Geological Survey of the Naples Bay (CARG Project) G. Aiello1 , F. Budillon1 , A. Conforti1 , B. D’Argenio1,2 , L. Ferraro1 , E. Marsella1 , N. Pelosi1 , M.L. Putignano1 , R. Tonielli1 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, Department of Earth Sciences, University of Napoli “Federico II”, Napoli, Italy gemma.aiello@iamc.cnr.it Abstract The results of the geological and geomorphological survey of the Naples Bay, carried out by the IAMC-CNR Institute of Naples on behalf of the National Geological Survey of Italy (now ISPRA) and by the Geological Survey of the Campania Region are here presented. The geological maps produced through the marine geological survey (at scales varying from 1:50.000 to 1:10.000) are: n. 446-447 “Napoli”, n. 464 “Ischia”, n. 465 “Procida”, n. 466 “Sorrento nord” and n. 484 “Capri”. The geological survey consisted in the acquisition of multibeam and single-beam bathymetric data, Sidescan Sonar acoustic imagery, reflection seismics, sea bottom samples and piston cores. A detailed marine geological survey, at the scale 1:10.000, has also been merged with underwater geological survey (up to the - 30 m of water depth), based on a direct sampling of both rocky outcrops and mobile sea bottoms, allowing to draw key geological and morpho-stratigraphic sections. The physiographic domains, representative of tectono-sedimentary and/or volcanic processes, include the continental shelf, the abrasional or depositional marine platforms, often in correspondence to coastal cliffs, the relic or preserved morphologies of monogenic volcanic edifices, the hummocky chaotic deposits, accumulated as a consequence of gravity instability processes. Moreover, the continental slope appears deeply incised by canyons evolving downslope to tributary channels. The CARG-MARE project has evidenced the regional geological structure of the Naples Bay, located along the western margin of the southern Apenninic chain and representing a Pleistocenic halfgraben, similarly to the peri-tyrrhenian basins, which form the seaward prolongation of the coastal plains of the Campania Region (Campania Plain and Volturno Plain).
1
Introduction
The CARG-MARE Project is among the most significant research activities carried out since the 1999 at the Institute for Marine Environment, National Research Council, Naples. The geologic understanding of the continental margins and the technological know-how of the Institute has
been concurrently growing in performing acquisition surveys, processing and geologic interpretation. This was needed by the marine geologic cartography of the Naples Bay (geologic maps n. 446-447 “Napoli-Ercolano”, n. 464 “Ischia”, n. 465 “Procida”, n. 466 “Sorrento nord”, n. 484 “Capri” and n. 485 “Termini”), a large project just completed that started in
Marine Geology
Figure 1: Digital Elevation Model of the Naples Bay recorded in the frame of the CARG project (reported from [1]) 2003 funded by the Geological Survey of Italy (ISPRA) and was enlarged by Region Campania (Settore Difesa Suolo). The geologic mapping has been carried out through a large amount of Multibeam and Single-Beam bathymetric data, Sidescan imagery, multichannel and single-channel (including Subbottom Chirp) profiles, sea bottom samples and cores, collected during several oceanographic cruises. Moreover, a large amount of data collected during Scuba surveys has been linked with geophysical data. Among the results of the projects it is worth to mention a high resolution Multibeam bathymetry of the Naples Bay ([1]; Figure 1), that has been used to construct a DEM (Digital Elevation Model) of the area. In this context, it is worth mentioning that the profiles used for the pho566
tomosaic construction of the sea bottom, cover the whole area. The latter, together with bathymetry, has been the base for the geological survey. The integrated geological interpretation of Multibeam bathymetry and Sidescan Sonar imagery, tied to sea bottom samples in correspondence to significant backscatter variations has been realised in a GIS environment. The morpho-structures and the seismic sequences, both volcanic and sedimentary, which characterize the stratigraphic architecture of the Naples Bay, have been also integrated with marine magnetic survey [2, 3, 4, 5, 6, 7]. Starting from the 2003 the IAMC-CNR of Naples has increased its activities through a Convention with the Campania Region (Soil Defence Sector) for a geolog-
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ical survey at the 1:10.000 scale of the whole coastal belt of the Campania Region (Naples and Salerno bays and Cilento Promontory) inside the – 200 m isobath. The marine geological survey at the scale 1:10.000 has been also merged with the underwater geological survey (down to the – 30 m isobath), with the direct sampling of the rocky substratum and subordinately, of the mobile sea bottom, for which lithologic and sedimentologic information have been collected. All the above data allowed to prepare geological and morpho-stratigraphic sections. The underwater geological survey allowed also to map the evidences of sea level stillstands (marine abrasion platforms, pocket beach deposits, often bioclastic, sandstones and beach rocks). Moreover the maps enclose the geometric and chronological relationships between underwater Quaternary landforms and deposits (with reference to the data available for the emerged sectors) and the morpho-stratigraphic elements, with reference to the recent evolution of the coastal belt.
2
Cartographic criteria and methodologies
The marine geological maps show the distribution of the main litho-stratigraphic units cropping out at the sea bottom, and of the main morphological and structural lineaments, according to the CARG standards expressed in the “Guidelines for the geological surveys of the marine areas” [8]. The main stratigraphic units recognized through the analysis of sediments cropping out at the sea bottom belong to the Late Quaternary Depositional Sequence. In this sequence the evolution and the migration
of the depositional environments may be recognised on the base of the variations of the depositional space during the last 4th order glacio-eustatic cycle, ranging between the 128 ky and the present-day (isotopic stage Q5e of [9]). The cartographic representation is chiefly based on the lithofacies associations, whose grouping form several depositional elements (portions of Systems Tract), and their individuation allows to restore the dynamics of the sedimentary environments. In particular, the Systems Tract of the Late Quaternary sequence, being limited by time-transgressive surfaces, are considered equivalent to the Unconformity-Bounded Units (Synthems – UBSU) and constitute the basic units of the cartographic representation. The non-linear geological features are represented by two superimposed levels: the textural classes distinguished according to Folk (1954), graphically represented with the full colour; the depositional elements distinguished with symbols. The superimposition of the environmental information on the textural one allows for a more complete reading of the cartography. The morphological elements are both areal and linear and represent another level of graphic superimposition on the geologic informations.
3
Marine sediments of the Late Quaternary Depositional Sequence
In the geological maps already published or in publication (like the map n. 465 “Procida”) the physiographic domains, mostly developed during the last Highstand Systems Tract, are characterized by spe567
Marine Geology
cific sedimentary processes and/or volcanic events. They include the littoral zone, the continental shelf (marine platform, abrasional or depositional, often at the base of coastal cliffs, relic or preserved morphologies of monogenic volcanic edifices, hummocky chaotic morphologies due to gravitational instability) and the continental slope (the latter deeply incised by canyons, evolving downslope in tributary channels). The littoral zone is composed of the toe of cliff deposits, submerged beach deposits, relic deposits, submarine slide deposits and beach rocks. The toe of cliff deposits are represented by heterometric blocks, often volcanic, having dimensions from metric to decametric, alternating with poorly sorted heterometric gravels, with heterogeneous texture. The submerged beach deposits are composed of gravels, gravelly sands and coarse-grained sands with pebbles (tuffs and lavas), as well as by coarse to medium grained sands, often with gravels and, subordinately, pebbles and isolated blocks (tufaceous and subordinately lavic in nature). The relic beach deposits are composed of well-rounded heterometric volcanic pebbles, immersed in a sandysilty matrix. Beach rock, sandstones and conglomerates have also been mapped. The inner shelf zone includes inner shelf deposits and bioclastic deposits. The inner shelf deposits are composed of gravels, sandy gravels and coarse grained lithobioclastic sands. The lithic component is made up of volcanic elements (tuff and other pyroclastic particles). The bioclastic component is formed by fragments of Molluscs and Echinoderms. The sandy fraction is represented by heterometric gravels, gravelly sands and bioclastic sands in a scarce pelitic matrix.
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The outer shelf environment includes clastic deposits and bioclastic deposits. The bioclastic deposits are composed of organogenic sands and gravels with scarce pelitic matrix, bio-detritic coarse-grained sands with scarce pelitic matrix (“coastal detritic”), draping on pelitic and sandypelitic sea bottoms. The muddy coastal detrital assemblage is formed by pelites with bioclastic gravels and sands. At places rhyzomes of marine Phanerogams have also been found. The slope zone includes the deposits of the Transgressive Systems Tract., of the Lowstand Systems Tract and Pleistocene relic marine deposits of the Falling Stage Systems Tract. Pelites with a prevailing sand fraction are diffused in correspondence of the canyons heads. The Transgressive Systems Tract is composed of volcanoclastic pebbles, alternating with thick pumice levels, often in a pelitic matrix. These deposits have a mounded external morphology, exhibit an acoustically-transparent seismic facies and crop out on the outer shelf, offshore the Naples town, in a bathymetric interval ranging between – 140 m and – 180 m. The Lowstand Systems Tract includes relic littoral deposits, composed of wellsorted sands and gravels with bioclastics, as well as middle to fine sands, organised as coastal prisms overlying the shelf break. The Pleistocene marine units are represented by coarse to fine grained marine deposits, forming relic morphologies of beach and continental shelf environments. Undifferentiated substratum, both carbonatic (mostly Mesozoic) in nature and volcanic, and forming the base of the Late Quaternary sequence has also been mapped.
Marine research at CNR
569 Figure 2: Regional geological sections constructed as layout elements at the scale 1:50.000 for the map n. 465 “Procida�.
Marine Geology
4
Tectonic setting
The CARG-MARE project has contributed to a detailed reconnaissance of the regional geological structure of the Naples and Salerno Bays and of the Cilento Promontory, that share most of the geologic and geomorphologic characteristics of the whole southern Tyrrhenian coastal belt. The marine geology in the surveyed area appears as complex as the onshore geology, even though they complement each other, being the latter characterised by older rocks exhumed by erosion, while the marine areas hold the history of the last few thousand of years. The Naples Bay is a Pleistocene halfgraben, comparable, from a geodynamic point of view, to the peri-tyrrhenian basins [10, 11, 12, 13]. This bay is part of a belt of coastal tectonic depressions: the Campania, Volturno and Sele Plains. The Campania Plain is bounded by marginal normal faults active along its borders and is filled by several thousand of meters of sed-
iments and volcanics, erupted by the Vesuvius, Phlegrean Fields and the Ischia and Procida islands. The Sorrento Peninsula represents a NESW trending horst, about 20 kilometers long, characterised by several listric faults, downthrowing northwards blocks of MesoCenozoic carbonates. The Naples Bay is crossed by two main systems of normal faults, with a NE-SW and NW-SE direction, along which the Nisida and Pentapalummo volcanic banks are located. Moreover, a complex fault system with a NNE-SSW trend offshore the Vesuvius has also been individuated [4]. The tectonic lineaments separating the north-western sector of the Bay from the south-eastern one, may be located along the Dohrn canyon, separating an area, where volcanic structures prevail, from another one, where sedimentary processes are dominant. According to several authors this lineament corresponds to a magmatic pathway, both onshore and offshore [3, 14].
References [1] B. D’Argenio, G. Aiello, G. De Alteriis, A. Milia, M. Sacchi, et al. Digital Elevation Model of the Naples Bay and adjacent areas, Eastern Tyrrhenian sea. Atlante di Cartografia Geologica scala 1:50:000 - progetto CARG - Servizio Geologico d’Italia (APAT), 32 Int. Cong. Firenze, 2004. [2] G. Aiello, F. Budillon, G. Cristofalo, B. D’Argenio, G. De Alteriis, M. De Lauro, L. Ferraro, E. Marsella, N. Pelosi, M. Sacchi, and R. Tonielli. Marine geology and morpho-bathymetry of the Naples Bay. pages 1–8, 2001. [3] G. Aiello, A. Angelino, E. Marsella, S. Ruggeri, and A. Siniscalchi. Carta magnetica di alta risoluzione del Golfo di Napoli (Tirreno meridionale). Bollettino della Societ`a Geologica Italiana, 123:333–342, 2004. [4] G. Aiello, A. Angelino, B. D’Argenio, E. Marsella, N. Pelosi, S. Ruggeri, and A. Siniscalchi. Buried volcanic structures in the Gulf of Naples (Southern Tyrrhenian sea, Italy) resulting from high resolution magnetic survey and seismic profiling. Annals of Geophysics, 48:1–15, 2005. 570
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[5] G. Aiello, E. Marsella, V. Di Fiore, and C. D’Isanto. Stratigraphic and structural styles of half-graben offshore basins in Southern Italy: multichannel seismic and Multibeam morpho-bathymetric evidences on the Salerno Valley (Southern Campania continental margin, Italy). Quaderni di Geofisica, 77:31, 2009. [6] G. Aiello, E. Marsella, and S. Passaro. Submarine instability processes on the continental slopes off the Campania Region (Southern Tyrrhenian sea, Italy): the case history of the Ischia island (Naples Bay). Bollettino di Geofisica Teorica Applicata, 50(2):193–207, 2009. [7] S. Ruggieri, G. Aiello, and E. Marsella. Integrated marine geophysical data interpretation of the Naples Bay continental slope (southern Tyrrhenian sea, Italy). Bollettino di Geofisica Teorica Applicata, 48(1):1–24, 2007. [8] Gruppo di Lavoro per la Geologia Marina del Servizio Geologico Nazionale. Nuove linee guida per il rilevamento geologico delle aree marine ricadenti nei fogli CARG. Linee Guida per la Cartografia Geologica Marina - APAT, 2002. [9] R. Catalano et al. Linee guida al rilevamento delle aree marine del Servizio Geologico Nazionale. Commissione per la Cartografia Geologica Marina ISPRA (exAPAT), 1996. [10] R. Bartole, D. Savelli, M. Tramontana, and F.C. Wezel. Structural and sedimentary features in the Tyrrhenian margin off Campania, southern Italy. Marine Geology, 55:163–180, 1983. [11] A. Cinque, P.P.C. Aucelli, L. Brancaccio, R. Mele, A. Milia, G. Robustelli, P. Romano, F. Russo, M. Russo, N. Santangelo, and I D. Sgambati. Volcanism, tectonics and recent geomorphological change in the Bay of Napoli. Supplemento Geogr. Fis. Dinam. Quat., 3:123–141, 1997. [12] F. Trincardi and N. Zitellini. The rifting of the Tyrrhenian basin. Geomarine Letters, 7(1):1–6, 1987. [13] M. Mariani and R. Prato. I bacini neogenici costieri del margine tirrenico: approccio sismico-stratigrafico. Memorie della Societ`a Geologica Italiana, 41:519–531, 1988. [14] C. Scarpati, P. Cole, and A. Perrotta. The Neapolitain Yellow Tuff. A large volume multiphase eruption from Campi Flegrei, Southern Italy. Bulletin of Volcanology, 55:343–356, 1993.
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The Regional Geological Structure of the Naples Bay Inferred by New Multichannel Seismic Reflection Profiles G. Aiello1 , A.G. Cicchella2 , V. Di Fiore1 , E. Marsella1 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, University of Napoli “Federico II”, Napoli, Italy gemma.aiello@iamc.cnr.it Abstract The deep regional geological structure of the Naples Bay is controlled by two NE-SW and NNE-SSW (Anti-Apenninic) normal faults, the Dohrn canyon fault and the Capri-Sorrento fault. Strong downthrows of Meso-Cenozoic sedimentary sequences, representing the acoustic basement, have been observed in correspondence to both the faults. Main regional morpho-structures are: the Banco di Fuori , a morpho-structural high of the Meso-Cenozoic carbonatic acoustic basement, bounding southwards the Naples Bay, whose flanks and top are draped by the Pleistocene deposits of the Late Quaternary sequence; the Dohrn canyon, separating the eastern side of the Bay, where sedimentary seismic sequences crop out, from the western one, where volcanic seismic units prevail; the Magnaghi canyon, draining the volcanic and volcano-clastic input of the Ischia and Procida islands during the Late Quaternary and eroding the sediments of the Mg volcanic unit, characterized by chaotic reflectors; the Capri Basin, a deep basin localised in the Tyrrhenian bathyal plain south of the Naples Bay in the southern sector of the Dohrn canyon, infilled by Pleistocenic-Holocenic sediments overlying the carbonatic basement; the Salerno Valley, a half-graben filled by three seismic units corresponding to Quaternary marine and continental deposits, laterally grading to chaotic sequences related to the “Flysch del Cilento” Auct; the Volturno Basin, infilled by four marine to deltaic seismic sequences, frequently alternating with volcanoclastic levels, overlying deep seismic units, correlated with Miocene flysch deposits (sands and shales) and MesoCenozoic carbonates.
1
Introduction
Some new data on the deep geological structure of the Naples Bay based on regional multichannel seismic profiles parallel and perpendicular to the continental margin off Campania are here presented. Similarly to other back-arc basins, the Tyrrhenian sea is an area of ongoing extension inside large-scale convergence be-
tween the continental plates of Europe and Africa. Tyrrhenian extension started about 10 Myr ago, leading to the Pliocene formation of oceanic crust [1, 2, 3, 4]. Three passive continental margins, Sardinia, Northern Sicily and Southern Italy border the Southern Tyrrhenian bathyal plain. This area is seismically active and experienced strong horizontal and vertical movements. The Southern Tyrrhenian continental mar-
Marine Geology 14°00’
14°30’
15°00’
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r Ty n ia en rh se
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a
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rn
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Vesuvius ISCHIA
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Shallow water carbonatic successions and basinal units (Mesozoic)
Pyroclastic deposits and lavas (Quaternary)
Siliciclastic synorogenic units of piggy back basins (Miocene)
Marine and continental deposits (Quaternary)
Normal faults
Volcanic calderas
20 km
Figure 1: Sketch location map of the three processed and interpreted seismic profiles on the geological map of the western margin of the Apenninic chain. Main regional faults have been also reported. Key: Shallow water carbonatic successions and basinal units (Mesozoic). Siliciclastic synorogenic units of piggy back basins (Miocene). Pyroclastic deposits and lavas (Quaternary). Marine and continental deposits (Quaternary). Normal faults. Volcanic calderas. gin off Campania owes its complex stratigraphic architecture to the interaction between volcanic and sedimentary processes during the Late Quaternary. Multichannel reflection seismics collected by the CNR-IAMC Institute of Naples, Italy during several oceanographic cruises, starting from the 1996, coupled to high resolution reflection seismics, Multibeam bathymetry and magnetic profiles have contributed to the knowledge of the tectono-stratigraphic setting of the margin, both to a regional and to a local scale [5, 6, 7, 8, 9, 10, 11, 12, 13, 14, 15, 16].
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Basin formation processes controlled by the interplay of extension and compression and their tectonic expression on the basin fill have been intensively studied in the frame of the International Lithosphere Task Force “Origin of Sedimentary Basins” [17]. Another intriguing research theme is represented by the relationships between deeper lithospheric processes and near-surface tectonics of sedimentary basins [18]. Understanding the development of sedimentary basins on continental margins requires insight into the processes responsible both for basin subsi-
Marine research at CNR
dence and for distribution and preservation of sediments within the basin. It is widely accepted that passive margin basins subside as a result of lithospheric extension, assuming pure or simple shear extension, uniform or depth-dependent [19]. According to sedimentologic and geomorphologic data, during the post-rift evolution of extensional basins the surface processes (erosion and sedimentation) modify the topography and thickness of sedimentary fill at rates comparable with the rates of tectonic uplift/subsidence [20]. A coupling between the surface and the tectonic processes can be expected. Important results on the topographic evolution of the European continental margins and related basins have been recently collected in the frame of the Topo-Europe Research Programme, initiating a number of novel studies on the quantification of rates of vertical motions, related tectonically-controlled river evolution and land subsidence in carefully selected natural laboratories in Europe [21]. A grid of multichannel seismic lines was acquired in the 1999 on the continental margin of Campania by a joint project between Geomare Sud (Naples, Italy), the Marine Geology Group of the University of Palermo and the Department of Tectonics of the Vrije University of Amsterdam., the Netherlands (SISTER99 oceanographic cruise [22, 23]). The main goal of the cruise (Seismic Investigations in the South Tyrrhenian Extensional Regions) was to acquire data to constrain the kinematics of formation of the Tyrrhenian margin. The projects aims at a quantification of vertical and horizontal movements experienced by the continental margin off Campania through the processing, the interpretation and the depth-conversion of seismic lines. Images derived from these seismic
lines provide a kinematic frame for the development of the Tyrrhenian system in last 10 MY, allowing for the linking between basin-scale and crustal-scale phenomena/processes. During the cruise regional seismic lines parallel to and across the continental margin of Campania from Gaeta to Capo Palinuro have been acquired. Multichannel profiles clearly document the geometries of normal faults running parallel to the margin. Two major faults defining fault blocks 10-20 km wide and dipping towards the ocean are defined by the interpretation of seismic profiles [22]. The profile parallel to the continental margin clearly depict the main extensional regions of the Campania offshore, as the Salerno graben and the Sapri basin to the south. Shortening related tectonic features also appear, testifying inversion processes at a basin scale. In this paper we present three processed profiles (SISTER4 2, SISTER9 1 and SISTER7 2, Figure 1) localised in the Naples Bay. Two of them finish in correspondence to the Campania-Latium Tyrrhenian margin (Volturno Basin), the third one towards the sedimentary basins of the Salerno Valley, localised southwards of the Sorrento Peninsula [9].
2
Geological setting
The Naples Bay lies in the southern part of a tectonic depression, the Campania Plain, produced from the back-arc extension that accompanied the NE-verging accretion of the Apenninic thrust belt during the roll-back of the subducting foreland plate. The period of wedge accretion goes from Middle-Late Miocene to the end of the Early Pleistocene. In the Campania Plain, the first stage of lower575
Marine Geology
ing and submersion (Early Pleistocene in age) was probably controlled by a NW-SE extension like the following one (Middle Pleistocene). Normal faults, inherited from both these stages were then reactivated during Late Pleistocene and Holocene phases of subsidence and uplift, especially in the zones affected by volcano-tectonics [24]. The Phlegrean Fields are a volcanic district surrounding the western part of the Naples Bay, where volcanism has been active for at least 50 ky [25]. Its present morphology refers to events which occurred after the emplacement of the Campanian Ignimbrite, a huge pyroclastic flow erupted 35 ky ago, when the area experienced a first phase of calderization. The volcanic districts of Procida, Vivara and Ischia islands represent the seaward prolongation of the Phlegrean Fields area, that also includes a number of submerged volcanic banks (Miseno, Nisida and Pentapalummo). The Somma-Vesuvius volcanic activity and the sedimentary processes in the Sarno-Sebeto coastal plain controlled the development of erosional and volcano-sedimentary processes in the eastern sector of the Naples Bay. Previous geophysical studies on the deep structure of the Bay refer, in particular, to the carbonatic acoustic basement, [26, 27], underlying the Pleistocene sedimentary sequences of the basin fill. The carbonatic acoustic basement dips towards NW with angles of about 10° and is characterised by extensional tectonics. Its extensional tectonic syle, related to the opening of the Tyrrhenian sea, controlled the development of a half-graben structure along the western Apenninic margin. Several regional faults have been singled out, i.e.a N110° trending normal fault, responsible for the downthrowing of the carbonatic basement under the Vesuvius volcano, a N10° trending nor576
mal fault, along which five submarine volcanoes are aligned and a N70° normal fault, separating the Naples Bay from the Salerno Bay. The first two faults are arranged radially with respect to the magma chamber of the Phlegrean Fields and are interpreted as the main cause of magma uprising. The carbonatic basement has not been recognised under the Phlegrean Fields volcanic complex. A major morpho-structural high (Banco di Bocca Grande or Banco di Fuori) separates the Dohrn from the Magnaghi canyon and is presumably formed by a Mesozoic carbonate block, that resulted from the regional uplift and tilting of the carbonatic acoustic basement, cropping out in the Sorrento Peninsula-Capri island structural high. The carbonatic nature of the Banco di Fuori high is suggested by its location along the structural alignment Capri island – Sorrento Peninsula and confirmed by the lack of significant magnetic anomalies [28, 8].
3
Seismic data processing
The employed instruments and the advanced techniques of multichannel seismic data acquisition have allowed to obtain high quality data also in the Campania volcanic area, where the occurrence of pyroclastic levels and buried volcanic bodies produce a strong scattering of the acoustic energy. In particular, new Airguns, a new 48-channel streamer and a new system of data acquisition (Geometrics), in dotation to the IAMC-CNR of Naples, Italy have been used. The acquisition parameters are represented by the seismic source, by the length of the seismogram, by the sample interval, by the distance between sources and by the dis-
Marine research at CNR
tance between hydrophones. The seismic source consisted of two GI guns SI/Sodera (210 cubic inch); the length of the seismogram was of 5 sec; the sample interval was of 1 msec; the distance between sources was of 25 m; finally, the distance between hydrophones was of 12.5 m. The initial procedure of seismic data processing consisted of data quality check and field geometry assignment. Trace editing aimed to detection and removal of dead or very noise traces and spikes, that may induce problems with forward Fast Fourier Transform (FFT). A top muting allowed to eliminate the signal above the first arrivals of the seismic traces. Automatic Gain Control (AGC) allowed a trace normalization. Seismic data processing was aimed to reduce random noise in the data and to improve the resolution of the seismic wavelet with spiking deconvolution algorithms. The velocity analysis was carried out to remove the moveout on the CDP gathers in order to define the velocity of the different reflectors and to obtain the final stacked section. A post-stack deconvolution has been applied to the data to remove multiple arrivals. After deconvolution we also applied a single bandpass filter in order to improve the interested signal along the sections. The seismic interpretation was carried out taking into account the occurrence of multiples and noises and carefully eliminating them from geologic interpretation. The technique utilised for the multiple removal, mainly the sea bottom multiples has been defined after several tests; the velocity analysis and the predictive deconvolution strongly improved the resolution of the seismic stacked sections, on which the geologic interpretation has been carried out.
4
Regional geological interpretation
The seismic interpretation has revealed that the deep structure of the Naples Bay is controlled by two normal faults, NE-SW and NNE-SSW (Counter-Apenninic), i.e. the Dohrn canyon fault and the Capri-Sorrento fault (Figure 2, see also [9]). Strong downthrows of the Meso-Cenozoic carbonatic sequences, representing the acoustic basement and extensively cropping out in correspondence to the structural high CapriSorrento Peninsula have been observed in correspondence to these faults. The Dohrn canyon fault bounds the south-western flank of the Banco di Fuori and downthrows the Meso-Cenozoic carbonates under the canyon, with a vertical throw in the order of one thousand of metres. The syntectonic nature of the deposits of the prograding wedge of the Middle Late Pleistocene (unit B) in correspondence to this fault allows to date its activity, Pleistocenic in age. Moreover, the fault seems to be suturated by the Holocene deposits, which appear to be relatively undeformed. Main regional morpho-structures evidenced by the seismic interpretation are: 1. the Banco di Fuori, a morpho-structural high of the Meso-Cenozoic carbonatic acoustic basement, bounding the southern sector of the Naples Bay, whose flanks and top are draped by the Pleistocene deposits of the Late Quaternary depositional sequence (Figures 2 and 3); 2. the Dohrn canyon, a main canyon crossing the Naples Bay and separating the eastern sector of the Bay, where the sedimentary seismic units crop out, from the western one, where the volcanic seismic units prevail. It is articulated in 577
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two branches, the eastern one and the lying the carbonatic acoustic basement; western one, merging in a thalweg hav- 5. the Salerno Valley (Figure 2), a halfing an Anti-Apenninic (NE-SW) direcgraben basin characterised by three tion, bounded southwards by the Capri main seismic units corresponding to Basin (Figures 2 and 3); Quaternary marine and continental sed3. the Magnaghi canyon, draining the voliments, laterally grading to chaotic canic and volcanoclastic input coming deposits related to the “Flysch del from the eruptive activity of the IsCilento� Auct; chia and Procida islands during the Late 6. the Volturno basin, hosting the northQuaternary and carving the sediments ern sector of the Campania Plain and of the volcanic Mg unit, characterized the surrounding offshore, whose sediby reflectors having a chaotic distribumentary filling consists of four marine tion (Figure 2); and deltaic seismic sequences, alternat4. the Capri basin (Figure 3), a deep basin ing with volcanoclastic levels, overlying localised in the Tyrrhenian bathyal plain the deep seismic units of the acoustic southwards of the Dohrn canyon, filled basement, related with Miocene flysch by a thickness of 0.7 sec (twt) of deposits (sands and shales) and MesoPleistocene-Holocene sediments overCenozoic carbonates.
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Marine research at CNR
Line SISTER9_1 SISTER 7_2 CDP
twt (sec)
NW
SE
0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 1.1 1.2 1.3 1.4 1.5 1.6 1.7 1.8 1.9 2.0 2.1 2.2 2.3 2.4 2.5 2.6 2.7 2.8 2.9 3.0
SISTER 7_2 Dohrn fault
NW Banco di Fuori Magnaghi Basin
SE
Capri-Sorrento fault
Capri structural high Dohrn canyon
B
MC
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Salerno Valley
A Mg
Hol A Pls
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FC
MC
Figure 2: Multichannel seismic profile SISTER9 1 and corresponding geologic interpretation (see the Figure 1 for the location). The main morpho-structural features are the Magnaghi basin, the Banco di Fuori, the Dohrn canyon, the Capri structural high and the Salerno Valley. Note the occurrence of the regional master faults, i.e. the Dohrn fault, bounding the south-eastern flank of the Banco di Fuori high and the Capri-Sorrento fault, bounding the south-eastern flank of the Capri structural high. Key: MC: Acoustic basement. Meso-Cenozoic carbonates cropping out offshore the Sorrento Peninsula and the Capri island. FC: Acoustic basement. Cenozoic siliciclastic deposits related to the �Flysch del Cilento� Auct., underlying the sedimentary filling of the Salerno Valley. A: Early Pleistocene relic prograding wedge representing the lower unit in the stratigraphic architecture of the Naples Bay, characterised by obliquous prograding clinoforms B: Late Pleistocene prograding wedge representing the upper unit in the stratigraphic architecture of the Naples Bay characterised by low angle sigmoidal to obliquous clinoforms, supplied by the palaeo-Sarno river mouth. Mg: Late Pleistocene volcanic seismic unit characterised by an acoustically transparent seismic facies and constituting the stratigraphic architecture of the Ischia offshore under the Magnaghi canyon. Pls: Late Pleistocene seismic unit representing the upper unit of the basin filling of the Salerno Valley composed of marine sediments. Hol: Holocene highstand drape. 579
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LINE SISTER 7_2
SISTER 9_1
SISTER 4_1
twt (sec)
SW
NE
0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 1.1 1.2 1.3 1.4 1.5 1.6 1.7 1.8 1.9 2.0 2.1 2.2 2.3 2.4 2.5 2.6 2.7 2.8 2.9 3.0
DOHRN CANYON’S THALWEG
CAPRI BASIN
BANCO DI FUORI
OUTER SHELF OF THE NAPLES BAY
Hol Hol
Hol1 Pls
F2
Pls
Mc
Hol2
F1
B
Hol1 A Hol Pls
?
?
?
Mc
Figure 3: Multichannel seismic profile SISTER7 2 and corresponding geologic interpretation (see the Figure 1 for the location). The main morpho-structural lineaments identified on the seismic profile are: the Capri Basin, the Dohrn canyon, the Banco di Fuori and the outer shelf of the Naples Bay. Key: MC. Acoustic basement. Meso-Cenozoic carbonates cropping out offshore the Sorrento Peninsula and the Capri island. A: Early Pleistocene relic prograding wedge representing the lower unit in the stratigraphic architecture of the Naples Bay, characterized by obliquous prograding clinoforms. B: Late Pleistocene prograding wedge representing the upper unit in the stratigraphic architecture of the Naples Bay, characterized by low angle sigmoidal to obliquous clinoforms, supplied by the palaeo-Sarno river mouth. Pls: (a) Late Pleistocene seismic unit corresponding to the Forced Regression System Tract and the Transgressive System Tract of the Late Quaternary depositional sequence overlying the top and the flanks of the Banco di Fuori structural high. (b) Late Pleistocene seismic sequence representing the basin filling of the Capri basin. Hol: Holocene deposits, characterised by parallel and continuous seismic reflectors in the Capri basin; (1) Holocene deposits, characterised by an acoustically transparent seismic facies in correspondence to channel levee complexes in the Dohrn canyon’s thalweg; (2) Holocene deposits on the outer shelf of the Naples Bay, characterized by alternances of marine and volcanoclastic sediments; (F) slide deposits.
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References [1] E. Patacca, R. Sartori, and P. Scandone. Tyrrhenian basin and Apenninic arcs: kinematic relations since late Tortonian times. Memorie della Societ`a Geologica Italiana, 45:425–451, 1990. [2] E. Patacca and P. Scandone. Geology of the Southern Apennines. Bollettino della Societ`a Geologica Italiana, 7:75–119, 2007. [3] R. Sartori. The Tyrrhenian back-arc basin and subduction of the Ionian lithosphere. Episodes, 26(3), 2004. [4] R. Sartori, L. Torelli, N. Zitellini, G. Carrara, M. Magaldi, and P. Mussoni. Crustal features along a W-E Tyrrhenian transect from Sardinia to Campania margins (Central Mediterranean). Tectonophysics, 383:171–192, 2004. [5] G. Aiello, F. Budillon, G. de Alteriis, O. Di Razza, M. De Lauro, E. Marsella, N. Pelosi, F. Pepe, M. Sacchi, and R. Tonielli. Seismic exploration of the perityrrhenian basins in the Latium-Campania offshore. International Congress 8th Workshop Task Force “Origin of Sedimentary Basins”. June 1997, Palermo, Italy, Extended Abstract, 1997. [6] G. Aiello, F. Budillon, G. de Alteriis, E. Marsella, G. Pappone, and M. Sacchi. Late Neogene tectonics and basin evolution along the Tyrrhenian margin of Southern Italy. International Congress 8th Workshop Task Force “Origin of Sedimentary Basins”, June 1997, Palermo, Italy, Extended Abstract, 1997. [7] G. Aiello, F. Budillon, G. Cristofalo, B. D’Argenio, G. De Alteriis, M. De Lauro, L. Ferraro, E. Marsella, N. Pelosi, M. Sacchi, and R. Tonielli. Marine geology and morphobathymetry of the Naples Bay. In structures and Processes of the Mediterranean Ecosystems, Chapter 1, pages 1–8, 2001. [8] G. Aiello, A. Angelino, B. D’Argenio, E. Marsella, N. Pelosi, S. Ruggieri, and A. Siniscalchi. Buried volcanic structures in the Gulf of Naples (Southern Tyrrhenian sea, Italy) resulting from high resolution magnetic survey and seismic profiling. Annals of Geophysics, 48:1–15, 2005. [9] G. Aiello, V. Di Fiore, E. Marsella, and C. D’Isanto. Stratigraphic and structural styles of half-graben offshore basins in Southern Italy: multichannel seismic and Multibeam morpho-bathymetric evidences on the Salerno Valley (Southern Campania continental margin, Italy). Quaderni di Geofisica, 77:1–33, 2009. [10] G. Aiello, E. Marsella, and S. Passaro. Submarine instability processes on the continental slopes off the Campania region (Southern Tyrrhenian sea, Italy): the case history of the Ischia island (Naples Bay). Bollettino di Geofisica Teorica Applicata, 50(2):193–207, 2009.
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[11] E. Marsella, G. Aiello, A. Angelino, P.P. Bruno, V. Di Fiore, F. Giordano, N. Pelosi, A. Siniscalchi, C. D’Isanto, and S. Ruggieri. Shallow geological structures and magnetic anomalies in the Gulf of Naples. An integrated analysis of seismic and magnetometric profiles. Bollettino di Geofisica Teorica Applicata, 52(1-2):292– 297, 2002. [12] A. Siniscalchi, A. Angelino, S. Ruggieri, G. Aiello, E. Marsella, and M. Sacchi. High resolution magnetic anomaly map of the Bay of Naples (Southern Tyrrhenian sea, Italy). Bollettino di Geofisica Teorica Applicata, 42:99–104, 2002. [13] M. Secomandi, V. Paoletti, G. Aiello, M. Fedi, E. Marsella, S. Ruggieri, B. D’Argenio, and A. Rapolla. Analysis of the magnetic anomaly field of the volcanic district of the Bay of Naples, Italy. Marine Geophysical Researches, 24:207– 221, 2003. [14] S. Ruggieri. Applicazione di metodologie di geofisica marina allo studio dei margini continentali in aree vulcaniche attive: carta magnetica di alta risoluzione del margine continentale campano-laziale tra i Golfi di Gaeta e di Napoli (Tirreno centro-meridionale). Tesi di Dottorato di Ricerca in Scienze ed Ingegneria del Mare, Universit`a degli Studi di Napoli “Federico II”, 2006. [15] S. Ruggieri, G. Aiello, and E. Marsella. Integrated marine geophysical data interpretation of the Naples Bay continental slope. Bollettino di Geofisica Teorica Applicata, 48:1–24, 2007. [16] A.G. Cicchella. Analisi ed elaborazione di profili sismici a riflessione nel Golfo di Napoli (Tirreno meridionale). Tesi di Dottorato di Ricerca in Scienze ed Ingegneria del Mare, Universit`a degli Studi di Napoli “Federico II”, 2009. [17] S. Cloetingh, B. D’Argenio, R. Catalano, S. Horvath, and W. Sassi. Interplay of extension and compression in basin formation: an introduction. Tectonophysics, 252:1–5, 1995. [18] G. Quinlan, J. Walsh, J. Skogseid, W. Sassi, S. Cloetingh, L. Lobkovsky, C. Bois, H. Stel, and E. Banda. Relationship between deeper lithospheric processes and near-surface tectonics of sedimentary basins. Tectonophysics, 226:217–225, 1993. [19] P.A. Allen and J.A Allen. Basin analysis. Principles and applications. Blackwell Scientific Publications, 1990. [20] E. Burov and S. Cloetingh. Erosion and rift dynamics: new thermomechanical aspects of post-rift evolution of extensional basins. Earth and Planetary Science Letters, 150:7–26, 1997. [21] S. Cloetingh, P.A. Ziegler, P.J.F. Bogaard, P.A.M. Andriessen, I.M. Artemieva, G. Bada, R.T. Van Balen, F. Beekman, Z. Ben Avraham, J.P. Brun, H.P. Bunge,
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E. Burov, E. Carbonell, C. Faccenna, A. Friedrich, J. Gallart, A.G. Green, O. Heidbach, and A.G. Jones. TOPO-EUROPE: the geoscience of coupled deep-earth surface processes. Global and Planetary Change, 58:1–118, 2007. [22] G. Bertotti, E. Marsella, N. Pelosi, F. Pepe, R. Tonielli, and the Sister 99 Shipboard Scientific Party. Sister99: a seismic campaign to investigate the kinematics of the South Tyrrhenian extensional regions. Giornale di Geologia, 61(3):25–36, 1999. [23] A. Korevaar, A. Pagano, V. Vandervejer, G. Bertotti, E. Marsella, and F. Pepe. Regional seismic lines across and along the Campania passive continental margin. EUG General Assembly, Abstract, 2000. [24] A. Cinque, P.P.C. Aucelli, L. Brancaccio, L. Mele, A. Milia, G. Robustelli, P. Romano, F. Russo, N. Santangelo, and D. Sgambati. Volcanism, tectonics and recent geomorphological change in the Bay of Napoli. Suppl. Geogr. Fis. Dinam. Quat., 3:123–141, 1997. [25] M. Rosi and A. Sbrana. Phlegrean Fields. Quaderni De La Ricerca Scientifica, CNR, Italy, 1987. [26] N. Fusi. Structural setting of the carbonatic basement and its relationships with magma uprising in the Gulf of Naples (Southern Italy). Annali di Geofisica, 39(3), 1996. [27] G. Berrino, G. Corrado, and U. Ricciardi. Sea gravity in the Gulf of Naples: a contribution to delineating the structural pattern of the Vesuvian area. Journal of Volcanology and Geothermal Research, 82:139–150, 1998. [28] G. Aiello, A. Angelino, E. Marsella, S. Ruggieri, and A. Siniscalchi. Carta magnetica di alta risoluzione del Golfo di Napoli (Tirreno meridionale). Bollettino della Societ`a Geologica Italiana, 123:333–342, 2004.
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Ecosystem Variability in the Strait of Sicily: Evidence from Major and Trace Elements in Sediment Box-Cores G. Tranchida1 , A. Bellanca2 , M. Angelone3 , A. Bonanno1 , C. Buscaino1 , L. Langone4 , R. Neri2 , B. Patti1 1, Institute for Coastal Marine Environment, CNR, Capo Granitola (TP), Italy 2, Faculty of Mathematical, Physical and Natural Sciences, University of Palermo, Palermo, Italia 3, Technical Unit for Environmental Characterization, Prevention and Recovery, ENEA Roma, Italy 4, Institute of Marine Sciences, CNR, Bologna, Italy giorgio.tranchida@iamc.cnr.it Abstract Results obtained from the geochemical high-resolution study carried out on boxcore sediments from the Strait of Sicily permitted to reconstruct the spatial and temporal distribution of major and trace elements. Sediments were collected along two onshore-offshore transects in front of the towns of Sciacca and Pozzallo. Most samples generally display Ti/Al and K/Al ratio values around 0.06 and 0.02, respectively, that are consistent with riverine and aeolian detrital sources from surrounding land. Deviation from these values might reflect high energy depositional environments. Geochemical spatial variability along the Sciacca- Pantelleria transect reflects the irregular morphology of the Adventure Bank. Based on 210 Pb chronology, the vertical distribution of heavy metals in the cores well records the effects of industrialization, agricultural activities, and urbanization that affected the southern coast of Sicily since the begin of the 20th 22 century with an evident enhancement starting around the 1960s. Some variations of major and trace element contents are interpreted as effect of geothermal and/or magmatic activities in the Strait of Sicily. The dynamic and the chemical-physic features of the water mass, combined with the bottom morphology, are invoked to explain the higher sediment accumulation rate and increasingly high contents of some trace metal measured in the sediments near the Pozzallo coast.
1
Introduction
Despite of its reduced dimensions, the Strait of Sicily plays an important role in the physical and dynamical processes evolving in the whole Mediterranean [1, 2]. According to various authors [3, 4, 2], the sill topography in the Sicily Strait has a
principal role in the water mass circulation. In fact, the Atlantic Ionian Stream (A.I.S.; [5]) encircles two cyclonic vortices over the Adventure Bank and off Cape Passero, and describes a pronounced anticyclonic meander in between the Maltese Channel. In this context, the influence of the anthropogenic impact in the
Marine Geology
Figure 1: Bathymetric map of the investigated area with sampling sites. area needs to be considered. In fact, the southern belt of Sicily, but also other parts of the island, since 1950 suffered the consequences of an intensive building speculation along the coasts; the widely, often wild, hydraulic arrangement of most rivers; the intensive agricultural utilization, especially in the southeastern belt of Sicily; and the activity of an important industrial centre in the town of Gela producing mineral oil since 1960. The main objectives of this study were: (i) to reconstruct the spatial and temporal geochemical variability of recent sediments in the Sicily Channel; (ii) to assess the detrital sources and sediment accumulation rates; (iii) to define the distribution of heavy metals, distinguishing between anthropogenic and natural inputs; (iv) to reconstruct the anthropogenic impact history in the Strait of Sicily during the last 150-200 years.
586
2
Sampling
Sampling sites were arranged along two transects (Sciacca-Pantelleria and Pozzallo-Malta) perpendicularly to the southern coast of Sicily (Figure 1). Sediments were collected by box-corer sampler. Cores were sealed, labelled and stored at -20 째C until the analysis. In the lab, defrosted cores were extruded and cut at 1-2 cm intervals with a stainless steel saw and sediment slices dried at 50 째C. A total of 191 samples were obtained.
3
Analytical methods
In order to characterize the geochemistry of the two transects, analyses of major and trace elements were performed on sediment by X-ray fluorescence spectrometry (XRF). Data reduction was achieved fol-
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lowing the method described by Franzini et al. [6]. Bulk sediment mineralogy was determined by powder X-ray diffraction (XRD). The relative proportions of minerals were determined according to methods and data of Schultz [7] and [8]. The CaCO3 content was determined by means of a classic gas-volumetric technique reported in Husselman [9]. To determine chronology and mass accumulation rate of the studied sediments, radiometric analysis of 210 Pb activity was carried out on selected cores. The age of sediments was determined by using the constant rate of supply (CRS) model [10, 11, 12, 13]. According to this model, the flow of 21068 Pb from the column water towards the seafloor is considered constant, independently from the sedimentation rate [14, 15]. In order to determine pseudo total heavy metals, samples were digested with aqua regia in Teflon bombs using a microwave mineralizer. This method is widely used in environmental geochemistry studies [16, 17, 18, 19]. The obtained solutions were analysed by inductively coupled plasma mass spectrometry (ICP-MS).
4 4.1
Results and discussion Sediment mineralogy
For the Sciacca-Pantelleria transect, XRD analysis highlights some differences among each site. In the sites 164, 272, 440, the silicate phases are dominant with more abundant clay minerals (averagely 50% in site 440 and up to 70% in site 272). In site 440, a moderate feldspar content (average value 9%) is present. The sediments of onshore site 64 and offshore site 611 are characterized by more abundant carbonates
with average values close to 60% for site 64 and to 80% for site 611, where Mg-calcite (mean value 23%) occurs. In addition to calcite, cores 64 and 611 display the presence of aragonite with mean values of 24% and 29%, respectively. The main characteristic of the Pozzallo-Malta transect is the decreasing trend of clay minerals (from 45% to 86 38%) and the increasing trend of carbonate phases from onshore to offshore sediments.
4.2
Major element geochemistry
Results highlight two different sample populations corresponding to onshore and offshore box-cores. The onshore sediments generally exhibit Al2O3 and Fe2O3 contents lower than those offshore. However, some exceptions are sediments from the more distal stations that yield higher CaO and MgO values, increasing from the eastern to western sites.
4.3
Element/Al ratios
According to various authors [20, 21, 22], element data were normalized to Al to compensate for carbonate dilution and changeable grain-size. Consistently with the XRD results, El/Al ratios varying within roughly comparable intervals indicate a general vertical and spatial chemical homogeneity within and among all transect cores. In most samples, K/Al ratios fluctuate around a mean value of 0.2 and Ti/Al ratios mostly show values of about 0.06, close to the values measured for Late Quaternary Mediterranean sediments [23, 24, 25] and inferred to be representative of detrital export from surrounding land to the Strait of Sicily. Occurrence of high K/Al 587
Marine Geology
ratios in a sediment are commonly associated to an illite-rich, smectite-poor clay fraction. Thus, a reason for anomalously high K/Al values (from 0.27 to 0.37) in cores 64 and 611 could be the fact that, at these locations, sediments are deposited under a high-energy regime favouring resuspension of fine, smectite-rich particles. This is supported by the presence of coarse biogenic detritus (core 64 and 611). The clear positive shift in the lower portion of Ti/Al profile for the site 440 in correspondence of a negative excursion in the K/Al curve may be ascribed to magmatic activity implying emission of volcanic materials. This is consistent with Hawaiites and Ol-basalts from Pantelleria showing Ti/Al and K/Al ratios of 0.12-0.27 and 0.10.12, respectively [26].
4.4
Dating and sedimentation rate
Measurements of 210 Pb radiometric activity of four box-cores (two for each transect) were performed in order to calculate the sediment accumulation rates in the studied area. 210115 Pb analysis results were extrapolated to the rest of the studied sediments on the basis of matching depth multimetal profiles. For the sites 272 and 440 of the SciaccaPantelleria transect, similar average accumulation rates of 0.17 and 0.16 cm·yr−1 118 were respectively calculated, highlighting, probably, a quite similar sediment source, coherently with the mineralogical and chemical data. For the Pozzallo-Malta transect, the average accumulation rate is quite high, with 0.40 cm·yr−1 120 , at site 134 and as low as 0.05 cm·yr−1 at site 407. The markedly different sedimentation rate val588
ues between nearshore and offshore sediments of the Pozzallo-Malta transects are probably due to flow currents influenced by environmental conditions (depth, sea-floor morphology) and to a high detrital contribution from southeastern Sicily. The higher sediment accumulation rate recorded at site 134 (Pozzallo-Malta) is probably linked principally to the water mass circulation in the Strait of Sicily. As previously observed, the A.I.S. describes a cyclonic vortex in the Capo Passero area [2] with a consequent slowing down of the stream that determines a decrease of the sediment load and also an increase of the accumulation rate. Successively, the AIS stream with a cyclonic encirclement departs from the Capo Passero area seawards with a reduced sediment load capacity, which probably explains the low sedimentation rate (0.05 cm·yr−1 ) measured at site 407.
4.5
Sub-total heavy metal contents
Concentrations of Co, Cd, Cr, Cu, Ni, Pb, V, Mn, Zn, Sb, and As were measured in sediment box-cores recovered along the Sciacca-Pantelleria and Pozzallo-Malta transects. Arsenic contents were measured only in the box-core 440 of the SciaccaPantelleria transect and in all box-cores of the Pozzallo-Malta transect. In the Sciacca-Pantelleria transect, sediments of sites 64 and 611 display low contents of most heavy metals accounting for dilution effect of an abundant carbonate component. Throughout the cores, metal concentrations are moderately to widely fluctuating, with higher values in the upper portion at sites 272 and 440. The sediments of the Pozzallo-Malta transect are characterized
Marine research at CNR
by a lesser variability of the trace metal contents, with generally higher values at nearshore site 134 and lower at site 188. Interestingly, the Mn profile for the box-core 521 exhibits a clear abrupt increase (at 12.5 cm) from 750 to 2420 mg¡kg−1 , maintaining persistently high values in the uppermost sediments. For the same site, other trace metals describe opposite patterns. After removing the carbonate dilution effect by normalization to Al, trace metal concentrations appear quite similar for different sites of the Sciacca-Pantelleria transect and vertical profiles more clearly illustrate an upward increasing trend for most elements. Al normalization poorly affects elemental distribution patterns for the Pozzallo-Malta cores.
More significant results are represented in Figures 2 and 3. EFs are generally less than 1 for Cd, Cr, Cu, Co, V, Zn, and Ni. In the Sciacca-Pantelleria and Pozzallo-Malta transects Sb, Pb, As, and Mn display low to moderate enrichment, with highest EF values at site 64 for Sb and Pb, at site 296 for As, and at site 521 for Mn. Nevertheless, EF values generally less than 1 for Cr, Ni, and Zn suggest that relatively high contents of these elements may be ascribed to the natural background rather than to human activities.
4.8
Temporal distribution and source of trace metals
210
4.6
Background levels
To establish the heavy metal background level of the Strait of Sicily sediments, the method of maximum likelihood of lognormal distribution parameter was applied to sediments older than 1920, separated from the more recent ones by using the calculated sedimentation rates. The older sediments are reasonably thought to be poorly affected by intensive coastal urbanization and industrialization, because these have a major impact on natural coastal dynamics and took place in the south of Sicily only after the 1950 [27].
4.7
Evaluation of sediment contamination
Once obtained the background values, following various authors [28, 29, 30, 31, 32, 22, 33, 34], the Enrichment Factor (EF) was calculated as: (Metal/Al)sample EF = (Metal/Al)background
Pb radiometric analysis permitted to reconstruct for the first time the trace metal evolution 169 in sediments from the Strait of Sicily. In Figures 2 and 3 the chronological profiles of EFs for Sb, Pb, and As are represented, because of the sources of these elements appear to be influenced by contributions other than the main terrigenous input. Sb EF curves display that the element increase as from approximately the 1975 at the nearshore stations of PozzalloMalta and, especially, Sciacca-Pantelleria transects. Antimony is ubiquitously present in the environment as a result of nature processes and human activities and its presentday enrichment factor (relative to the Scnormalized value of typical crustal rocks) is in the order of 70 times [35]. The recent Sb contamination of coastal sediments of the two transects could be ascribed to intensive agricultural activities implying plastic waste incineration practices. On the other hand, positive excursions of Sb EF are recorded across the 1900 for both transects at sites 296, 407, and, 589
Marine Geology
Figure 2: EF age profiles of Sb, Pb, and As in sediments of the Sciacca-Pantelleria transect. markedly, 272 (Figures 2, 3). EF profiles for Pb and As exhibit similar coeval shifts, suggesting that these elements can be derived in part by a natural source probably related to the well known geothermal activity in the Strait of Sicily [36, 37]. Intriguingly, the most recent paroxysismic activity in the Pantelleria area was the submarine eruption of 1891 [38, 39]. Anthropogenic sources must be probably invoked to explain the progressive increase of Pb and As starting from about 1930s at many stations of both transects with further enrichments at some sites since 1960s. Interestingly, Pb EF values for sediments
590
younger than 1920 highlight that lead is homogeneously and moderately enriched in the studied area. This could suggest a no-point source for the element; in fact, it could be a Pb input linked to a major maritime route that affects the offshore of the Strait of Sicily. In the Strait of Sicily, the Sicilian coasts and the Sicilian shelf north of Malta are influenced by eastward flowing of the surface water of Atlantic origin (Modified Atlantic Water, MAW). Due to the bottom morphology, the area immediately in front of Pozzallo might be a zone of damping of current strength and velocity that justifies an enhanced partic-
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Figure 3: EF age profiles of Sb, Pb, and As in sediments of the Pozzallo-Malta transect. ous factors that control the spatial and temporal distribution of major and trace elements. Ti/Al and K/Al ratios, generally around 0.06 and 0.2, well represent the geochemical imprint of the detrital (riverine and aeolian) input from surrounding regions. Deviation from these values are referable to high energy depositional environments. Chronological profiles, traced on the basis 5 Concluding remarks of 210204 Pb radiometric data, suggest that Sb, As, and Pb are enriched in the sedi198 The high resolution geochemical study ments with respect to a background value performed on box-core sediments recov- calculated by elaboration of the element ered in the Strait of Sicily highlights variulate setting. On the whole, local oceanographic conditions are favourable to accumulation in this site of contaminant from local sources and, in addition, of pollutants derived from agricultural and industrial (importantly, a petrochemical plant at Gela town) activities along the southeastern Sicilian coasts.
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Figure 4: Maps representing the dynamic height and geostrophic velocity (a), temperature (b), and salinity (c) of the superficial water of the Strait of Sicily (July 2003; [40]). In (a) arrows indicate the geostrophic current velocity. contents in sediments older than 1920, with an increasing trend starting from approximately the 1960s. Point and no-point sources of contaminants might be intensive urbanization of the coast, a petrochemical plant, increasingly intensive agricultural activities in the southeastern Sicily and, for lead, maritime routes involving oil tanker traffic. At some stations, marked fluctuations of concentrations of major elements and trace metals along a +150 y record are interpreted as signals of geother-
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mal and/or magmatic activities in the Strait of Sicily. It is of interest that the core 134, in front of the Pozzallo coast, has a sediment accumulation rate as high as 0.40 cm¡yr−1 , relatively high Al contents proxy for abundant fine detritus and shows a recent increasing enrichment for many trace metals. An explanation for this could be that the eastward flowing of the AIS surface water in the Malta Shelf undergoes a damping of current strength and speed with a consequent higher particulate set-
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ting. This is due to a combined effect of haline barrier found by Mazzola et al. [40] the bottom morphology and of the thermo- at the passage from the Strait of Sicily to the Ionian Sea (Figure 4).
References [1] J. Garcia Lafuente, A. Garcia, S. Mazzola, L. Quintanilla, J. Delgado, A. Cuttitta, and B. Patti. Hydrographic phenomena influencing early life stages of the Sicilian Channel anchovy. Fish. Oceanogr., 11:31–44, 2002. [2] J. Garcia Lafuente, J.M. Vargas, F. Criado, A. Garcia, J. Delgado, and S. Mazzola. Assessing the variability of hydrographic processes influencing the life cycle of the Sicilian Channel anchovy, Engraulis encrasicolus, by satellite imagery. Fish. Oceanogr., 14:32–46, 2005. [3] P.F.J. Lermusiaux and A.R. Robinson. Features of dominant mesoscale variability, circulation pattern and dynamics in the Strait of Sicily. Deep-Sea Res., 48:1953– 1997, 2001. [4] K. Beranger, L. Mortier, G.P. Gasparini, L. Gervasio, M. Astrali, and M. Cr´epon. The dynamics of the Sicily Strait: a comprehensive study from observations and models. Deep-Sea Res. II, 51:411–440, 2004. [5] A.R. Robinson, J. Sellschopp, A. Warm-Varnas, W.G. Leslie, C.J. Lozano, P.J. Haley Jr., L.A. Anderson, and P.F.J. Lermusiaux. The Atlantic Ionian Stream. Journal of Marine System, 20:129–156, 1999. [6] M. Franzini, L. Leoni, and M. Saitta. Revisione di una metodologia analitica per fluorescenza X basata sulla correzione completa degli effetti di matrice. Rend. Soc. Ital. Mineral. Petrol., 31:365–378, 1975. [7] L.G. Schultz. Quantitative interpretation of mineralogic composition from X-ray and chemical data for the Pierre Shale. U.S. Geol. Surv. Prof. Pap, 1964. [8] E. Barahona, F. Huertas, A. Pozzuoli, and J. Linares. Mineralogia e genesi dei sedimenti della provincia di Granada (Spagna). Mineralog. Petrogr. Acta, 26:61– 90, 1982. [9] J. Husselman. On the routine analysis of carbonates in unconsolidated sediments. J. Sediment. Res., 36:622–625, 1966. [10] E.D. Goldberg. Geochronology with 210 Pb. In: Radioactive dating, I.A.E.A. Symp., pages 121–131, 1963. [11] P.G. Appleby and F. Oldfield. The calculation of 210221 Pb dates assuming a constant rate of supply of unsupported 210 Pb to the sediment. Catena, 5:1–8, 1978.
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[12] P.G. Appleby and F. Oldfield. The assessment of 210 Pb data from sites with varying sediment accumulation rates. Hydrobiologia, 103:29–35, 1983. [13] F. Oldfield and P.G. Appleby. Empirical testing of 210 Pb dating models for lake sediments. In: E.Y. Haworth and J.W.G. Lund (Eds.), Lake Sediments and Environmental History. Leicester University Press, Leicester, pages 93–124, 1984. [14] W.W. Dickinson, G.B. Dunbar, and H. McLeod. Heavy metal history from cores in Wellington Harbour, New Zealand. Environmental Geology, 27:59–69, 1996. [15] R.A. Ligero, M. Barrera, M. Casas-Ruiz, D. Sales, and F. L`opez-Aguayo. Dating of 281 marine sediments and time evolution of heavy metal concentrations in the Bay of C`adiz. Spain. Environ. Pollut., 118:97–108, 2002. [16] J.C. Varekamp. Trace element geochemistry and pollution history of mudflat and marsh sediments from the Connecticut coastline. Coastal Res., 11:105–123, 1991. [17] I. Cac¸ador, C. Vale, and F. Catarino. Accumulation of Zn, Pb, Cu, Cr, and Ni in sediments between roots of the Tagus estuary salt marshes, Portugal. Estuar. Coast. Shelf Sci., 42:393–403, 1996. [18] L.S. Chan, C.H. Yeung, W.W.S. Yim, and O.L. Or. Correlation between magnetic susceptibility and distribution of heavy metals in contaminated sea-floor sediments of Hong Kong Harbour. Environ. Geol., 36:77–86, 1998. [19] B. Rubio, K. Pye, J.E. Rae, and D. Rey. Sedimentological characteristics, heavy metal distribution and magnetic properties in subtidal sediments, Ria de Pontevedra, NW Spain. Sedimentology, 48:1277–1296, 2001. [20] P.W. Balls, S. Hull, B.S. Miller, J.M. Pirie, and W. Proctor. Trace metals in Scottish 225 estuarine and coastal sediments. Mar. Pollut. Bull., 34:42–50, 1997. [21] S. Covelli and G. Fontolan. Application of a normalization procedure in determining regional geochemical baselines. Environ. Geol., 30:34–45, 1997. [22] B. Rubio, M.A. Nombela, and F. Vilas. Geochemistry of major and trace elements in sediments of the Ria de Vigo (NW Spain): an assessment of metal pollution. Mar. Pollut. Bull., 40:968–980, 2000. [23] R. Wehausen and H.-J. Brumsack. Cyclic variations in the chemical composition of eastern Mediterranean Pliocene sediments: a key for understanding sapropel formation. Marine Geology, 153:161–176, 2000. [24] M.E. B¨ottcher, J. Rinna, B. Warning, R. Wehausen, M.W. Howell, B. Schnetger, R. Stein, H.J. Brumsack, and J. Rullokotter. Geochemistry of sediments from the connection between the western and the eastern Mediterranean Sea (Strait of Sicily, ODP Site 963). Palaeogeogr. Palaeoclimatol. Palaeocol., 190:165–194, 2003.
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[25] D. Rocca. High-resolution geochemical records from Late Quaternary Mediterranean sediments: palaeoclimatic and palaeoceanographic reconstruction of the Sapropel S1. Ph.D thesis, Univ. Palermo, 2005. [26] L. Civetta, M. D’Antonio, G. Orsi, and G.R. Tilton. The geochemistry of volcanic rocks from Pantelleria Island, Sicily Channel; petrogenesis and characteristics of the mantle source region. Journal of Petrology, 39:1453–1491, 1998. [27] A.B. Cundy, P.E. Collins, S.D. Turner, I.W. Croudace, and D. Horne. 100 years of environmental change in a coastal wetland, Augusta Bay, southeast Sicily: evidence from geochemical and palaeoecological studies. In: Black, K.S., Paterson, D.M. and Cramp, A. (Eds.), Sedimentary Processes in the Intertidal Zone. Special Publication, Geol., Soc., London, 139:243–254, 1998. [28] J.H. Trefrey and B.J. Presley. Heavy metals in sediments from San Antonio Bay and N.W. Gulf of Mexico. Envir. GeoL, 1:282–294, 1976. [29] S.A. Sinex and G.R. Helz. Regional geochemistry of trace elements in Chesapeake Bay sediments. Environ. Geol., 3:315–323, 1981. [30] G.R. Helz, S.A. Sinex, K.L. Ferri, and M. Nichols. Processes controlling Fe, Mn and Zn in sediments of 270 northern Chesapeake Bay. Estuar. Coast. Shelf Sci., 21:1–16, 1985. [31] H.L. Windom, S.J. Schropp, F.D. Calder, J.D. Ryan, R.G. Smith, L.C. Burney, F.G. Lewis, and C.H. Rawlinson. Natural trace metal concentrations in estuarine and coastal marine sediments of the southeastern United States. Environmental Science and Technology. 23:314–20, 1989. [32] C.L. Lee, M.D. Fang, and M.T. Hsieh. Characterization and distribution of metals in surficial sediments in south-western Taiwan. Marine Pollut. Bull., 36:464–471, 1998. [33] P. Woitke, J. Wellmitz, D. Helm, P. Kube, P. Lepom, and P. Litheraty. Analysis and assessment of heavy metal pollution in suspended solids and sediments of the River Danube. Chemosphere, 51:633–642, 2003. [34] K.M. Huang and S. Lin. Consequences and implication of heavy metal spatial variations in sediments of the Keelung River drainage basin, Taiwan. Chemosphere, 53:1113–1121, 2003. [35] M. Filella, N. Benzile, and Y.W. Chen. Antimony in the environment: a review focused on natural waters. I. Occurrence. Earth Sci. Rev., 57:125–176, 2002. [36] G. Etiope, A. Caracausi, and R. Favara. Methane emission from the mud volcanoes of Sicily (Italy). Geophysical Research Letters, 29(8):pp. 1215, 2002.
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[37] C.W. Holland, G. Etiope, A.V. Milkov, E. Michelozzi, and P. Favali. Mud volcanoes discovered offshore Sicily. Marine Geology, 199:1–6, 2003. [38] H.S. Washington. The submarine eruptions of 1831 and 1891 near Pantelleria. Am. J. Sci., 27:131–150, 1909. [39] N. Calanchi, P. Colantoni, P.L. Rossi, M. Saitta, and G. Serri. The Strait of Sicily continental rift systems: physiography and petrochemistry of the submarine volcanic centres. Marine Geology, 87:55–83, 1989. [40] S. Mazzola, B. Patti, A. Bonanno, A. Cuttitta, G. Basilone, G. Buscaino, S. Gontcharov, A.R. Vergara, G. Cosimi, V. Palumbo, M. Cancemi, L. Rollandi, G. Morizzo, C. Cavalcante, A. Arig`o, P. Sposito, C. Patti, L. Calise, R. Angotzi, R. Bellanca, Neri, M. Sprovieri, R. Di Leonardo, D.P. Rocca, and G. Tranchida. Applicazione e sviluppo di tecnologie idroacustiche avanzate per lo studio della dinamica di organismi marini mobili (Astamar). MIUR, Piano Ambiente Marino, Progetto n. 1, Cluster 10, Final Report, page pp. 205, 2004.
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Carbonate Sedimentation and Hydrodynamical Pattern on a Modern Temperate Shelf: the Bonifacio Straits (Western Mediterranean) G. De Falco1 , T. Batzella2 , A. Cucco1 , A. Ribotti1 , M. Borghini3 , S. De Muro2 1, Institute for Coastal Marine Environment, CNR, Oristano, Italy 2, Department of Earth Sciences, University of Cagliari, Cagliari, Italy 3, Institute of Marine Sciences, CNR, Pozzuolo di Lerici (SP), Italy giovanni.defalco@cnr.it Abstract The sedimentary features of the shelf of the Bonifacio straits, between Sardinia and Corsica (Western Mediterranean) were analyzed in order to evaluate the relationships between the production and transport of biogenic carbonate sediments and the morphological and hydrodynamical characteristics of the basin. Shallow seismic data (3.5 kHz) were acquired and superficial sediments were collected and analyzed for texture, mineralogical composition and skeletal carbonate taxonomical classification. A tridimensional hydrodynamic model was performed in order to simulate the current velocity at the seafloor. Two sectors of production of biogenic carbonate sediments in the Bonifacio straits were identified: the Posidonia oceanica meadows, and the flat relief located at 50 m depth characterized by Ma¨erl beds. Carbonate sediments produced in P. oceanica meadows are mostly trapped inside the meadow and partly exported outside the lower bathymetric limit forming a narrow belt. Sediment produced in the Ma¨erl beds are transported and accumulated in restricted basins. A wide belt, E-W trending, of mixed relict siliciclastic and biogenic carbonate sediments is located in the southern sector of the strait. The sedimentary pattern is strictly related to bottom current velocity, which shows lower values associated to the Ma¨erl beds and higher values associated to the mixed siliciclastic-carbonate sedimentary facies.
1
Introduction
Cool-water carbonate depositional environments are recognized for their primary importance for the continental shelves accounting for one third of the volume of global carbonate [1]. Modern carbonate deposits on temperate shelves have been described by many author in the last decades, from different world regions [2,
3]. Carbonate depositional environments are present in many sites in the Mediterranean [4, 5, 6] and are characterized by heterozoan skeletal assemblages [7, 8]. The production of biogenic carbonate debris is related to different benthic ecosystems of infralittoral and circalittoral zones [9, 8, 7]. Posidonia oceanica seagrass meadows are
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recognized as a major ecosystem producing biogenic carbonate sediments in the infralittoral from the shoreline up to 40 m depth [9, 10]. The carbonate production from P. oceanica epiphytes is generally low (0.19 - 0.43 g · m−2 · day−1 , equivalent to 69 - 157 g · m−2 · year−1 ) if compared to other Mediterranean coastal benthic ecosystems [9] or other tropical seagrass [11]. However, it has been observed that the sediments collected inside the P. oceanica meadows in different Mediterranean sites have high percentage of biogenic carbonate due to the fauna – e.g. gastropods, foraminifers, bivalves, echinoids, bryozoans - associated with the ecosystems (Fornos and Ahr, 1997). Biogenic carbonate particles have been found to be associated to the sandy fraction of sediments [12], and can affect the composition of adjacent beach sediments [13]. Considering the contribution of the vagile fauna associated to the meadow, the carbonate production rate from Posidonia meadows in western Sardinia was evaluated in the range 421-1,262 g CaCO3 m−2 · year−1 [10], one order of magnitude higher than those measured from P. oceanica epiphytes (69 - 157 g · m−2 · year−1 , [9]) and are among higher values between seagrass systems [11]. In the circalittoral the carbonate factory of mobile substrate is mainly related to the production from red algae which typically form two types of sediments: the Ma¨erl facies, with free living red algal braches and the ‘facies a` pralines’ dominated by rhodolits [4, 8]. Those facies are generally associate with biocalstic sediments deriving from the reworking of carbonate fragments, in the mobile substrate termed as ‘detritique coteir’ by [7]. Carbonate sediments are generally present where the input of terrigenous sedi598
ments is low, otherwise mixed carbonatesiliciclastic facies may occur due to the mixing of intrabasinal biogenic carbonate production with external input especially in the nearshore [4]. Hydrodynamics is an important factor which control shelf carbonate sedimentary processes, however few studies have compared the hydrodynamical pattern and carbonate sedimentation. It was shown that wind waves are the main intrabasinal forcing which influence carbonate production from P. oceanica meadows from central-western Sardinia [10]. Hydrodynamics were also cited as the main factor affecting the stability of algal nodules in the circalittoral of the Pontian islands [14], however no current measurements were reported. The aim of this paper is to analyze the relationships between the distribution of biogenic carbonate sediments and the hydrodinamical and morphological characteristics of the Bonifacio straits shelf (Western Mediterranean) in order to evaluate the factors which controls the production and transport of biogenic sediments.
2
Geological setting
The Corsica and Sardinia block rotated counter-clockwise between the Aquitanian (23 Ma) and the Burdigalian following the opening of the Ligurian Basin resulting in an oceanic basin that widens southwards, towards the Balearic Abyssal Plain. The geology of North Sardinia and South Corsica is dominated by the Corsican-Sardinian Hercynian batholiths, of upper Carboniferous-Permian age, outcropping for a total length of 400 km and for a width of 50 km between the two islands [15], characterized by calco-alkaline
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granitic formations with fields of basic dikes. The post Hercynian formations are limited in northen Sardinia to quaternary deposits of coastal areas and alluvial plains and very limited outcrops of Miocene limestone (Capo Testa). Marine sedimentary deposits of Miocene age (upper Burdigalian lower Langhian) outcrops for a surface of ca. 25 km2 in the southern Corsica from Bonifacio to Santa Amanza bay. The Bonifacio strait separates two distinct basins (Thyrrenian and western Mediterranean) and is an area of great relevance from the environmental point of view. It is part of the Bocche di Bonifacio International Marine Park. In the study sector the depth does not exceed 90 m. The seabed morphology in the northern shelf sector, toward the Corsica island, is characterized by tabular relief located at ca. 50 m depth, with incisions forming channels converging toward the center of the basin (Figure 1). Small basins delimited by morphological relief sometimes forming small islands (Lavezzi archipelago) are present toward the eastern sector. East-West oriented comet marks bedforms have been reported in the deepest area of the study sector toward the south [16].
3 3.1
Methods Numerical Modelling
Water circulation in the Bonifacio strait area was investigated by means of numerical modelling techniques. A tridimensional hydrodynamic model was applied to reproduce the wind and tide induced water currents in the whole Bonifacio strait area. The model uses finite elements for horizontal spatial integration and a semi-implicit
algorithm for integration in time. The finite element method is highly flexible due to the subdivision of the numerical domain in triangles varying in form and size. It is especially suited to reproduce the geometry and the hydrodynamics of complex coastal areas. Velocities are computed in the centre of each triangular element, whereas scalars are computed at each node. Vertically the model applies Z layers with varying thickness. Most variables are computed in the centre of each layer, whereas stress terms and vertical velocities are solved at the interfaces between layers. The model resolves the primitive equation, vertically integrated on each layer. Details of the numerical treatment are given in [17, 18, 19] and [20]. A set of numerical simulations have been carried out considering the main meteomarine forcing the water circulation in the interested areas. In particular, both astronomic tide and the main wind regimes, such as Mistral and Grecale winds, were taken into account to provide a complete characterization of the local hydrodynamics. Open boundary conditions were provided from tidal gauges and both meteorological and oceanographic operational model of the area.
3.2
Sediment sampling and geophysical data acquisition
A survey conducted in 2008 by R/V Urania of the CNR allowed to collect ca. 100 km of sub bottom record lines by using ship’s Chirp profiler at 3.5kHz (Figure 1). Chirp data were processed by using SeisPrho 1.2 software developed by ISMARCNR SEISPRHO: An interactive computer program for processing and interpretation 599
Marine Geology
Figure 1: Map of the Bonifacio strait with reported bathymetry, reflection seismic tracks and sediment sampling stations. of high-resolution seismic reflection profiles [21]. Sediment samples were collected in 95 station 2.5 km spaced by a Van Veen grab during year 2008 (Figure 1). Sediments were washed with distilled water and subsampled by quartering for different analysis. Grain size distribution was determined by dry sieving, carbonate content by Dietrich-Fruhling calcimeter. The mineralogical composition and the classification of bioclastic grains were quantitatively performed by analysis at binocular microscope with separation and weighting of the different components (mineral species, lithoclasts, bioclast of different taxonomical groups). Sediment samples were classified by mul600
tivariate statistics using cluster and factor analysis for facies classification [10].
4
Results and discussion
The distribution of current velocity (v [m¡s−1 ]) along the whole area was computed from numerical modeling in order to compare the hydrodynamics with sediment distribution (Figure 2). Two scenarios were taken into account: (i) current induced by Grecale wind (North-East, Figure 2a) and (ii) current induced by Mistral wind (North-West, Figure 2b). Those winds directions account the main wind events and the more intense wind storms in the Bonifacio straits. For both wind
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Figure 2: Current velocity at seabed level induced by Grecale (A) and Mistral (B) winds. directions higher values of current velocity at seabed are evident along an EastWest oriented belt which connects the eastern and western basins. Higher values of current velocity are also associated to the shallower areas (Lavezzi archipelago). The Mistral scenario (Figure 2b) also exhibits higher values of current velocity in the North-West sector, directed toward the South East direction. Three east-west trending chirp sections are reported in Figure 3. In the western sector
the seabed is mainly opaque and characterized by a tabular morphology (left side of the AB section, Figure 3). This morphology is probably related to the outcropping of the rocky sedimentary basement of Miocene age [16]. Toward the East a sedimentary layer covers the rocky basement filling the depressions between tabular reliefs. In the eastern side of the AB section the sedimentary layer is continuous and the basement morphology is evident below the superficial sediment deposit. The de-
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Figure 3: Selected seismic sections (chirp). The map shows the location of seismic lines. posit of sediment is limited toward east by the granitic outcropping of the of Lavezzi archipelago (CD section, right side, Figure 3). Another sedimentary deposits is evident in the EF section between two granitic ridges. Chirp data show that sediment deposition is confined into small basins controlled by the basement morphology, the Miocene sedimentary rocks to the west and the granitic outcropping to the east. The sedimentary deposits show limited thickness (<10 m). In particular, chirp data allow to identify two sectors of sediment accumulation, in the middle of study area, to the west from Lavezzi island, where small channels converge from north, and in the eastern area between two granitic outcrops (Figure 3, section EF). 602
The sediment distribution is shown in Figure 4. At shallower depth, along the coastline of Sardinia and Corsica, and in the marine sector between the minor islands, the seabed is mainly colonized by Posidonia oceanica seagrass (pattern 1 in Figure 4). Outside the deeper limit of the meadow a narrow belt of biogenic sands mainly characterized by association of bivalve and bryozoans was found (2, Figure 4). Those sediments were interpreted as deriving from the exporting of skeletal grains produced in the adjacent meadows. Ma¨erl facies were mainly found (3, Figure 4) associated to the tabular relief mainly locate at north. P. oceanica meadows and Ma¨erl beds can be considered as the main areas of production of biogenic carbonate
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Figure 4: Sediment distribution in the Bonifacio straits: (1) Posidonia oceanica meadows; (2) biogenic sands with bivalve and bryozoans; (3) Ma¨erl beds; (4) mixed biogenicsiliciclastic sands; (5) mixed biogenic-siliciclastic sandy gravels. sediments in the Bonifacio straits, similarly to other Mediterranean sites [10, 4, 5, 6, 8]. Carbonate sediments produced in P. oceanica meadows are mostly trapped inside the meadow [12] and partly exported outside the lower bathymetric limit forming a narrow belt. Sediment produced in the Ma¨erl beds are transported and accumulated in restricted basins limited by the morphology of the rocky basement. The small channels between tabular relief
are filled by mixed siliciclastic-biogenic sands (4, Figure 4). Those sediments are interpreted as deriving by the reworking of Ma¨erl facies, mixing with terrigenous components and deposition in the deeper sectors. Those mixed sands were transported toward the south and deposited in the small basins identified by chirp data. In the central sector a wide belt of mixed siliciclastic biogenic sandy gravels was found (5, Figure 4). This facies is constituted
603
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by coarse relict siliciclastic sediments with lithoclasts coated by encrusting algae. The comparison of sediment distribution and seabed current velocities shows that the mixed silicicalstic-biogenic sandy gravels are related to higher current values both from Mistral and Grecale winds. This finding highlights that current velocity at the seabed is a factor which limits the carbonate production in the circalittoral [14].
ated to Posidonia oceanica meadows and Ma¨erl beds. Sediments produced within the Posidonia meadows are mostly trapped inside and limited exported outside the deeper limit. Sediments produced in Ma¨erl beds are reworked transported and deposited in small depressions. Sectors with high seabed current velocity are associated with mixed relict siliciclastic-biogenic sediments mainly formed by encrusting algae. Seabed current velocity and morphology of the basement are the factors which controls 5 Conclusions the intrabasinal production and distribution The production of biogenic carbonate sed- of biogenic carbonate sediments in the ciriment in the Bonifacio strait can be associ- calittoral of Bonifacio straits.
References [1] C.S. Nelson. An introductory perspective on non-tropical shelf carbonates. Sedimentary Geology, 60:3–12, 1988. [2] H.M. Pedley and G. Carannante. Cool-water carbonate ramps: a review. Geological Society, London, Special Publications, 255:1–9, 2006. [3] N.P. James. The cool-water carbonate depositional realm. In: James, N.P., Clarke, J. (Eds.), volume 56. 1997. [4] M. Brandano and G. Civitelli. Non-seagrass meadow sedimentary facies of the Pontinian Islands, Tyrrhenian Sea: A modern example of mixed carbonate–siliciclastic sedimentation. Sedimentary Geology, 201:286–301, 2007. [5] F. Toscano and B. Sorgente. Rhodalgal–Bryomol temperate carbonates from the Apulian Shelf (Southeastern Italy), relict and modern deposits on a current dominated shelf. Facies, 46:103–118, 2002. [6] J.J. Forn´os and W.M. Ahr. Temperate carbonates on a modern, low-energy, isolated ramp: the Balearic Platform, Spain. Journal of Sedimentary Research, 67:364–373, 1997. [7] J.M. P´er`es and J. Picard. Nouveau manuel de bionomie benthique de la Mer M´editerran´ee. Rec. Trav. Stn. Mar. Endoume, Marseille. Fasc. Hors. Ser. Suppl., 31(47):5–138, 1964. [8] G. Carannante, M. Esteban, J.D. Milliman, and L. Simone. Carbonate lithofacies as paleolatitude indicators. Sedimentary Geology, 60:333–346., 1988. 604
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[9] M. Canals and E. Ballesteros. Production of carbonate particles be phytobentic communities on the Mallorca–Menorca shelf, northwestern Mediterranean Sea. Deep Sea Research, 44(3-4):611–629, 1997. [10] G. De Falco, M. Baroli, A. Cucco, and S. Simeone. Intrabasinal conditions promoting the development of a biogenic carbonate sedimentary facies associated with the seagrass Posidonia oceanica. Continental Shelf Research, 28(6):797–812., 2008. [11] E. Gacia, C.M. Duarte, N. Marba, J. Terrados, H. Kennedy, M.D. Fortes, and N.H. Tri. Sediment deposition and production in SE-Asia seagrass meadows. Estuarine Coastal and Shelf Science, 56:909–919, 2003. [12] G. De Falco, S. Ferrari, G. Cancemi, and M. Baroli. Relationships between sediment distribution and Posidonia oceanica seagrass. Geo marine letters, 20:50–57, 2000. [13] G. De Falco, E. Molinaroli, M. Baroli, and S. Bellacicco. Grain size and compositional trends of sediments from Posidonia oceanica meadows to beach shore, Sardinia, Western Mediterranean. Estuarine Coastal and Shelf Science, 58(2):299– 309, 2003. [14] D. Basso. Deep rhodolith distribution in the Pontian Islands, Italy: a model for the paleoecology of a temperate sea. Palaeogeography, Palaeoclimatology, Palaeoecology, 137:173–187, 1998. [15] L. Carmignani, G. Oggiano, S. Barca, P. Conti, A. Funedda, S. Pasci, and I. Salvadori. Geology of Sardinia (Explanatory notes of the Geologic map of Sardinia at 1:200,000 scale). page 247 pp, 2008. ´ [16] F. Pluquet. Evolution r´ecente et s´edimentation des plates-formes continentales de la Corse. page 300 pp, 2006. [17] G. Umgiesser and A. Bergamasco. Outline of a primitive equation finite element model. In Rapporto e Studi Istituto Veneto di Scienze, Lettere ed Arti, Venezia, XII:291–320, 1995. [18] G. Umgiesser, D.M. Canu, A. Cucco, and C. Solidoro. A finite element model for the Venice Lagoon. Development, set up, calibration and validation. Journal of Marine Systems, 51:123–145, 2004. [19] A. Cucco and G. Umgiesser. Modeling water exchanges between the Venice Lagoon and the Adriatic Sea. 2005. [20] C. Ferrarin, G. Umgiesser, A. Cucco, T.W. Hsu, A. Rol;, and C. Amos. Development and validation of a finite element morphological model for shallow water basins. Coastal Engineering, 55(9):716–731, 2008.
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[21] L. Gasperini and G. Stanghellini. SEISPRHO: An interactive computer program for processing and interpretation of high-resolution seismic reflection profiles. Computer & Geoscience, 35(7):1497â&#x20AC;&#x201C;1507, 2009.
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Relationships Between Multibeam Backscatter Sediment Grain Size and Seagrass Distribution from the Western Sardinia inner Shelf (Mediterranean Sea) G. De Falco1 , R. Tonielli1 , G. Di Martino1 , S. Innangi1 , S. Simeone2 , I.M. Parnum1 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, International Marine Centre, Oristano, Italy 3, Curtin University, Perth, Australia giovanni.defalco@cnr.it Abstract The use of multibeam acoustic data, including bathymetry, backscattering intensity and angular response, to classify seabed was evaluated in the inner shelf of central western Sardinia (Mediterranean Sea), a site characterized by sandy and gravelly sediments, Posidonia oceanica seagrass beds growing on sedimentary and hard substrates. A multibeam survey was carried out and ground-truth data were collected from 57 stations. 41 ground-truth stations provided data on sedimentary substrate without benthic vegetation, while the others provided data on Posidonia oceanica beds. Sediment samples were analyzed for grain size. Backscatter intensity is directly correlated to the weight percent of coarse fraction in sediments (1-16 mm) and inversely correlated to the weight percent of finer sediments (0.016-0.5 mm). Three sedimentary facies were identified, sandy gravels (SG), gravelly sands (GS) and slightly gravelly muddy sands (SGMS). Backscatter intensity significantly differed between GS, SG and SGMS. The range of backscatter intensity from Posidonia oceanica meadows overlapped values recorded for GS, and consequently it was not possible to distinguish Posidonia beds and gravelly sands only on the basis of backscatter data. Seabed mapping, including sediments and seagrass distribution was finally obtained by a combination of information deriving from backscatter data and morphological features from multibeam bathymetry validated by ground truth-data.
1
Introduction
Acoustic methods for seafloor mapping have been widely developed over the last decades. In particular, the development of swath bathymetry has allowed to obtain detailed maps of seabed morphology and the analysis of related acoustic backscatter has
made it possible to classify sediment types and habitat typology ([1, 2, 3], and references therein). Backscatter intensity data, the measure of sound scattered back toward the transmitter by acoustic reflection, have been shown to be related to sediment properties [4, 5, 3].
Marine Geology
Coarser sediments are more likely to result in higher backscatter intensity due to lower porosity, higher density and sound velocity, and larger roughness of the sediment-water interface [4]. In sandy sediments backscatter has been found to decrease with mean grain size [4]. Exceptions to those relationships can be due to inhomogeneities, the presence of near-surface gas or bioturbation in superficial sediments [6, 7, 8]. The acoustic response from the seabed can also be related to other properties of the sea-floor such as roughness [4], or to the presence of benthic organisms [1]. Consequently, multibeam bathymetry and related backscatter have been used to map benthic habitat [9]. The angular response is a measure of the variation of backscatter strength along the across-track direction [10, 11]. Backscatter strength near nadir is generally higher than values recorded in the outer swath because of differences between acoustic reflection (near nadir) and scattering (outer swath) [4]. The shape of the curve of angular response can be used to distinguish seabed typologies. Particularly, it was found that backscatter intensity in seabed covered by dense seagrass is almost independent from the angle of incidence, because seagrass scatters acoustic waves in all directions [12]. In contrast sandy and muddy sediments exhibit a strong decrease of backscatter strength toward the outer swaths [12]. Previous calibration studies between backscatter data and ground-truth analysis were carried out mainly in extra Mediterranean Seas, while few studies reported examples of multibeam backscatter calibration from the Mediterranean region. In particular the inner shelf of the Mediterranean Sea can be characterized by a complex and characteristic seabed with the presence of the en608
demic seagrass specie Posidonia oceanica, widely distributed across the whole basin up to 30-40 m in clear waters [13]. This seagrass can be associated both to sedimentary and rocky substrate. In this paper we evaluated the capability of multibeam acoustic data, including bathymetry and backscattering intensity, to distinguish different seabed typologies in a site of the Mediterranean Sea characterized by a complex seabed that includes sandy-gravelly sediments, Posidonia oceanica seagrass on hard (i.e bio-construction) and sedimentary substrate. Acoustic data have been statistically correlated to sedimentary data and calibrated by using ground truth on benthic habitat distribution in order to evaluate the acoustic variables which better discriminate seabed typologies.
2 2.1
Material and Methods Study area
The study area is the inner shelf of the central-western coast of Sardinia (Western Mediterranean Sea) from the Gulf of Oristano to the open sea toward the western direction (Figure 1). The gulf of Oristano has a surface of approx. 150 km2 , the maximum depth is ca. 24 m. About 70% (100 km2 m3 ) of the seabed surface of the gulf is covered by two Posidonia oceanica meadows on sedimentary substrate (Figure 1). The wider meadow of the central sector of the gulf of Oristano, partially included in the acquisition area (Figure 1) shows a sedimentary substrate composed by siliciclastic gravelly sands and sandy gravels [14]. The unvegetated area between the two meadows is a wide trough, considered the paleovalley of
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Figure 1: Map of the study area showing location of multibeam survey and the distribution of Posidonia oceanica meadows. the Tirso river, the major tributary of the gulf, during the last glaciations [15]. Toward the west this area is characterized by coarse sediments covered in the inner sector by muddy deltaic sediments [14]. A morphological threshold separates the gulf from the open sea [14], and has been interpreted as the result of beach rock formations [15]. This sector is characterized by the presence of hard substrate colonized by Posidonia oceanica (Figure 1) as shown by the Posidonia oceanica meadow mapping data available on the website of the Italian Ministry of Environment. The inner shelf outside the gulf toward the west is characterized by sands and gravelly sands with variable (siliciclastic and bio-
genic) mineralogical compositions [15].
2.2
Acoustic data acquisition and processing
A multibeam acoustic survey was carried out along the central western inner shelf of Sardinia (western Mediterranean, Figure 1) in June 2007. The survey was conducted aboard the R/V Tethis of the Italian National Research Council (CNR). The acquisition area falls within a bathymetric range between c.a -15 m and -50 m, from the meadow of Posidonia oceanica located inside the gulf of Oristano up to the deeper shelf toward the west direction. Multibeam data were acquired by using the 609
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Reson Seabat 8111 device at 100 kHz and following the acquisition standard of International Hydrographic Organization [16]. Acquisition was carried out along parallel track routes with 50%, or more, of overlapping between adjacent bands. Sound speed profiles in the water column were acquired every 8 hours. The acoustic footprint varies in size between ca 0.5 up to 1.3 m, depending on depth, the swath width was 150°. The Reson Seabat 8111 device provides the side scan and snippet option functionality that allows to acquire backscatter data. Snippet data are backscatter signal envelopes taken in each beam around the bottom detection time. The gain and the pulse length were maintained as constant as possible in order to optimize backscatter data. Backscatter data were processed by using the Centre for Marine Science and Technology (CMST)’s Multibeam sonar processing toolbox [17] which allowed to clear the backscatter intensity from distortion due to the different responses between the nadir and the external sector of the acquisition band. A 2 m spaced grid of backscatter strength data was computed. This tool also allows to extract the across track Angular Response values. The tool was developed in MATLAB® [17] and allows to carry out adjustments in order to normalize backscatter intensity. The applied corrections to Backscatter intensity depend on the frequency. The frequency of the system we used (Reson Seabat 8111R) is 100 kHz, an intermediate values among the various multibeam devices [17]. The geometric and physical parameters are, however, closely linked to the instrument and strongly influence the acoustic strength. One of the parameters on which backscatter strength mainly depends is the 610
angle of incidence. As illustrated in Parnum [17], the used tool first converts XTF data (extends Triton Format) to Matlab® format. It then calculates the space coordinates (XYZ) and the angle of incidence for every beam and every ping. After that the algorithm calculates the peak intensity and energy of backscatter corrected for the transmission power and energy respectively and for receive gain. The time varying gain (TVG) applied to the backscatter amplitude in the sonar system is removed and the backscatter intensity is then corrected for the actual transmission loss, including the acoustic energy spreading and in-water attenuation losses. Finally, the backscatter intensity and energy are corrected for the insonificaction and beam footprint areas respectively in order to estimate the seafloor backscatter strength. The applied mathematical functions are described in Parnum [17] and relevant literature.
2.3
Ground truth information
Based on multibeam bathymetry and backscatter data, ground truth estimation was carried out for a total of 57 georeferenced stations (Figure 2b). For each station a sampling with Van Veen Grab was carried out. Sediment samples were collected from 41 stations. In the remaining stations direct observations by divers were carried out in order to verify seabed typology. Sediment samples were transported in laboratory and analyzed for grain size distribution by using dry sieving and Galai CIS 1 laser system [14]. The coarser fraction (>8 mm) was analyzed by measuring the three axes of all pebbles with a manual vernier caliper. Data deriving from caliper measurements were converted into volume of the equivalent sphere and successively
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Figure 2: Acoustic data related to survey area reported in Figure 1. (A) Digital terrain model derived from multibeam bathymetry. Sectors 1, 2 and 3 are described in the text. (B) Backscatter intensity and position of the ground truth stations classified on the basis of seabed type. into weight, assuming the density of quartz for the pebbles, in consideration of their siliciclastic composition. The final grain size distribution was obtained with subsequent proportional recombination of the individual data sets.
2.4
Comparison of acoustic data and ground truth information
For each ground truth station backscatter intensity values were extracted from a radius of 10 m surrounding the position of the station. It has been shown that extraction radii in the range 8-20 m are required to achieve a strong acoustic calibration [3].
A total of 70 backscatter intensity values from each ground truth station were extracted. The mean values and standard deviation of backscatter intensity were then computed for each station. For each ground truth station the Angular Response (AR) curves were also extracted [10]. The AR curve can be divided in three domains (D1, D2 and D3) and for each domain different numerical variables can be computed in order to describe the shape of the AR curve [10]. In our case we computed the D1 and D2 mean backscatter intensity and the D2 slope values. Multivariate statistics was applied in order to analyze the relationships between acoustic variables and grain size data [4],
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with the Factor Analysis method. Mean backscatter intensity and variables describing the angular response curve (D1 and D2 mean backscatter intensity and D2 slope values) were used as acoustic variables. The weight percent of the grain size classes at 1 phi intervals of the gravelly and sandy fractions and silt+clay % content were used as sediment grain size variables. A total of 41 cases, including all ground truth stations with sedimentary substrate, were used for the analysis. To test the hypothesis that backscatter intensity varies with seabed typologies, an Analysis of Variance (ANOVA) was performed [18]. Six ground truth stations for each seabed typology were randomly chosen. At each station backscatter intensity data were collected in 30 points randomly chosen in a 10 m radius surrounding the ground truth station. The differences between backscatter intensities were tested by using the ANOVA with the following design: seabed typology (four levels), stations (six levels). The first factor was considered orthogonal and fixed while stations were considered random and nested in interaction with seabed typology. Prior to analysis, the homogeneity of variance was tested using Cochranâ&#x20AC;&#x2122;s test [19]. Significant differences among seabed typologies detected by ANOVA, were further analyzed using a posteriori Studentâ&#x20AC;&#x201C;Newmanâ&#x20AC;&#x201C; Keuls (SNK) tests [18].
pattern related with the distribution of Posidonia oceanica on sedimentary substrate as reported from previous maps (see Figure 1 for comparison). Several morphological features are clearly evident in the central sector. A flat relief irregularly shaped can be identified (2 in Figure 2a), while five linearly shaped subparallel reliefs, NNW-SSE oriented, are clearly visible (3 in Figure 2a). This central sector was reported as colonized by Posidonia oceanica on rocky substrate in previous maps (Figure 1). Toward the west the seafloor exhibits a more regular morphology with increasing depth up to ca. 60 m. The map of backscatter intensity, mapped at 5 db intervals, is reported in Figure 3b. Toward the east, sector 1, identified from the multibeam bathymetry, shows backscatter values between -115 db and 100 db. In the central sector the morphological features previously described (2 and 3) are not clearly related to variation of backscatter intensity, while toward the west patches of lower backscatter from the seabed (<-115 db) are clearly evident (Figure 2b).
3.2
Ground truth data
Based on seabed morphology and backscatter data, ground truth data were validated in the stations located in Figure 2b. Ground truth stations were set up to collect information on different patches detected by backscatter intensity data and on different morphological features. Par3 Results and discussion ticularly in correspondence of the linearly 3.1 Seabed morphology and shaped relief (3 in Figure 2a) ground truth map of backscatter intensity stations were positioned both at the top and in the valleys between adjacent reliefs. The mutlibeam bathymetry of the study site Sediment samples were collected in 41 stais reported in Figure 2a. In the eastern sec- tions (out of 57) reported in Figure 2b. tor (1) the seafloor relief shows an irregular In the remaining stations grab samples of 612
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Figure 3: Plot of factor scores deriving from factor analysis (see Table 2 for the significance of factor scores). Three sediment groups were identified based on their position on the plot. Posidonia oceanica leaves and coralligenous assemblages were collected. Direct observation by divers showed that the stations located in the eastern sector were characterized by the presence of Posidonia oceanica on sandy substrate, while the stations in the central sector were characterized by a hard substrate mainly bio-constructed and covered by Posidonia oceanica (Figure 2b). Along the linearly shaped relief (sector C, Figure 2a), patches of Posidonia oceanica on bio-constructed hard substrate at the top of the reliefs alternated to sedimentary substrate in the valleys between two adjacent reliefs were detected. Grain size analysis showed that sediments were mainly sandygravelly. Silt+Clay content, with an excep-
tion of one sample, did not exceed 20%, while 10 samples showed a gravel content >50%.
3.3
Relationships between seabed typology and acoustic response
Factor analysis allowed the extraction of two factors accounting for 76% of the total variance (Table 1). Factor 1, accounts for 44% of the total variance, and shows that backscatter intensity and energy variables of the Angular Response curve (D1 and D2 mean backscatter intensity) are correlated to the gravelly fractions and inversely correlated to medium-fine sands. Factor 2 (32% of the total variance) showed that the 613
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Table 1: Score coefficient of Factor analysis. Significant coefficients are in bold. slope of the central domain of the angular response curve (D2 slope) was related to coarse sand/very fine gravel fraction (0.54 mm) and inversely correlated to the finer fractions (<0.124 mm). The distribution of ground truth stations in the plot of Factor scores is reported in Figure 3. Stations can be grouped in three groups characterized by different grain size and acoustic responses. Stations in right side of the plot (positive values of Factor 1 score) are characterized by higher backscatter intensity values. Sediments from those stations are mainly Sandy Gravels (SG). Factor 2 separates stations characterized by lower D2 slope and finer grain size in the lower part of the plot (SGMS - slightly gravelly muddy sands) from stations with higher D2 slope and composed by Gravelly Sands (GS).
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Factor analysis allowed to identify groups characterized by different sediment typology and acoustic response. A further seabed typology, not comprised in the Factor Analysis, was the Posidonia oceanica meadows. Backscatter intensity from Posidonia oceanica on sedimentary substrate (eastern sector of the acquisition area) was in the range of -110/-115 db (Figure 2b). Backscatter intensity values from Posidonia oceanica on hard substrate was in the same range. Consequently the acoustic response from Posidonia oceanica, was compared to acoustic response from sedimentary seafloor regardless of seagrass substrate (sediments or hard bioconstructions). The differences in backscatter intensity between seabed typologies, including Posidonia oceanica, were tested by analysis of
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Table 2: Summary results of analysis of variance for backscatter intensity.
Table 3: Mean grain size composition and acoustic parameters of seabed typologies. variance (ANOVA). The analysis of variance showed significant differences for backscatter intensity for the three groups identified by factor analysis (SNK test, Table 2). Backscatter intensity values from the Posidonia oceanica meadow significantly differ from SGMS and SG, while any differences were not found between Posidonia oceanica and GS (Table 2).
3.4
Seabed mapping
The findings of this study show that the different seabed typologies of the Mediterranean inner shelf can be mapped using a combination of data, including backscat-
ter intensity, multibeam bathymetry and ground truth validation. Backscatter intensity alone was not able to distinguish benthic habitats. Multivariate statistics allowed to classify ground truth stations in three sedimentary seabed typologies: sandy gravels (SG), gravelly sands (GS) and slightly gravelly muddy sands (SGMS), characterized by different acoustic responses, in terms of backscatter intensity and shape of the angular response curves. The relationships between grain size and backscatter intensity deriving from multivariate statistics confirms the results of correlations previously found by other authors [4, 5]. In order to obtain a map of seabed typol615
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Figure 4: Whiskers plot of backscatter intensity for different seabed typologies (threshold values are indicated on the right side of the plot). ogy from acoustic data it is necessary to establish threshold values of backscatter intensity which allow to identify the spatial boundary between seabed type (sedimentary facies and/or benthic habitat). Results from multivariate statistics and ANOVA showed that this is only partially possible for the seabed typologies analyzed in this study. Threshold values can be identified as indicated in Figure 4, and can be used to separate three groups, (i) SG, (ii) SGMS (iii) GS + Posidonia oceanica beds (both on hard and soft substrate). The map of backscatter intensity limited to three classes separated by threshold values previously identified, is reported in Figure 5a. Based on this map it is possible to 616
identify the spatial boundaries of SGMS (<-116.8 db) and SG (>-109.8 db), while Posidonia oceanica beds and GS are represented in the same classes (-106.8/-109.8 db). A map of seafloor typologies can be created by combining backscatter data, multibeam bathymetry and ground truth validation (Figure 5b). The eastern sector (sector 1 in Figure 2a) is characterized by intermediate backscatter values, and irregular seabed morphology. In this sector ground truth validation showed the presence of Posidonia oceanica on sedimentary substrate, the lower western limit of which can be defined by changes in backscatter values and seabed morphology (Figure 2a, Figure 5ab).
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Figure 5: A) Map of backscattering using threshold values and B) seabed map obtained by using a combination of data, including backscatter intensity, multibeam bathymetry and ground truth validation. In the central sector Posidonia oceanica on hard substrate and coralligenous assemblages were associated to the top of seafloor reliefs, while in the depression between them, coarse sediments, distinguished in SG and GS on the basis of backscatter data, were found. The western sector is well resolved by the backscatter data, which allowed to identify the boundaries between SGMS, GS and SG.
4
Conclusions
Based on multibeam backscatter three sediment typologies can be discriminated: gravelly sands, sandy gravels and slightly gravelly muddy sands. Threshold values of backscatter intensity can be established in order to map the spatial distribution of those sediment types. Backscatter intensity from Posidonia oceanica beds showed the same range of gravelly sands. Backscatter intensity over Posidonia oceanica beds is independent from the seagrass substratum (sedimentary or hard bio-constructed) and is mainly due 617
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to the scatter and reflection of acoustic signal from leaves. The different seabed typologies analyzed in this study, gravelly and sandy sediments, Posidonia oceanica beds, can be mapped using a combination of data, including backscatter intensity, multibeam bathymetry and ground truth validation. Backscatter intensity alone was not able to distinguish all benthic habitats.
5
Acknowledgments
This work was funded by SIGLA project, from the Italian Ministry from Scientific Research. We gratefully acknowledge the captain and the crew of R/V Thetis for their effort during the data acquisition. We would like to thank Dr. Massimiliano Di Bitetto coordinator of the project, Gianluca Solinas for his contribution to grain size analysis and Patricia Sclafani for the revision of English text. We particularly thank Alexander Gavrilov for the critic revision of the manuscript.
References [1] K.W. Holmes, K.P. Van Niel, B. Radford, G.A. Kendrick, and S.L. Grove. Modelling distribution of marine benthos from hydroacoustics and underwater video. Continental Shelf Research, 28:1800– 1810, 2008. [2] T. Medialdea, L. Somoza, R. Le´on, M. Farr´an, et al. Multibeam backscatter as a tool for sea-floor characterization and identification of oil spills in the Galicia Bank. Marine Geology, 249:93–107, 2008. [3] T.F. Sutherland, J. Galloway, R. Loschiavo, C.D. Levings, et al. Calibration techniques and sampling resolution requirements for groundtruthing multibeam acoustic backscatter (EM3000) and QTC VIEW classification technology. Estuarine Coastal and Shelf Science, 75:447–458, 2007. [4] V.L. Ferrini and R.D. Flood. The effects of fine-scale surface roughness and grain size on 300 kHz multibeam backscatter intensity in sandy marine sedimentary environments. Marine Geology, 228:153–172, 2006. [5] J.A. Goff, B.J. Kraft, L. Mayer, S.G. Schock, et al. Seabed characterization on the New Jersey middle and outer shelf: correlatability and spatial variability of seafloor sediment properties. Marine Geology, 209:147–172, 2004. [6] J.C. Borgeld, J.E. Hughes Clarke, J.A. Goff, L. Mayer, and J.A. Curtis. Acoustic backscatter of the 1995 flood deposit on the Eel Shelf. Marine Geology, 154:197–210, 1999. [7] L. Fonseca, L. Mayer, D. Orange, and N. Driscoll. The high frequency backscattering angular response of gassy sediments: model/data comparison from the Eel River Margin, California. Journal of Acoustic Society of America, 111(6):2621–2631, 2002. 618
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[8] R. Urgeles, J. Locat, T. Schmitt, and J.E. Hughes Clarke. The July 1996 flood deposit in the Sanguenay Fjord, Quebec, Canada: implications for sources of spatial and temporal backscatter variations. Marine Geology, 184:41–60, 2002. [9] V.E. Kostylev, B.J. Todd, G.B.J. Fader, R.C. Courtney, et al. Benthic habitat mapping on the Scotian Shelf based on multibeam bathymetry, surficial geology and seafloor photographs. Marine Ecology Progress Series, 219:121–137, 2001. [10] J.E. Hughes Clarke, B.W. Danforth, and P. Valentine. Aerial seabed classification using backscatter angular response at 95 kHz. Shallow Water. NATO SACLANTCEN, Conference Proceedings Series CP, 45:243–250, 1997. [11] I.M. Parnum, P.J.W. Siwabessy, and A.N. Gavrilov. Identification of Seafloor Habitats in Coastal Shelf Waters Using a Multibeam Echosounder. 2004. [12] P.J.W. Siwabessy, I.M. Parnum, A.N. Gavrilov, and R.D. McCauley. Overview of coastal water habitat mapping research for Coastal CRC. 86, 2006. [13] G.A. Kendrick, N. Marba, and C.M. Duarte. Modelling formation of complex topography by the seagrass Posidonia oceanica. Estuarine Coastal and Shelf Science, 65:717–725., 2005. [14] G. De Falco, M. Baroli, A. Cucco, and S. Simeone. Intrabasinal conditions promoting the development of a biogenic carbonate sedimentary facies associated with the seagrass Posidonia oceanica. Continental Shelf Research, 28(6):797–812, 2008. [15] S. Carboni, L. Lecca, and C. Ferrara. La discordanza versiliana sulla piattaforma occidentale della Sardegna. Bollettino Societ`a Geologica Italiana, 108:503–519, 1989. [16] IHO International Hydrographic Organization. IHO Standards for Hydrographic Surveys. page 28, 2008. [17] I.M. Parnum. Benthic Habitat Mapping using Multibeam Sonar System. page 213, 2007. [18] A.J. Underwood (Ed.). Experiments in Ecology: Their Logical Design and Interpretation using Analysis of Variance. 1997. [19] B.J. Winer, D.R. Brown, K.M. Michels, and (Eds.). Statistical Principles in Experimental Design. 1991.
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Cold Seeps, Active Faults and the Earthquake Cycle: New Perspectives from Marine Geological Studies and Seafloor Observatories L. Gasperini1 and the SN-4 Team (P. Favali1,4 , G. Marinaro2,4 , G. Etiope2,4 , F. Furlan3,4 , F.Gasparoni3,4 ) 1, Institute of Marine Sciences, CNR, Bologna, Italy 2, National Institute of Geophysics and Volcanology, Roma, Italy 3, Tecnomare s.p.a, Venezia, Italy 4, SN-4 Team, Italy luca.gasperini@ismar.cnr.it Abstract A connection between cold seeps and active fault systems have been recognized since long time. This phenomenon could take the form of different diagnostic sedimentary features, such as pockmarks or mud and sand volcanoes, or the formation of carbonate crusts, or patches of reduced material at the seafloor. It is also known that an earthquake event could control gas emissions at cold seeps, and precursor gas emissions have been observed along several seismogenic faults, either on land or under water. The seafloor constitutes an interesting place to study these relationships, because evolution in time and space of these phenomena could be recorded in the sedimentary sequence more continuously than in other environments. Marine geological techniques have progressed both in the accuracy of geophysical imaging applied to seismogenic features, and in the development of multi-parameter geochemicalgeophysical-oceanographic monitoring system that could be deployed over relatively long period of time at the seafloor, the so called â&#x20AC;&#x153;seafloor observatoryâ&#x20AC;?. This leads to the opportunity of quantitatively study changes in the cold seep activity and how they could be controlled by faults in the earthquake cycle. The final purpose of these studies includes the mitigation of seismic hazards in coastal areas. An experiment of this kind is in progress in the Sea of Marmara, along the submerged track of the North Anatolian Fault system.
1
Introduction
A cold seep is fluid manifestation at the seafloor typically enriched with methane. often in the form of a brine pool. Unlike hydrothermal vents, which are associated to volcanism, cold seeps can be found in a number of geodynamic setting, including convergent and transform margins.
Another important difference between cold seeps and hydrothermal vents is that cold seeps emit at a slow and dependable rate, this could be an advantage once we decide to monitor emissions at a single vent. Association between cold seeps and active fault systems could take the form of different diagnostic sedimentary features, such
Marine Geology
Figure 1: Morphotectonic map of the North Anatolian Fault (NAF) in the Sea of Marmara compiled using data from different sources, including: onland topography (Shuttle Radar Topography Mission, NASA, JPL Lab, http://www2.jpl.nasa.gov/srtm/) Multibeam bathymetry [1, 2]. as pockmarks or mud and sand volcanoes, or the formation of carbonate crusts or patches of reduced material at the seafloor. It is also known that an earthquake event could control gas and fluid flow at a cold seep, and an increasing in gas emissions can be observed along several seismogenic faults, either on land or under the water, before and during and earthquake. One point not completely understood concerns coupling between seismic stress release and pore pressure changes during large earthquakes. Interactions between fluids and crustal strain (seismic and aseismic) have been widely studied onland (e.g.: [5, 6]), while the importance of fluids in the dynamics of submarine faults has been recognized in relatively recent times, together with the progress of deep seafloor exploration, in particular with the observation that fluid emission sites are often associ-
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ated with active faults (e.g. [7, 8, 9]). Fluid outflow sites have been widely observed associated to active faults in the Sea of Marmara, along the submerged portion of the North Anatolian Fault (NAF) northern strand (Figure 1), under the form of carbonate crusts, black patches, and bacterial mats (e. g. [10, 11, 4]). Free gas emissions are common, and appear to be influenced by the occurrence of earthquakes [12]. Due to these observations and the to its high geohazard potential, the Sea of Marmara has been identified as a unique natural laboratory to study relationships between fluids and seismicity through the EC funded ESONET Network of Excellence (European Seafloor Observatory Network). Part of this project is focused on studying and monitoring the western end of the Izmit Gulf, not far from the epicenters of the 1999 destructive earthquakes of Izmit
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Figure 2: Tectonic map of the Izmit Gulf derived from high-resolution multibeam bathymetric data (from [3] mod.). and Duzce (Mw 7.4 and 7.2, respectively) that caused heavily damages and a large number of casualties. The main reason for focusing our studies on this segment of the NAF is that it has been suggested that the Izmit earthquake caused an increase in tectonic loading in this area, that will be probably dissipate through a large earthquake in the next decades close to Istanbul.
2
The NAF below the Gulf of Izmit (Sea of Marmara)
The E-W aligned strike-slip deformation pattern characterizing the NAF system can be followed on land for about 1200 km from the Karliova triple junction to the Marmara Sea by analysing aerial and satellite images, as well as by using digital terrain models that highlight the presence of a relatively narrow principal displacement zone (Figure 1). However, the northern branch of the NAF, that accommodates over 80% of the Anatolia-Eurasia relative motion [10, 13], disappears below the
Sea of Marmara and cannot be traced using field and remote sensing observations. Prior to 1999, marine geological data in the Sea of Marmara and in the Gulf of Izmit were scant. Geophysical data, including high-resolution bathymetric maps and seismic reflection profiles have since been extensively collected, in the Gulf of Izmit and in the deep Marmara basins, as a consequence of the strong international effort that followed that disasters. The morphobathymetric map of the Izmit Gulf compiled with high-resolution mutibeam echosounder data [3] is among those results (Figure 2). It shows a complex pattern of releasing and restraining bends along the submerged trace of the NAF, causing the subsidence of three main basins separated by sills (Figure 2). The principal displacement zone of the NAF system forms, in the Izmit Bay, a series of steps and ridges, oriented E-W (strike-slip) NW-SE (transtensional) and SW-NE (trans-pressional) particularly in the center of the gulfs, and close to the Hersek Peninsula, where the deformation zone appears wider. The principal deformation zone converge again at the entrance of the Izmit Gulf, where it 623
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Figure 3: Underwater photograph of the Marmara seafloor close to the epicentre of the 1999 Izmit earthquake, collected using a ROV on board of R/V Urania during Marmara2001 expedition. Surface rupture is visible, as well as the effects of fluid seepage that produced black coloured sediments. focuses in a single, narrow E-W oriented furrow (Figure 2). A common pattern observed in the three basin is that deformation zones is wider at the basin centers and get narrow and more focused at the their edges. This has been particularly important to consider when we mapped the surface rupture caused by the 1999 Izmit earthquake, that was analysed using different type of geophysical data, as well as direct observations of the seafloor carried out mainly using ROV (Remote Operated Vehicles) system (Figure 3).
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3
Seismicity seeps
and
cold
The Izmit earthquake have shown that a strong correlation between seismicity and fluid and gas emission does exist. In fact, repeated surveys of the Gulf of Izmit before and after the event allowed to observe that the intensity of methane emissions from the seafloor increased after the event [14]. Gas emissions from the seafloor where observed analyzing acoustic backscatter from different sources, including echo-sounder and chirp-sonar profiles as well as sidescan sonar images. The acoustic waves used to insonify the seafloor, ranging in frequency from few kHz to hundreds of kHz, are â&#x20AC;&#x153;scatteredâ&#x20AC;? by gas bubbles in the water column, showing a typical pattern in the
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Figure 4: Typical acoustic backscatter pattern indicating the presence of gas in the water column. records (Figure 4). Because these data are collected to image the seafloor (multi- or single beam echosounder, side-scan sonar system) or the shallow sub-seafloor (chirp sonar or subbottom profiling systems) it becomes possible, and relatively easy correlating gas emissions and the presence of tectonic structures. The occurrence of gas and fluid emission and the availability of accurate morphotectonic maps of the NAF below the Sea of Marmara, stimulated by the strong international effort that followed the 1999 earthquakes, give us the opportunity of quantitatively study the relationship between cold seep, active faults and the earth-
quake cycle, within well constrained tectonic models. This could be carried out by repeated surveys or, more effectively by using seafloor observatories, that could continuously monitor different parameters at the seafloor, including seismic activity (broad band seismometers) and changes in the rate of emission of fluid and gas from the subsurface (methane and other geochemical sensors) during relatively long periods (years).
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Figure 5: Example of “black patch” from the Marmara seafloor. Photo collected using the Nautile submersible during MarNaut cruise (from [4]).
4
The black patches
Typical expressions of gas and fluid emission at the seafloor of the Sea of Marmara are the so called “black patches” (Figure 5), that mark unevenly the NAF track [10, 4]. Such features are typical indicators of relatively continuous emission of dissolved methane at the seafloor. The anaerobic oxidation of methane triggers a suite of geochemical reactions that ultimately result in production of black Fe and Mn sulfide mineral assemblages. Methane originates from below from microbial degradation of organic matter or from deeper thermogenic hydrocarbon generation and passes upwards as a dissolved component in pore fluid advection or as a buoyant gas phase. The sulfate source is the overlying water column, with sulfate diffusing across the sediment-water boundary. Secondary reactions include precipitation of 626
authigenic calcium carbonate, precipitation of Mn sulfides, and production of a variety of Fe. The latter dominate the sulfidic sediment of methane seeps and give it it’s black color (Figure 5). Fe oxyhydroxides often make up a significant percentage of ocean margin seafloor sediment and these are prone to biogeochemical reduction, often by a microbial intermediary. Black sediment patches are indicators of rapid, prolonged and current methane emissions. The emissions must be rapid enough to keep the bottom boundary layer completely anoxic. Flow measurements made in such environments have indicated outfow rates of centimeters to meters per day [15]. Sulfide oxidation is rapid [16] and in the absence of outflow, oxygenated bottom water will interact with the sediment and it will quickly (days) lose its black colour. One control on how long the site has been active is the presence or absence of carbonates.
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Figure 6: Deployment site of the SN-4 observatory in the Sea of Marmar.a The anaerobic oxidation of methane reaction leads to the production of authigenic calcium carbonate at and just below the seafloor. While there are many factors affecting the accumulation rate, as a general rule the lack of significant carbonate structures relates to a relatively short lifetime of venting activity on the order of some tens of years or less.
5
The SN-4 experiment
Several lines of evidence suggest that the Gulf of Izmit is a key observation point to understand the behaviour of the NAF. First, it is a site where intense transtensive deformation replaces the linear, almost pure strike-slip pattern that characterizes that regional tectonic boundary to the east; second, it has been the epicenter of the last de-
structive earthquake in the region; finally, gas emission correlated to a major earthquake has been recently observed. After several years of geo-marine studies we can conclude that that the Gulf of Izmit is characterized by the presence of a wide (>10 km) deformation zone across the trace of the NAF system [3, 17]. However, during the last 10,000 years, where accurate stratigraphic constraints are available, we are able to locate areas where the principal displacement zone of the submerged strand of the NAF is extremely narrow and shows an almost pure strike-slip deformation pattern. In some of these areas, due to the presence of piercing points, we were able to estimate very accurately horizontal and vertical deformation rates along these fault strands during the Holocene [3]. Moreover, recent analyses of geophysical
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Figure 7: The SN-4 seafloor observatory. imagery (multibeam bathymetry and sidescan sonar mosaics) and direct observations carried out using ROV system and the Nautile submersible (during MARNAUT cruise) suggest that the surface rupture of the 1999 terminated close to the Izmit Gulf entrance [18]. All these observations suggested to attempt an experiment: deploy a seafloor observatory in a key location of the NAF system to monitor seismicity and fluid seepage (Figure 6, [2]). Data collected during several oceanographic expeditions (MARMARA 20002001-2005 and 2009, guided our choice for the best location of the experiment, carried out using the SN-4 observatory [19]. SN-4 is a GEOSTAR (GEophysical and Oceanographic STation for Abyssal Research)-class seafloor observatory [20]. It is equipped with a number of sensors, including a broad-band, three component seismometer and a methane sensor. The selected location lay at â&#x2C6;ź200 m of water
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depth at the entrance of the Izmit Gulf, inside a canyon abandoned after the last episode of sea level rise, and since then passively displaced by the the NAF northern strand. The chosen SN-4 location appears particularly favourable for several reasons: 1. it is close to the western end of the surface rupture associated with the the 1999 Izmit earthquake, thus, it is the most probable area where the next earthquake affecting the fault strand towards Istanbul will enucleate. 2. it is an area characterized by gas and fluids emission related to the fault activity, as documented by acoustic images of the water-column and direct observations carried out using ROVs, and thus can be used to study correlations between these phenomena and mechanical behavior of faults; these information could be suddenly used for seismic risk assessments and to define early-warning strategies. This occur-
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rence is also confirmed by the presence of “black patches” at the seafloor observed during MarNaut cruise [18]. 3. gas emissions are not continuous and relatively weak. Most of gas remains trapped and accumulate in the sediment. This is a perfect condition for monitoring gas outputs related to seismicity. In fact, if the gas emission would be continuous and vigorous, it would have an intrinsic variability (flux pulsations) which would be impossible to distinguish, by sensors, from an increased emission eventually induced by an earthquake. 4. it is an area characterized by a “focusing” of the NAF principal deformation zone, which is constituted here by a single NAF, strike-slip segment; we measured strike-slip rate during geolog-
ical time span (10,000 years) along this strand. 5. it is a relatively accessible area, due to the moderate water depth (200 m) and the vicinity to the coastline. This is an extremely important point, once we will eventually decide to go toward a permanent (or semi-permanent) observatory. The objective of the SN-4 deployment will be to analyse possible relationships between fluid emission and seismicity and evaluate the use of these processes as earthquake precursors. The observatory has been deployed in September 2009 and its recovery is scheduled after 1 year. The time-series collected will hopefully shed light on correlation between seismicity and emission of gas and fluid from the subsurface along an active, seismogenic fault segment.
References [1] X. Le Pichon, A.M.C. Sengor, E. Demirbag, C. Rangin, et al. The active main Marmara fault. Earth Planet. Sci. Lett., (192):595–616, 2001. [2] L. Gasperini et al. MARM2009: marine geological study of the North-Anatolian Fault beneath the Sea of Marmara. Cruise Report. ISMAR-CNR Bologna Technical reports. 2009. [3] A. Polonia, L. Gasperini, A. Amorosi, E. Bonatti, et al. Holocene slip rate of the North Anatolian Fault beneath the Sea of Marmara. Earth Plan. Sci. Lett., 227:411– 426, 2004. [4] L. Geli et al. Gas emissions and active tectonics within the submerged section of the North Anatolian Fault zone in the Sea of Marmara. Earth Plan. Sci. Lett., 274:34–39, 2008. [5] R. Muir-Wood and G.C.P. King. Hydrological Signatures of Earthquake Strain. J. Geophys. Res., 98:22,035–22,, 1993. [6] M. Triquet et al. Radon emanation and electric potential variations associated with transient deformation near reservoir lakes. Nature, 399:137–141, 1999.
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[7] L. G´eli, P. Henry, T. Zitter, S. Dupre, et al. Gas emissions and active tectonics within the submerged section of the North Anatolian Fault zone in the Sea of Marmara. Earth Planet. Sci. Lett., 274(1-2):34–39, 1990. [8] Le Pichon et al. Fluid venting activity within the Eastern Nankai trough accretionary wedge — a summary of the 1989 Kaiko-Nankai results. Earth Plan. Sci. Lett., (109):303–318, 1992. [9] P. Henry, S. Lallemant, K. Nakamura, et al. Surface expression of fluid venting at the toe of the Nankai wedge and implications for flow paths. Marine Geology, 187:119–153, 2002. [10] Armijo et al. Submarine fault scarps in the Sea of Marmara pull-apart (North Anatolian Fault): Implications for seismic hazard in Istanbul. Geochem. Geoph. Geosys., 6, 2005. [11] T.A.C. Zitter, P. Henry, G. Aloisi, G. Delaygue, et al. Cold seeps along the main Marmara fault in the Sea of Marmara (Turkey). Deep-Sea Res., 2008. [12] Tary et al. Sea Bottom Observations from the Western Escarpment of the Sea of Marmara. BSSA, 2010. [13] B.J. Meade, B.H. Hager, S.C. McClusky, R.E. Reilinger, et al. Estimates of seismic potential in the Marmara region from block models of secular deformation constrained by GPS measurements. Bulletin of the Seismological Society of America, 92:208–215, 2002. [14] I. Kusc¸u, M. Okamura, H. Matsuoka, E. Gokasan, et al. Sea floor gas seeps and sediment failures triggered by the August 17, 1999 earthquake in the Eastern part of the Gulf of Izmit, Sea of Marmara, NW Turkey. Marine Geology, 215:193–214., 2005. [15] M.D. Tryon, K.M. Brown, and M.E. Torres. Fluid and chemical fluxes in and out of sediments hosting hydrate deposits on Hydrate Ridge, OR, II: Hydrological processes. Earth and Planetary Science Letters, 201:541–557, 2002. [16] F. Wang and P.M. Chapman. Biological implications of sulfide in sediment. A review focusing on sediment toxicity. Environmental Toxicology and Chemistry, 18:2526–2532, 1999. [17] M.-H. Cormier, L. Seeber, C.M.G. McHugh, A. Polonia, et al. North Anatolian Fault in the Gulf of Izmit (Turkey): Rapid vertical motion in response to minor bends of a nonvertical continental transform. Journal of Geophysical Researches B04102, doi: 10.1029/2005JB003633, 111:80–86, 2006. [18] L. Gasperini, A. Polonia, P. Henry, X. Le Pichon, et al. How far did the surface rupture of the 1999 Izmit earthquake reach in Sea of Marmara? 2010.
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[19] G. Marinaro, G. Etiope, P. Favali, L. Beranzoli, L. Gasperini, et al. SN-4 seafloor observatory in Marmara Sea. Rendiconti online Soc. Geol. It., 2:1–3, 2008. [20] P. Favali, L. Beranzoli, G. D’Anna, F. Gasparoni, et al. A fleet of multiparameter observatories for geophysical and environmental monitoring at seafloor. Ann. Geophys, 49:659–680, 2006.
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The Response of Benthic Foraminifera to Pollution: an Example from the Naples Harbour (Southern Italy) L. Ferraro1 , M. Sprovieri2 , D. Salvagio Manta2 , E. Marsella1 , S. Sammartino3 , P. Rumolo1 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, Institute for Coastal Marine Environment, CNR, Capo Granitola (TP), Italy 3, University of M´alaga, M´alaga, Spain luciana.ferraro@iamc.cnr.it Abstract Foraminifera are the most abundant marine protozoa. They show high biodiversity and abundance as an effect of their different ecological requirements. They may respond to environmental changes, leading to high production of specimens under favourable environmental conditions and to their disappearance under strongly unfavourable conditions. This paper represents a detailed investigation carried out on the benthic foraminiferal assemblages, in the highly contaminated marine sediments of the Naples Harbour (southern Campania).The aim of the research is to assess the response of foraminifera to organic and inorganic pollution. Combined use of multivariate canonical, cluster and principal component analysis provided an appropriate approach to explore the role played by contaminants (Heavy metals, PAHs, VOCs, TRPHs, PCBs) and physical parameters (grainsize, pH, Eh, TOC) on the spatial distribution of the benthic foraminifera. Obtained results provide evidence for a high and systematic non-linear response of the biota to the effects of the different pollutants. The reduced number of specimens per sample, their small dimensions and low diversity, testify the negative effects of a highly contaminated environment. In particular the results underline an important effect of the VOC on the distribution of some species, but also demonstrate that any kind of oversimplification would cancel the strong complexity of the biotic response to combined effects of different contaminants.
1
Introduction
Studies on pollution effects on benthic foraminifera and of the possible use of these organisms as proxies were initiated in 1960s, with the pioneer papers of Resig [1], Watkins [2] and Boltovskoy [3]. Since then, benthic foraminifera are increasingly used as environmental bioindicators, espe-
cially in polluted marine-coastal areas. In the last decades, a substantial number of investigations [4, 5, 6, 7, 8, 9, 10, 11, 12, 13, 14, 15, 16, 17, 18] confirmed the reliability of this organisms as useful tracers of marine sediments contamination. However, current literature lacks detailed information on benthic foraminiferal assemblage distribution modes and wide spectra
Marine Geology
of organic and inorganic contaminants, being generally restricted to analysis of a limited number of pollutants in different marine areas. That renders the potential attribution of the effects of the single classes or groups of analytes, on the distribution of benthic foraminiferal species extremely unlikely. The dataset here presented shows results of a large range of contaminants and chemical-physical parameters from the highly polluted area of the Harbour of Naples, combined to a detailed analysis of benthic foraminiferal assemblage. The Naples harbour is located in the southeastern Tyrrhenian Sea margin (Gulf of Naples, Figure1) and represents one of the larger of southern Europe. Because of its geographical position, it represents a natural receptor of municipal and industrial discharges coming from the city of Naples. In addition, it is a site of important commercial and tourist ship traffic and port activities (e.g., shipbuilding and good stocking). For this reason it represents a complex and dynamic environment, and can be considered a suitable natural laboratory to assess the potential of benthic foraminifera as tracers of anthropic pollution. In this study we have focused the attention to a more definitive and appropriate evaluation of the potential toxic effects of a variety of pollutants on the distribution patterns of benthic foraminifera species to verify the potential of these organisms as appropriate biomonitors of marine environmental pollution.
ure1), using a hydraulic 6 meter long vibrocorer with an inner diameter of 10 cm. The sediment was homogenized with a plastic spoon, placed into pre-cleaned highdensity polyethylene bottles for chemical analyses and stored at -18째C on board within an hour of collection. Benthic foraminiferal analysis were carried out on a sub-sample sieved with mesh width of 90 microns. The entire dried residue was microscopically analyzed and all the specimens were counted and percentages utilized for statistic analysis. Benthic species were identified following the generic classification of Loeblich and Tappan [19]. Samples for grainsize analysis were treated with H2 O2 solution, then washed and dried at 40째C. Grainsize analyses were carried out to establish the percentage of silt and clay (<63 mm) by a Laser Particle-Size Analyzer, and sand fraction (2 mm to 63 mm) by means of a microsieve. Analytical procedures for Geochemical analyses (Total Organic Carbon, heavy metals, Polycyclic Aromatic Hydrocarbons and Polychlorinated biphenyls) are reported in detail in Sprovieri et al., 2007. For analysis of the total recoverable petroleum hydrocarbons (TRPHs) was used the infrared spectrophotometric method ISO/TR 11046 (2005). Principal Component Analysis (PCA) and Canonical Correlation Analysis (CCA) are widely used to synthesize (reduce the dimensions of system) and interpret (evidence variability patterns of subsets of information), in order to extract as much more information as possible from such a non linear complex system. Furthermore, cluster analysis (with Kmeans method) is employed in order to in2 Material and Methods tegrate multivariate analysis with the aim Surface sediment samples (0-20 cm) were to improve consistency of analyses outcollected in the harbour of Naples, dur- put. Spatial georeferenced data were proing November 2004, from 84 stations (Fig- cessed with the ISATIS geostatistical soft634
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Figure 1: Location map of the study area with sampling stations.
Table 1: Basic statistics of VOCs, TRPHs and of 12 main analysed VOC congeners. ware package [20]. All the information was biphenyls (PCBs) are reported in detail in managed in a GIS georeferenced environ- [21]. Eh values range between -507 and ment, using ArcGis 9.2 software package. 498 mV with a median of -170mV, while pH maintains a very stable 7,6 value with an associated standard deviation of about 0,3 ([21]). Among the metals, Zn (173 Results 7234 mg · kg−1 ), Cu (12-5743 mg · kg−1 ), Pb (19-3083 mg · kg−1 ), V (37-2114 3.1 Chemical contaminants mg · kg−1 ), Cr (7-1798 mg · kg−1 ) and Results of grainsize, pH, redox poten- Ni (4-362 mg · kg−1 ) dominate with lower tial (Eh), Total Organic Carbon (TOC), concentrations measured for Sn (1-265 Heavy metals (Cr, Cu, Ni, Pb, V, Zn, Co, mg · kg−1 ), As (1-1121 mg · kg−1 ), Co Sn, Cd, Hg and As), Polycyclic Aromatic (1-30 mg · kg−1 ), Cd (0-3 mg · kg−1 ) and Hydrocarbons (PAHs) and Polychlorinated 635
Marine Geology
Figure 2: Median concentration values (µg · g−1 ) of the 57 analysed congeners of VOCs measured in the surface sediments of the harbour of Naples. Hg (0,01-139 mg · kg−1 ). Total concentrations of the analysed priority 16 US-EPA (1997) [22] PAHs range from 22 to 25,440 ng · g−1 dry wt with phenanthrene, fluoranthene, pyrene, benzo(a)antracene, chrysene and benzo(a)pyrene measured as the most important congeners and the 3 and 4 rings PAHs dominating in the studied sediments. In the following we present a more detailed description of concentration of TRPHs and VOCs, unprecedentedly reported. TRPHs show a wide range of variability with a median value of 266 µg · g−1 , a maximum of about 17140 µg · g−1 and minima corresponding to the detection limit (∼1 µg · g−1 , Table 1). VOCs show a median value of 1,715 µg · g−1 with a maximum of 5,105 µg · g−1 and standard deviation of 0,962 µg · g−1 (Table 1). Six of the 57 analysed congeners (specifically
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Methylene chloride, 2,2 Dichloropropeno, Toluene, Bromomethane, Chloromethane and 1,1,2,2 Tetrachloroethane) as reported in Table 1, constitute more than 40% of the total of VOCs content, while Benzene, Toluene, Ethyl-benzene and Xylens averagely represent about 15% of the total of volatile compounds (Figure 2).
3.2
Benthic foraminiferal assemblage
A total of 17 benthic species were recognized. The abundance values of the main benthic foraminifera species are reported in Table 2. Deformed individuals are present, with very low percentages in only three species (Table 2). Ammonia tepida (65,2%) represents the most abundant benthic species followed by Quinqueloculina spp. (24,4%) and Elphid-
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Table 2: Abundance values of benthic species, total of individuals, Shannon index, 0 Hmax , Relative divergence and Pielou index, with main basic statistics (mean, median, min, max, proportion, % non empty samples, empty samples and Skewness). ium spp. (10,5%) (Table 2). Benthic foraminifera show low mean abundance with A. tepida presenting 1,18 BF/g and Quinqueloculina spp. and Elphidium spp. presenting 0,44 and 0,19 BF/g, respectively (Table 2). Shannon diversity index (Hâ&#x20AC;&#x2122;) reveals a mean value of 1,03 for all the samples, with an average maximum theoretical 0 value of 3,22 (median value of Hmax Table 2). Computing the relative median divergence 0 from equiprobability (median (Hmax â&#x2C6;&#x2019; 0 0 H )/Hmax ) we obtain 67,2%, that is a median divergence of two thirds from the maximum biodiversity (Table 2). Such a result is mostly due to the scarce abundance of the species (a median value of 19% of present species of the 21 found) rather than to a real low evenness.
3.3
Multivariate statistical analysis on the Benthic foraminiferal assemblage
Principal Component Analysis (PCA) was applied to the standardized benthic foraminiferal dataset to verify the dominance of a reduced number of species. The PCA reveals the closeness of the species A. beccarii, A. tepida, S. fusiformis, Bulimina spp., Elphidium spp. and Quinqueloculina spp. with the variable Total Foraminifera (Figure 3A). However these findings call for a careful interpretation: firstly, PCA application is based on correlation parameter (regarding mean and standard deviation) while the percentages of species in the samples are computed in relation to median values (thus correctly allowing for asymmetry in the distribution) secondly, 637
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Figure 3: (A) Projection of the variables (benthic foraminifera species) considering the whole dataset. (B) Plot of the four outlier samples considering the whole dataset. (C) Projection of the variables (benthic species) excluding the four outlier samples. the included species A. beccarii and Bulimina spp. show very low percentages of non-empty samples, with respectively 16,7% and 1,7% (Table 2) These two considerations reduce the dominant species in the studied samples to A. tepida, Elphidium spp. and Quinqueloculina spp.. The PCA, applied to samples, reveals the presence of 4 evident outliers (NO51, NO75, NO82 and NO90, Figure 3B). After filtering, only A. tepida maintains its leading position in the Total Foraminifera assemblage variable (Figure 3C).
3.4
Multivariate statistical analysis
A combined use of PCA and CCA has been used in this work. In particular, PCA was employed to inspect the presence of outliers in the available dataset and to enhance consistence of analysis, while the function of correlation pattern 638
recognition is granted by CCA. PCA reliably suggested filtering out samples NO44, NO51, NO75 and NO90 from the whole dataset and on the reduced set, Canonical Analysis was applied. This approach has been conveniently employed for analysis of organic compounds, heavy metals and physical/sedimentological parameters covariance modes with the three dominant benthic species. PCA applied to the benthic foraminifera and sediment grainsize (Figure 4) reveals a substantial absence of structured spatial variability and a reduced control of the sedimentological parameters on the distribution patterns of the benthic species. PCA applied to the benthic foraminifera with Eh and pH excludes evident covariance patterns (Figure 5). Only A. tepida shows a certain degree of opposite variance with pH in the sediments (Figure 5). An evident inverse correlation between A. tepida and TOC emerges by the analy-
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Figure 4: PCA applied to the whole dataset of benthic foraminifera and sediment grainsize. sis of the benthic foraminifera compared to the different classes of organic compounds (Figure 6) with a substantial independence with VOCs, TRPHs and TotalPCB, while Quinqueloculina spp., Elphidium spp. show an opposite covariance only with VOCs and an apparently limited effect of the other organic compounds (Figure 6). On the other hand, canonical analysis performed on the biotic dataset and on the group of heavy metals (Figure 7) indicates opposite distribution patterns between Quinqueloculina spp. and Co, Ni and V. Unexpected scarce influence is recorded for Hg, Cr, Zn, Sn, Cu, Cd, As and Pb on the distribution patterns of
A. tepida and Elphidium spp. (Figure 7).
4
Discussion
The reduced number of specimens in the benthic foraminiferal assemblage and the low species diversity in the studied samples clearly suggest a significant role played by the complex ensemble of organic and inorganic contaminants on the wildlife of the marine bottom sediments and per se can be considered as appropriate earlier guideline for assessment of environmental pollution of marine areas. However, surprisingly, the classic highly toxic contaminants (PAHs 639
Marine Geology
Figure 5: PCA applied to the whole dataset of benthic foraminifera and Eh-pH values. and several of the analysed heavy metals) do not seem to have significant influence on the distribution of the three dominant benthic species. Also grainsize seems to show an independent variability with respect to the three benthic species. On the other hand, the TOC content seems to have a strong impact on the distribution modes of A. tepida, that, conversely, seems highly tolerant at high concentration and potential availability [23] of toxic heavy metals as already reported by [14, 17, 7] and references therein. The absence of any correlation between Total PAH, but also the highly toxic Total PCB and TRPHs, and the benthic species, directly suggesting that these 640
classes of pollutants have no direct control on the abundance patterns of those benthic foraminifera or alternatively indicating specific tolerance of foraminifera to toxic influences of organic compounds present in marine sediments. Possibly, the hydrophobic character of these contaminants associated to an extremely reduced potential of transferring from sediments to the seawater phase, renders them few bioavailable and harmful for the benthic foraminifera. Quinqueloculina spp. and Elphidium spp. show a statistically significant sensitivity to the distribution patterns of VOCs thus suggesting a primary control of such a class of contaminants on part of the benthic foraminiferal assemblage. This class
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Figure 6: CCA of Ammonia tepida, Elphidium spp. and Quinqueloculina spp. with TOC, VOCs, ÎŁPAH, ÎŁPCB and TRPH values. of contaminants, generated by incomplete combustion of organic matter related to industrial processes and motor traffic, was demonstrated to cause acute toxic effects on human health, cancer, neurobehavioral effects and adverse effects on the kidney and many of the congeners in this class are listed as priority pollutants. Particularly in the harbour of Naples, the anticovariance between Quinqueloculina spp. and Elphidium spp. with the distribution of VOCs probably reveals an unpredicted dose-mortality effect of this class of contaminants on the benthic species. Though the volatility and comparatively high solubility of the many chlorinated VOCs, will not tend to partition to aquatic sediments in the same way as the chlorinated pesticides, they will however be present in some regions at least in concentrations which could exert an effect on benthic life and
organisms that in contact with the sediments over prolonged periods with potential toxic ramifications for the local ecosystem of unknown magnitude. On the other hand, as already suggested by [14], Quinqueloculina spp. appears the most sensitive genera to high concentration of some heavy metals in the sediments in comparison to A. tepida and Elphidium spp.. However, it appears extremely difficult to quantify the effects of the single heavy metals on the distribution of that species and to thus separate the effects of VOCs on the spatial distribution patterns of these organisms. Moreover, the apparent high sensitivity of Quinqueloculina spp. to Co, Ni and V that are generally present in the studied sediments at values close to the natural background ([21]) compared to the more toxic concentrations measured for Hg, Cd and Pb (with which this species in partic-
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Figure 7: CCA of Ammonia tepida, Elphidium spp. and Quinqueloculina spp. with heavy metals. ular covariates), suggests extreme caution in defining linear oversimplification in the interpretation of the effective toxic role of heavy metals on the benthic species. In its ensemble, the evident high-level of complexity which characterizes the available dataset, reveal some degree of indeterminacy in identifying clear patterns of correspondence among the different variables and this limits the reliability of definitive biotic-abiotic relationships. Furthermore, we cannot exclude that a number of other unexplored factors, related for example to a significant influence of a highly dynamical environment, may directly and in some case primarily influence and drive the distribution of the different benthic species.
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We consider these results an important contribution to the understanding of the potential role played by organic and inorganic contaminants on the distribution patterns of the benthic foraminiferal assemblage. A wider spectrum of reference sites, contaminated by a reduced number of specific classes of contaminants, is besides required to explicitly and definitively determine major causes of pollutant effects (decrease in number of specimens, total absence of selected species, etc.) on the benthic foraminiferal assemblage (at level of species or genera) and thus suitably propose their use as reliably bio-monitors of the marine environment.
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References [1] J.M. Resig. Foraminiferal ecology around ocean outfalls off southern California. Disposal in the Marine Environment, pages 104–121, 1960. [2] J.G. Watkins. Foraminiferal ecology around the Orange County, California, ocean sewe outfall. Micropaleontology, 7(2):199–206, 1961. [3] E. Boltovskoy. Los Foraminiferos Recientes. page 510, 1965. [4] A. Albani, R. Serandrei Barbero, and S. Donnici. Foraminifera as ecological indicators in the Lagoon of Venice, Italy. Ecological Indicators, 7:239–253, 2007. [5] E. Alve. Benthic foraminifera reflecting heavy metal pollution in Sørliord, Western Norway. Journal of Foraminiferal Research, 21(1):1–19, 1991. [6] E. Boltovskoy, D.B. Scott, and F.S. Medioli. Morphological variations of benthic foraminiferal tests in response to changes in ecological parameters: a review. Journal of Paleontology, 65:175–185, 1991. [7] V. Yanko, M. Ahmad, and M. Kaminski. Morphological deformities of benthic foraminiferal test in response to pollution by heavy metals: implications for pollution monitoring. Journal of Foraminiferal Research, 28(3):177–200., 1998. [8] R. Coccioni. Benthic foraminifera as bioindicators of heavy metal pollution. A case study from the Goro Lagoon (Italy). Environmental Micropaleontology,, pages 71– 103, 2000. [9] J.P. Debenay, E. Tsakiridis, R. Soulard, and H. Grossel. Factors determining the distribution of foraminiferal assemblages in Port Joinville Harbor (Ile d’Yeu, France): the influence of pollution. Marine Micropaleontolog, 43:75–118, 2001. [10] L. Bergamin, E. Romano, M. Gabellini, A. Ausili, and M.G. Carboni. Chemical, physical and ecological characteristics in the environmental project of a polluted coastal area: Bagnoli case study. Mediterranean Marine Science, 4(1):5–20, 2003. [11] R. Coccioni, F. Frontalini, A. Marsili, and F. Troiani. Foraminiferi bentonici e metalli in traccia: implicazioni ambientali. Quaderni del Centro di Geobiologia dell’Universit`a degli Studi di Urbino, 3:57–92, 2005. [12] A.M. Samir and A.B. El Din. Benthic foraminiferal assemblages and morphological abnormalities as pollution proxies in two Egyptian bays. Marine Micropaleontology, 41:193–227, 2001. [13] D.B. Scott, F.S. Medioli, and C.T. Schafer. Monitoring of Coastal Environments Using Foraminifera and Thecamoebian Indicators. page 176, 2001.
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[14] L. Ferraro, M. Sprovieri, I. Alberico, F. Lirer, L. Prevedello, and E. Marsella. Benthic foraminifera and heavy metals distribution: A case study from the Naples Harbour (Tyrrhenian Sea, Southern Italy). Environmental Pollution, 142:274–287, 2006. [15] L. Ferraro, S. Sammartino, M.L. Feo, P. Rumolo, D. Salvagio Manta, E. Marsella, and M. Sprovieri. Utility of benthic foraminifera for biomonitoring of contamination in marine sediments: A case study from the Naples harbour (Southern Italy). Journal of Environmental Monitoring, 11:1226–1235, 2009. [16] F. Frontalini and R. Coccioni. Benthic foraminifera for heavy metal pollution monitoring: A case study from the central Adriatic Sea coast of Italy. Estuarine, Coastal and Shelf Science, 76:404–417, 2008. [17] F. Frontalini, C. Buosi, S. Da Pelo, R. Coccioni, A. Cherchi, and C. Bucci. Benthic foraminifera as bio-indicators of trace element pollution in the heavily contaminated Santa Gilla lagoon (Cagliari, Italy). Marine Pollution Bulletin, 58:858–877, 2009. [18] E. Romano, L. Bergamin, M.G. Finora, M.G. Carboni, A. Ausili, and M. Gabellini. Industrial pollution at Bagnoli (Naples, Italy): Benthic foraminifera as a tool in integrated programs of environmental characterisation. Marine Pollution Bulletin, 56:439–457, 2008. [19] A.R. Loebelic and H. Tappan. In Foraminiferal Genera and their Classification. 2:970, 1988. [20] ISATIS On-Line Help. Geovariances. page 403, 2000. [21] M. Sprovieri, M.L. Feo, L. Prevedello, D. Salvagio Manta, S. Sammartino, S. Tamburrino, and E. Marsella. Heavy metals, polycyclic aromatic hydrocarbon and polychlorinated biphenyls in surface sediments of the Naples harbour (southern Italy). Chemosphere, 67:998–1009, 2007. [22] US-EPA (United States Environmental Protection Agency). Chemicals known to the state to cause cancer or reproductive toxicity. Environmental Protection Agency, 1997. [23] P. Adamo, M. Arienzo, M. Imperato, D. Naimo, Nardi G., and Stanzione D. Distribution and partition of heavy metals in surface and sub-surface sediments of Naples city port. Chemosphere, 61:800–809, 2005.
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Distribution of Benthic Foraminiferal Assemblages on the Southern Campanian Continental Shelf L. Ferraro1 , I. Alberico 2 , F. Lirer 1 , F. Budillon 1 , M. Vallefuoco 1 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, Interdepartmental Center for Environmental Resarch, University of Napoli â&#x20AC;&#x153;Federico IIâ&#x20AC;?, Napoli, Italy luciana.ferraro@iamc.cnr.it Abstract Benthic foraminifera in 250 samples collected on the Southern Campanian continental shelf were analysed. Q-mode cluster analysis resulted in the identification of six distinct foraminiferal assemblages, reflecting different environmental settings. The distribution of the assemblages shows a distinct zonation, that is manly attributed to grainsize and composition of sediment, bathymetry and to geographical distribution of the impact of the river supply. Cluster I (4-20 m), is related to the areas influenced by the rivers outflow on sandy-silt and very sandy silt substrates, dominated by the species Ammonia beccarii and Eggerella scabra; cluster II (10-30 m), associated with vegetated environments and/or sandy and gravelly sandy substrates, presents a characteristic epiphytic fauna composed by Cibicides lobatulus, Rosalina bradyi and Rosalina obtusa; cluster III (10-50 m) shows a typical infralittoral assemblage mainly related to sandy silty and very sandy silty substrates, not directly affected by the rivers outflow, with the dominance of Ammonia tepida and Elphidium granosum; cluster IV related to silty sandy substrates of the Sorrento Peninsula shelf, is chracterised by Elphidium crispum and Cibicides lobatulus; cluster V (30-100 m) is composed by silty substrates with the dominance of the opportunistic species Valvulineria bradyana while cluster VI is characterized by an outer-shelf assemblage on silty bottoms, with prevalent Cassidulina carinata
1
Introduction
Foraminifera are perfectly suitable for environmental studies, being recorders of environmental changes because of their wide distribution over all marine environments. A large number of physical and chemical parameters such as temperature, salinity, depth, sediment, oxygen, food and as well as biological interactions, influence the distribution of benthic foraminifera [1, 2, 3], making them useful tools for ecological
and environmental interpretations [4]. At present, only a few studies dealing with the distribution of benthic foraminifera assemblages on the southern campanian continental shelf, most of which analysed only confined areas of the shelf [5, 6, 7, 8, 9, 10]. In the present paper we tried to defined the spatial distribution of benthic foraminifera along southern campanian continental shelf from the Sorrento Peninsula to the Gulf of Policastro and their relationship with some environmental parame-
Marine Geology
Figure 1: Geological map of the investigated area showing the mainland outcrops with deposits and generalized bathymetry with morphology of marine area. ters.
2
General setting of the study area
The study area extends over different physiographic settings: the Sorrento Peninsula offshore, the Gulf of Salerno, the Cilento offshore and the Gulf of Policastro (Figure 1). The continental shelf surrounding the Sorrento Peninsula widens to the north and narrows in the southern sector, is bordered by a smooth shelf-break at about 140/170 m of water depth to the north and by a sharp shelf-break at about 100/120 m of water 646
depth to the south (Figure 1). According to [11], sediments in the northwestern sector are mainly composed by sands, in the depth range of 9/50 m and by silty-sands and sands in the outer part of the shelf, while in the southern sector the seafloor is mainly characterised by sands with Posidonia oceanica prairies and subordinately by silty-sands. The Gulf of Salerno is characterizes by a northern sector with a narrow shelf (about 1-2 km wide) bordered by a shelf-break at about 100/120 m of water depth and a steep slope, and by a southern sector which displays a shelf wide up to 35 km with a shelfbreak at about 180-200 m of water depth and a deep border (Figure 1). In this area
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Figure 2: Location map of the studied area with sampling stations. the continental shelf is strongly influenced by the Sele River input which represents the most important source of fresh water, nutrient and terrestrial continental organic matter. The Cilento offshore is rather irregular with a deep shelf-break (140 to over 200 m deep; Figure 1). The inner shelf around Licosa Cape and Palinuro Cape displays wide extensions of Posidonia oceanica meadows down to 30 m of water depth. The northern part of the continental shelf of the Gulf of Policastro is rather broad (up to 10 km), whereas the eastern and southeastern sectors are narrower in some places (only 2 km) and have slopes with up to 10% inclination (Figure 1). The rivers Lambro, Mingardo and Bussento, which flow into the northern part of the Gulf of Policastro, represent the main source of organic matter to the shelf.
3
Material and methods
A total of 250 bottom samples were collected by IAMC, CNR - Naples, during two oceanographic cruises (1998 and 2003), on the southern Campanian continental shelf (Figure 2). Sediments were collected using a Van Veen grab. For sedimentological and micropaleontological analyses two aliquots of undisturbed sediment were taken from the top 3-5 cm of seabed from each sample. In the laboratory the collected sediments for benthic foraminiferal analysis were sieved over sieves with 125 ¾m, dried at 60°C and then weight. If necessary, the samples were split with a microsplitter, and a minimum of 200 to 300 specimens were counted using a binocular microscope. Benthic foraminifera were classified according to [12] and [13]. The grain size analysis of samples was car647
Marine Geology
Figure 3: Distribution of sediment grain-size at the sea floor. ried out on sediment processed with peroxide solution, then washed and dried at 40°C. The >1000 µm granulometric fraction was analyzed using a microsieve (2000 µm, 1400 µm and 1000 µm), while the fraction smaller than 1000 µm has been analyzed by a Laser Particle-Size Analyzer. The percentages of sand, silt and clay fractions were calculated for each sample. Q-mode cluster analysis was performed on 34 selected species to assess the main composition of benthic foraminiferal assemblage. The cluster analysis has been applied to the selected foraminiferal data set and samples using the nearest neighbor method and Euclidian distance measure.
4
Results
The continental shelf of the study area is generally characterised by coarse sediments along the coast, while pelagic facies with fine sediments (silt) is present 648
to open sea (Figure 3). The texture and composition of sediments show that their main source is the Sele River (northern sector) and the Bussento River (southern sector), with other contributions come from some small rivers (Asa, Tusciano Alento, Lambro, Mingardo, Castocucco and Abatemarco). Particularly the shelf of the Sorrento Peninsula, also according to [11], is mainly characterized by silty sandy and sandy sediments. In the Gulf of Salerno the shelf is characterized by sandy silt and very sandy silt sediments in the infralittoral zone (050 m), with a decreasing in grain-size from 50 to 200 m, where the seabottom is dominated by silty sediments (Figure 3). From Agropoli to Punta degli Infreschi the seafloor is generally characterized by silty-sands and sands (0-50 m) with some patches of sands and gravelly sands with Posidonia oceanica off the Solofrone, Alento, Lambro and Mingardo
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Figure 4: Spatial distribution of the six defined clusters in the study area. river mouths (Figure 3). From 50 m seaward the seafloor is mainly composed by silty sediments (Figure 3). In the Gulf of Policastro (from the coast to about 50 m of water depth) the sediments off the Castrocucco, Lao and Abatemarco river mouths are composed by silty sands and sands down to 50 m depth, while in front of the Bussento river mouth sandy silty and very sandy-silty bottoms are dominant (Figure 3), with a decreasing in grainsize below this depth (silt). A total of 229 benthic foraminiferal species, belonging to 104 genera, were identified in the total assemblages. Generally, well-preserved foraminiferal tests dominate in all the samples. The Q-mode cluster analysis grouped the samples into six homogeneous clusters (Figure 4). Cluster I, grouped the samples between 4
and 20 m (Figure 4), mainly characterized by silty-sandy and sandy-silty sediments (Figure 7). The assemblage is composed of the species Ammonia beccarii and Eggerella scabra, and is mainly found close to the river mouth where the fauna presents a low density and biodiversity (Figure 5). In the Gulf of Salerno this cluster occurs mostly off the Asa and Sele river mouths while towards the south-eastern part of the Campanian coastal zone (from Capo Palinuro to the Gulf of Policastro) its distribution is recorded off the rivers Lambro, Mingardo, Bussento, Castrocucco, Lao and Abatemarco (Figure 4). Cluster II, is composed by the epiphytic species Cibicides lobatulus, Rosalina bradyi and Rosalina obtusa and is mainly located in the central sector of the study area from Agropoli to Capo Palinuro (10 - 30 m, Figure 4), where the sea floor is dominated by sands,
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silty sands and gravelly sands with Posidonia oceanica (Figure 7). Cluster III, is located between 10 and 50 m, and seems to be characteristic of sites less influenced by the river input (Figure 4); the assemblage is characterized by high biodiversity (Figure 5). Cluster IV, is confined to the western part of the study area (Sorrento Peninsula, Figure 3) with a fauna dominated by the species Elphidium crispum and Cibicides lobatulus (Figure 5). Cluster V, is located parallel to the coast, from 30 to about 100 m (Figure 4) and groups the samples es-
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sentially composed by silty sediments (Figure 7), probably characteristic of sites influenced by the river plume; the assemblage shows the highest biodiversity (Figure 5) with the dominance of Valvulineria bradyana. Finally cluster VI, groups the stations farthest away from the river mouth (Figure 4), the sediments of this area are all composed by silt (Figure 7); the fauna is characterized by Cassidulina carinata, Valvulineria bradyana and Bulimina marginata (Figure 5).
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Figure 5: Average Abundance Values (AV) of the 34 selected species within the six clusters: A16: Ammonia beccarii, A18: Ammonia inflata, A21: Ammonia tepida, A25: Amphycorina scalaris, A30: Asterigerinata mamilla, B2: Bigenerina nodosaria, B8: Bolivina alata, B16: Buccella granulata, B17: Bulimina aculeata, B18: Bulimina costata, B19: Bulimina elongata, B22: Bulimina marginata, C2: Cassidulina carinata, C8: Cibicides lobatulus, E2: Eggerella scabra, E5: Elphidium crispum, E6: Elphidium cuvilleri, E7: Elphidium granosum, E14: Elphidium punctatum, G1: Gavelinopsis praegeri, G10: Globocassidulina subglobosa, G19: Gyroidina umbonata, H3: Hyalinea baltica, M6: Melonis barleanum, P3: Planorbulina mediterranensis, Q15: Quinqueloculina seminulum, R5: Reussella spinulosa, R8: Rosalina bradyi, R11: Rosalina obtusa, S10: Sigmoilopsis schlumbergeri, S17: Sphaeroidina bulloides, T4: Textularia calva, U1: Uvigerina mediterranea, V1: Valvulineria bradyana. 651
Marine Geology
Figure 6: Comparison between the six clusters and mean depth values.
Figure 7: Composition of sediment in the six defined clusters.
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5
Discussion and conclusion
The southern campanian continental shelf shows a very diversified benthic foraminiferal fauna. The Q-mode cluster analysis allowed recognising of six assemblages (Figure 4), probably strong controlled by an interaction between organic matter, water depth and grainsize of seabottom. The distribution of benthic foraminiferal assemblages in the studied area has been already defined, by different authors, for selected sectors characterize by rivers outflows [5, 7, 6]. However, the present study underlines that the river contributions seem to have a relevant importance in the composition, structure and distribution of benthic foraminiferal assemblage along the southern campanian continental shelf. Generally, the distribution patterns of the clusters show a distinct tendency with the change in bathymetry. Indeed, as shown in Figure 6, the mean depth in each cluster increases moving from cluster I to cluster VI, with a significant shift from cluster III to cluster IV. Between 4 and 50 m water depth three different biofacies are present: an assemblage (with Ammonia beccarii and Eggerella scabra; Cluster I) on silty-sandy and sandy bottom, related to river contributions and located off of the main rivers mouths (Figure 4). This assemblage dominated by Ammonia beccarii and Eggerella scabra (Figure 5), is characterised by a low diversity attributed to the opportunistic behaviour of these species in the organic matter rich environment influenced by freshwater [14, 15, 16, 5, 6]. An assemblage (with Cibicides lobatulus, Rosalina bradyi and R. obtusa Cluster II; Figure 5) on silty-
sandy, sandy and gravelly-sandy bottoms with vegetation cover mostly Posidonia oceanica (Figure 3) [16, 17, 18]. This assemblage well fits with that found by [5] in the northern sector of the Salerno Gulf, by [6] in the Policastro Gulf and by [7] in different sectors of the continental shelf between Agropoli and Capo Palinuro (Southern Tyrrhenian Sea). A typical infralittoral assemblage on silty bottom with Elphidium granosum and Ammonia tepida (Cluster III; Figure 4), that suggests an environment with more stable salinity conditions, probably with lower fresh water input and not directly affects by the rivers outflows (Figure 4). The presence of Elphidium granosum reveal a distinct changes in percentage of organic matter as reported by [15] and [16] in the Adriatic Sea. Also A. tepida, which shows increasing values from cluster I to cluster III (Figure 5), prefers shallow, saline and brackish environments [19, 4, 16, 20, 21, 22]. Cluster IV is located on the shelf around the Sorrento Peninsula and groups stations from 9 to102 m water depth (Figure 3). The dominance of Elphidium crispum and Cibicides lobatulus (Figure 4) in the whole area can be attributed to the silty-sandy and sandy bottoms with Posidonia oceanica prairies [13]. The increasing biodiversity in the deepest biofacies, sheltered from the river influence (Cluster V and VI, Figure 5), suggests that the environmental conditions are not stressfull for the benthic foraminiferal fauna, which is dominated by Valvulineria bradyana and Cassidulina carinata; the distribution of these biofacies seem to mainly follow the bathymetry and sediment grain-size. However the relevant percentages of Valvulineria bradyana, within cluster V and of Cassidulina carinata within cluster VI (Figure 5), can be still 653
Marine Geology
related to the presence of percentage of organic matter at the seafloor probably dispersed to the outer shelf. In the Adriatic Sea, [16] decribes a minimum water depth of 40 m for V. bradyana, in the outer part of the organic carbon enriched clay belt. [23] shows that this species is a relevant marker of high productivity, while [24] and [25] observed that it is an opportunistic species, living in sediments containing high amounts of organic matter. Cluster VI, shows a very similar distribution to
cluster V but with a more significant positive correlation with depth (Figure 6). Here the fauna is dominated by Cassidulina carinata which can be found in areas with sustained organic input as reported for the continental shelf and open slope of East New Zealand [26] and for the French upper middle bathyal station in the Bay of Biscay, where it appears to respond quickly to labile organic matter input by a reproductive event [27].
References [1] G. Colom. Foraminiferos ibericos. Introducci´on al estudio de las esp´ecies bent´onicas recientes. Consejo Superior de Investigaciones Cientificas,Patrono Ju´ade la Cierva, Barcelona,, 38:245, 1974. [2] J. Murray. Ecology and Paleoecology of Benthic Foraminifera. page 398, 1991. [3] F.J. Jorissen. Benthic foraminiferal succession across Late Quaternary Mediterranean sapropels. Marine Geology, 153:91–101, 1999. [4] J.P. Debenay, B. Beck Eichler, J.J. Guillou, P. Eichler-Coelho, C. Coelho, and E. Porto-Filho. Comportement des peuplements de foraminif`eres et comparaison avec l’avifaune dans une lagune fortement stratifi´ee: La Lagoa da Conceic¸ao (SC, Br´esil). Revue Pal´eobiologie, 16(1):55–75, 1997. [5] F. Sgarrella and D. Barra. Distribuzione dei foraminiferi bentonici nel Golfo di Salerno (Basso Tirreno, Italia). Boll. Soc. Nat., 93:51–110, 1984. [6] F. Sgarrella, D. Barra, and A. Improta. The benthic foraminifers of the Gulf of Policastro (Southern Tyrrhenian Sea, Italy). Bollettino Societ`a Naturalisti, 92:77– 114, 1983. [7] M.G. Coppa, B. Russo, and G. Siani. The Holocene foraminiferal assemblages of the continental margin between Agropoli and Capo Palinuro (Tyrrhenian Sea). Boll. Soc. Paleont. Ital., 2:67–91, 1994. [8] M.G. Coppa and A. Di Tuoro. Preliminary data on the holocene foraminifera of the Cilento continental shelf (Tyrrhenian Sea). Rev. Esp. Paleont., 10(2):161–174, 1995. [9] L. Ferraro, T. Pescatore, B. Russo, M.R. Senatore, C. Vecchione, M.G. Coppa, and A. Di Tuoro. Studi di geologia marina del margine tirrenico: la piattaforma 654
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continentale tra Punta Licosa e Capo Palinuro (Tirreno Meridionale). Boll. Soc. Geol. It., 116:473–485, 1997. [10] L. Ferraro and F. Lirer. Morphological variations of benthonic foraminiferal tests from Holocene sediments of the Punta Campanella Shelf (south Tyrrhenian Sea). Proceedings of the Second and Third Italian Meetings on Environmental Micropaleontology. Grzybowski Foundation Special Publication, 11:45–59, 2006. [11] M. De Lauro, F. Budillon, G. Cristofalo, L. Ferraro, F. Molisso, A. Simioli, and C. Violante. Studio geomorfologico nell’area marina protetta ”Punta Campanella” tramite rappresentazione georeferenziata di alta risoluzione della batimetria del fondo. Rint.11, 2001. [12] A.R. Loeblich and H. Tappan. Foraminiferal Genera and their Classification. page 970, 1988. [13] F. Sgarrella and M. Moncharmont Zei. Benthic foraminifera of the Gulf of Naples (Italy): sistematics and autoecology. Boll. Soc. Paleontol. It., 32:145–264, 1993. [14] S. Donnici, R. Serandrei Barbero, and G. Taroni. Living benthic foraminifera in the Lagoon of of Venice (Italy): population dynamics and its significance. Marine Micropaleontology, 43:440–454, 1997. [15] S. Donnici and R. Barbero. The benthic foraminiferal communities of the northern Adriatic continental shelf. Marine Micropaleontology, 44:93–123, 2002. [16] F.J. Jorissen. The distribution of benthic foraminifera in the Adriatic Sea. Marine Micropaleontololy, 12:21–48, 1987. [17] M. Langer. Recent epiphytic foraminifera from Vulcano (Mediterranean Sea). Revue de Paleobiologie, 2:827–832, 1988. [18] M. Langer. Epiphytic foraminifera. Marine Micropaleontology, 20:235–256, 1993. [19] J.P. Debenay, B. Beck-Eichler, M. Fernandez Gonzalez, R. Mathieu, C. Bonetti, and W. Duleba. Les foraminif`eres paraliques des cˆotes d’Afrique et d’Am´erique du Sud, de part et d’autre de l’Atlantique: comparaison et discussion. G´eologie de l’Arique et de l’Atlantique sud, pages 463–471, 1996. [20] G.A. Seiglie. Foraminifers of Guayanilla Bay and their use as Environmental indicators. Rev. Esp. Micropaleontol, 7:453–487, 1975. [21] A. Almogi-Labin, R. Simav-Tov, A. Rosenfeld, and E. Debard. Occurence and distribution of the foraminifer Ammonia beccarii tepida (Cushman) in water bodies, Recent and Quaternary of the Daed Sea Rift, Israel. Marine Micropaleontology, 26:153–159, 1995.
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[22] A. Pascual, J. Rodriguez-Lazaro, O. Weber, and J.M. Jouanneau. Late Holocene pollution in the Gernika estuary (southern Bay of Biscay) evidenced by the study of Foraminifera and Ostracoda. Hydrobiologia, 475/476:477–491, 2002. [23] P.J.J.M. Verhallen. Late Pliocene to Early Pleistocene Mediterranean mud-dwelling foraminifera influence of a changing environment of community structure and evolution. Utrecht Micropaleontological Bulletin, 40:1–219, 1991. [24] L. Bergamin, L. Di Bella, and M.C. Carboni. Valvulineria bradyana (Fornasini) in organic matter-enriched environment. Il Quaternario, Ital. Jour. Quater. Sci., 12:51–56, 1999. [25] C. Fontanier, F.J. Jorissen, L. Licari, A. Alexandre, P. Anschutz, and P. Carbonel. Live benthic foraminiferal faunas from the Bay of Biscay: faunal density, composition, and microhabitats. Deep-Sea Research I, 49:751–785, 2002. [26] B.W. Hayward, H. Neil, R. Carter, H.R. Grenfell, and J.J. Hayward. Factors influencing the distribution patterns of recent deep-sea benthic foraminifera, east of New Zealand, Southwest Pacific Ocean. Marine Micropaleontology, 46:139–176, 2002. [27] C. Fontanier, F.J. Jorissen, G. Chaillou, C. David, P. Anschutz, and V. Lafon. Seasonal and interannual variability of benthic foraminiferal faunas at 550 m depth in the Bay of Biscay. Deep-Sea Research. Part 1, 50:457–494, 2003.
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The Dark Side of the Mediterranean Geological Record: the sapropel layers and a case study from the Ionian Sea L. Capotondi1 , L. Vigliotti1 , C. Bergami1 , F. Sangiorgi1,2 1, Institute of Marine Sciences, CNR, Bologna, Italy 2, Faculty of Science, Utrecht University, Utrecht, Nederlands lucilla.capotondi@bo.ismar.cnr.it Abstract A peculiar feature of the Neogene Mediterranean marine and land sequences is the quasi-cyclic occurrence of organic carbon-rich layers named sapropels. Their occurrences in the sedimentary record usually correspond to periods of enhanced monsoon rainfall during precession minima and summer insolation maxima. Nevertheless, the causal factors that led to their formation are still highly debated. Integrated multi-proxy investigations document that during sapropel deposition important changes occurred in the entire water column: freshwater lenses in the surface waters led to stratification of the water column and to hypoxic or totally anoxic bottom waters. Sapropels offer the unique opportunity to perform studies on climatic, oceanographic and environmental changes at an extraordinary resolution allowing detailed insights into short-scale climatic fluctuations. Micropaleontological and magnetic signatures demonstrate that oceanographic conditions conducive to sapropel formation were not confined to the eastern Mediterranean sea but occurred also and possibly simultaneously in the entire Mediterranean. The differences appear a consequence of different preservation, changes in water column depth and local hydrographic conditions. Here we report the main features characterizing the youngest Mediterranean sapropel (S1) deposited during the Holocene in the Ionian basin
1
Introduction
Neogene sediments of the Mediterranean Sea are characterized by the occurrence of organic carbon-rich (with TOC usually > 2%) layers named sapropels [1]. Their formation seems to be mainly controlled by astronomical forcing, usually corresponding to phases of precession-induced insolation maxima [2] leading to periods of wetter climate in the Mediterranean region. The word “sapropel” was introduced in literature by Potoni`e [3] to indicate dark
sediments with decomposing organism deposited under stagnant water. Sapropel is a contraction of the literal translation of the German words F¨aulniss and Schlamm into ancient Greek (sapros and pelos, meaning putrefaction and mud respectively). At first discovered in marine sediments in the ’50 [4, 5] the sapropels have been subject of a plethora of studies during the last decades and several models have been proposed to explain the mechanism leading to their deposition (e.g., [6]). In spite of the large numbers of studies
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performed in the last 40 years, and two Ocean Drilling Program expeditions (ODP legs 160 and 161 in 1995), the causes of sapropel formation are still debated. At present, Mediterranean sediments are characterized by low organic carbon content (<0.5% organic carbon) due to low surface water nutrient levels (hence generally low productivity) and oxic bottom waters due to a vigorous deep-water formation and circulation (details in [7]). Sapropels were instead likely deposited under hypoxic or anoxic deep water conditions strictly related to deep water stagnation due to a heavily reduced or halted deep (or even intermediate) water circulation. The deposition of organic-rich layers such as the sapropels must have hence required major changes to the present water circulation patterns. At first, their occurrence was linked to improved preservation of organic matter under anoxic bottom water conditions. The explanation for the anoxia was related to density stratification of the water column, limiting water circulation and supply of oxygen to deep water. At least for the youngest sapropel the water stratification takes into account increasing freshwater input originated from the Nile River during period of enhanced monsoon regime in the equatorial region [8, 9, 10, 11, 12, 13, 2]. Other proposed triggering mechanisms consider also the increasing organic matter accumulation related to the enhancement of primary productivity [14, 15, 16, 17, 18, 19, 2]. Certainly an increase in productivity in the Mediterranean cannot be achieved with the present-day oceanographical features and for this reason some authors invoked a reversal of the water circulation [16] or a shoaling of the density gradient (pycnocline) into the photic zone [17]. The debate concerning the roles and the importance of 658
productivity and preservation in sapropels formation is still on-going. Even if Total Organic Carbon (TOC) content was originally chosen as key parameter to identify sapropels, other proxies have been successively used. Among those, geochemical elements (e.g., Fe, Mn, Al, S, Ba, V, Mo, As, I) [16, 20, 21] magnetic parameters such as susceptibility, anysteretic (ARM) and isothermal remanence (IRM) [22, 23, 24, 25], microfaunal taxa [8, 26, 27] are sensitive to the anoxic conditions and can be considered indicators of the sediments deposited under anoxic conditions. Sapropels offer the unique opportunity to perform climatic, oceanographic and environmental reconstructions at an extraordinary resolution allowing detailed insights into short-scale climatic and environmental oscillations. Here, after a general discussion on the main features of sapropels, we present, as a case study, the sapropel S1 in a sediment core (ET 99-M11) collected in the Ionian Sea.
2
Sapropels across Mediterranean
the
Most of the literature on sapropels considers the eastern Mediterranean, where these layers have been at first discovered and defined. However organic rich layers occur also in the western part of the basin, although they appear more scattered and less developed [28]. This implies that the Sicily channel sill may act as a barrier against processes that favour sapropel deposition. The present-day anti-estuarine Mediterranean circulation, clearly influenced by the Gibraltar strait sill must have still played a role even at the time of the
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sapropel formation. TOC contents in the western Mediterranean show maximum values of up to 6% (in the Tyrrhenian Sea) and appear to decrease toward the western areas were TOC hardly reaches the 2-3% [28]. For this reason Murat [28] suggested to redefine a sapropel as an organic-rich lithologic layer (ORL) deposited in open sea, with at least 0.8% of TOC. The sapropels or organic rich layers observed in the Western Mediterranean are more scattered throughout the time even if the timing (at least for the late Quaternary <400 Ka) is synchronous with that of sapropels deposition in the eastern part of the basin [29]. The only exception is the Alboran Sea, where background TOC values are already above 0.4-0.5% [28, 26] and the timing of occurrence of ORL does not appear to match the timing observed for sapropels in the other parts of the basin. In particular the youngest ORL observed at ODP Sites 976 and 979 is strictly coincident with the Younger Dryas (∼12.5-11.5 Ka BP) which means it is older than Sapropel S1 deposited during the early Holocene (9.5-6.Ka BP). However the micropaleontological signal indicates that the planktonic foraminifer assemblage typical of a sapropel layer occurs above the ORL and is coincident with the timing of sapropel S1. It is also noteworthy that TOC peaks were observed in cores from the Alboran Sea at around 55 Ka BP corresponding to the rarely found sapropel S2 [26] and also in correspondence of the I-cycle 4-6 (3138 Ka BP) that never expresses sapropels in other areas. The presence of sapropels in both the western and eastern Mediterranean indicates that in some cases the entire basin responded in unison at precession minima. This seems true especially in correspondence of interglacial stages as in-
dicated by the finding of warm sapropels such as the S5 (∼125-119 Ka BP) throughout the basin.
3
Magnetic signature of sapropels
In sub-oxic/anoxic conditions as those typical of sapropel depositions, magnetic Fe oxides dissolve, resulting in a decrease of magnetic concentration coupled with an increase in magnetic grain-size and in coercivity [31, 32, 33, 24, 34]. The bacterial degradation of organic matter is a diagenetic process leading to sulphate reduction and methanogenesis that clearly occurs in sapropels. As evidenced during analysis of core from the ODP leg 160, the process can be so severe to be responsible of a magnetic enhancement observed in several sapropels recovered in the Eastern Mediterranean Sea [23]. A ferrimagnetic iron-sulphate phase is responsible of a high magnetization that is directly proportional to the organic carbon content found within the sapropels [23]. On the base of the magnetic properties Larrasoana [25] grouped the sapropels in the 3 different types corresponding to different anoxic conditions. The dissolution or enhancement of the magnetic signal within sapropel layers represents a distinctive feature that can be easily identified already in whole-core measurements (e.g. magnetic susceptibility, K) and represents a marker that can be used for tuning a sedimentary sequence to orbital scale (Figure 1). Concentration-related magnetic parameters such as K, ARM, IRM can be indicative of the presence of sapropel layers even in sediments where the lithologic expression is not clearly visible (“missing sapropels”, [22]). Possibly, the best indi659
Marine Geology
Figure 1: Magnetic parameters (ARM and NRM) from Borehole Prad1-2 (Adriatic Sea) and ODP Site 967 (eastern Mediterranean). Minimum ARM values related to magnetic dissolution occur in the sapropelitic layers recognized in the Adriatic Sea [30] and correlated with insolation maxima. Peaks in the NRM observed at ODP Site 967 reflect precipitation of iron sulphides in correspondence of sapropel layers. izes the sapropel layers consisting of an increase in the occurrence of warm subtropical species Globigerinoides ruber (var. rosea and alba) and the SPRUDTS group (see [36]), or, in many cases, in the exclusive presence of high productivity water indicators such as Neogloboquadrina dutertrei and Globigerina bulloides [8, 10]. Some authors (see for details [26] and [37]) noted that it is possible to recognize the late Quaternary sapropels on the basis of the quantitative and qualitative variations in planktonic foraminifera assemblage, either when sapropels are deposited in warm or in cold intervals. Generally, sapropel 4 Foraminifera signature layers deposited during warm intervals are characterized by peaks abundance of G. ruof sapropels ber and occurrence of G. ruber var. rosea Several studies have shown that an unusual (e.g. S1, S3, S5, S7, S9, and S10), while planktonic foraminiferal fauna character- abundance of N. dutertrei characterizes the cator of a sapropel layer is the ARM because it is more sensitive to the presence of fine grained ferrimagnetic materials. On the contrary, magnetic susceptibility could also represent an unclear indicator as it is also influenced by the presence of the paramagnetic contribution of the clay minerals. Another distinctive feature of the sapropel layers is the precipitation of Fe oxides at the oxygenation front that causes higher magnetic intensities [35] due to the formation of iron oxides and also Fe sulphides [23].
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sapropel layers deposited in cold intervals (e.g. S4, S6 and S8). Both these species inhabit surface waters and are well documented in literature to be related to low salinity [36, 38, 8, 39, 40]. Qualitative foraminifer analyses show that many taxa are characterized by an increase in size (e.g. Orbulina universa, G. ruber > 150 micron) and that the tests are thinner with large pores, often covered by diffuse pyrite crystals and sometimes entirely or partially filled by pyrite. During times of sapropels formation benthic foraminiferal abundances and diversity strongly decrease till the near exclusive presence of deep infaunal taxa. Microfauna even disappear in some levels suggesting extremely low oxygen values on the sea bottom [41, 27]. The diversity reported in the pre- and post sapropel benthic assemblages from different sites suggest that the evolution of the dysoxic-anoxic conditions, as well as the re-oxygenation pattern at the end of the stagnant period, were characterized by spatial and temporal variability, possibly controlled by basin physiography and, in particular, by the water column depth [27, 42, 43, 44]. In addition, high resolution benthic foraminiferal distributional trend during times of sapropel S5 and S6 deposition indicated that water column stratification and deep-water formation was unstable and reflected the climate fluctuations at millennial time scale. Based on these biological features the sapropel deposition appears the result of different oceanographic phases related to stratification of the water column, reduced ventilation in intermediate/deep water and changes in nutrient regimes.
5
Geochemical signature of sapropel
Sapropels are generally reported to contain higher concentrations of trace metals relative to the surrounding sediments (in particular Fe, Mn, Al, S, Ba, V, Mo, As, I) [16, 20, 45, 21]. Enrichment of redoxsensitive elements such as barium is considered the best indicator of sapropels as confirmed by the good correlation existing between this element and the organic carbon. This led to assume that Ba, present as biogenic barite, is a good paleoproductivity proxy and the best indicator of the sediments deposited under anoxic conditions. Another useful geochemical proxy in sapropel studies is a peak in Mn/Al, which usually delineates the post-depositional oxidation front and marks the thickness of the original sapropel. Post-depositional oxidative alteration in the sediment is documented at the top of many S1 intervals. These geochemical alterations modify the TOC profile but can be detected by studying redox-sensitive elements like Fe and Mn. During and after sapropel deposition, oxygen is consumed in situ by oxidation of organic matter. The oxygen-depleted conditions lead to remobilization and upward diffusion of Mn. When the organic carbon is consumed, oxygen diffuses into the sediments from above and oxidizes the reduced species, Mn and Fe oxyhydroxides which then precipitate at the oxidation front, enriching it in these elements.
6
Sapropel S1: case study from the Ionian Sea
Most of the information available concerning the sapropels is derived by high reso661
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Figure 2: Location map of the core ET 99 M11. lution studies performed on the most recent one (S1) deposited during the early Holocene as it can be easily recovered with simple coring equipment used at sea and its exact age determined by accurate radiocarbon dating methods. Multiproxy investigations have been used to identify the precise boundary of sapropel S1 in several cores covering the entire eastern Mediterranean basin demonstrating that the timing of its deposition occurred between 9.8 - 5.7 14C Ka BP (or 10.8- 6.1 Ka cal. BP) [46]. In this paper we discuss, as case study, the results of a sapropel S1 recovered in Core ET99-M11 (Ionian Sea, Lat 36°44’04”N, Long 15°50’94”E, 2600 m water depth) (Figure 2). This site was chosen because it is close to one of the present day sources of deep-
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water formation for the Eastern Mediterranean Sea, the Adriatic Sea, and it is close to the ODP Hole 964 where more than 50 sapropel layers were recovered in an excellent and complete sediment section spanning the interval from lower Pliocene to the Holocene [47]. Sapropel S1 is identified by about 38 cm of brown muddy sediment visually different from the surrounding sediments (Figure 3). The TOC content reaches a maximum value of 1.6% well distinct from the 0.2-0.3% background values. The chronological framework is constrained by three 14C AMS datings obtained from planktonic foraminifera. A multiproxy investigation carried out at high resolution (1cm sampling corresponding to about 100 years in time resolution) and including geo-
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chemical, planktonic foraminifer assemblage and rock-magnetic analyses indicates that three different environmental scenarios related to changes in anoxia, productivity and seasonal stratifications existed during the sapropel deposition as already proposed by Rolhing et al. [48] for the Adriatic Sea and observed in other cores from the eastern Mediterranean. 1. S1a sub-unit (from 9.8 to 8.2 Ka BP calibrated age). A clear and drastic change of several parameters marks the beginning of the sapropel deposition at about 9.8 ka. The interval is characterized by the strongest anoxic conditions as indicated by peak values in the Barium, Ba/Al ratio and TOC content, minima in the magnetic concentration (low K and ARM values) coupled with increasing magnetic grain size (low Karm/K) and coercivity (Figure 3). This is related to magnetite dissolution and reductive diagenesis of magnetic minerals as consequence of suboxic/anoxic conditions occurring during the sapropel deposition. The planktonic foraminifer assemblage is characterized by elevated percentages of the low salinity water indicator Globigerinoides ruber var. rosa and by the increase in frequency of spinose species typical of tropical and subtropical areas such as Globigerinella siphonifera, Globigerinella digitata, Globoturborotalita rubescens, Globoturborotalita tenella, and Globigerinoides trilobus suggesting warmest surface water conditions with low salinity lenses. The peak abundance of Globigerina bulloides opportunistic species, thriving in any eutrophic setting [49], documents that sapropel formation coincided with a marked increase in nutrient availability in the surface waters.
2. Sapropel interruption (8.2-7.9 Ka BP calibrated age). Both TOC and Ba content decrease for an interval of few centuries, at around 8 Ka BP, marking the interruption of Sapropel S1. This interval is characterized by the peak in frequency of Globorotalia inflata and Neogloboquadrina pachyderma species living in cool and well mixed layer with intermediate to high nutrient levels [49]. Their occurrences represent a short period of improved deep water oxygenation probably triggered by cold conditions synchronous to the 8.2 Ka event [48]. Magnetic parameters also indicate a decreasing dissolution (increasing Karm and Karm/K values) related to lower anoxic conditions (Figure 3). This interruption is characterized by colder water conditions coupled to fairly high productivity. 3. S1b sub-unit (7.9- 5.9 Ka BP calibrated age). The reestablishment of anoxic conditions is well defined by increasing TOC and Ba values (Figure 3). The gradual decrease of G. ruber together with the increase of N. pachyderma dextral and G. inflata (Figure 3) marks a significant change in the upper water column. In particular the distributional trend of the last species records the development of frontal systems in the surface/sub-surface water, leading to the demise of the stratification before the end of sapropel deposition. Magnetic parameters indicate low magnetic content with increasing grain size in the first part of this interval whereas an opposite trend related to the precipitation of Fe oxides is observed at the end of the sapropel layer. The timing of the sapropel boundaries is well defined by several proxies (TOC, Ba, magnetic parameters, planktonic 663
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Figure 3: Magnetic, Geochemical and Micropaleonttological record of sapropel S1 from core ET99-M11. ? symbol refers to 14 C dating points. Karm and Karm/K give indication of magnetic concentration and grain size respectively. Bariun and Ba/Al ratio are among the best indicator of Sapropels reflecting precipitation associated with primary productivity. Mass-specific magnetic susceptibility (x) was measured on discrete samples whereas Karm and Karm/K represent whole-core (U-channel) measurements. In the planktonic assemblages, the relative abundances of the deep-dwelling taxa N. pachyderma and G. inflata and the percentages of the warm subtropical species G. ruber allow to identify the paleoenvironmental sub-units of S1 layer (for details see text). foraminifera) indicating that the anoxic conditions started about 9.8 calibrated Ka BP. Minimum content of magnetic particles (low susceptibility and ARM) with large grain-size (lower Karm/K) indicate that reductive dissolution (i.e. anoxic conditions) reached a peak around 9.1 Ka BP. For an interval of about 3-4 centuries centered around 8.2 Ka BP the anoxic deposition was interrupted by a re-oxygenation phase. This is evident by a decrease in TOC and Ba content, an increase in magnetic content and in planktic foraminifer microfauna indicating deep seasonal mixed layer. After the interruption, the environmental conditions show a second interval characterized by anoxic conditions that appear less developed than those in sub unit S1a. At about 5.8 Ka BP, TOC and Ba
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return to values close to the background indicating the end of the sapropel. Biological proxies point out to a water mass which began to mix earlier than showed by end of the lithological sapropel. Magnetic parameters show an increase in fine grained minerals as an effect of precipitation of Fe-oxides at the top of the oxidation front.
7
Conclusions
In conclusion, our high resolution multiproxy analysis of sapropel S1 from the Ionian Sea shows similar features to synchronous sapropels recovered from the eastern Mediterranean sea. A first phase of bottom water anoxia and sea surface high productivity (due to increase freshwater discharge from the continents) is fol-
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lowed by a short interruption in the anoxic conditions provoked by mixing of upper and intermediate water, which still sustain fairly high productivity. The third phase of sapropel S1 deposition shows that upper water stratification was generally less intense than in the first phase and that the complete demise of stratification occurred before the lithological evidence for the end
of the sapropel S1. This study demonstrates that a combination of micropaleontological, magnetic and geochemical data is a good strategy to reconstruct the relative role of the factors (productivity, bottom waters ventilation, preservation) leading to sapropels formation.
References [1] R.B. Kidd, M.B. Cita, and W.B.F. Ryan. Stratigraphy of eastern Mediterranean sapropel sequences recovered during DSDP Leg 42A and their paleoenvironmental significance. Initial Reports of the Deep Sea Drilling Project, 42:421–444, 1978. [2] M. Rossignol-Strick. Mediterranean Quaternary sapropels, an immediate response of the African monsoon to variation of insolation. Palaeogeography, Palaeoclimatology, Palaeoecology, 49:237 – 263, 1985. ¨ [3] H. Potoni´e. Uber Faulschlamm-(Sapropel)-Gesteine. Sitz. Gesell. nat. F. Berlin, pages 243–245, 1904. [4] B. Kullenberg. On the salinity of the water contained in marine sediments. Meddelanden fran Oceanografiska institutet Goteborg, 21:1–37, 1952. [5] E. Olausson. Studies of deep-sea cores: Reports of the Swedish Deep-Sea Expedition. 8:323–438, 1947. [6] A. Cramp and G. O’Sullivan. Neogene sapropels in the Mediterranean: a review. Marine Geology, 153:11–28, 1999. [7] N. Pinardi and E. Masetti. Variability of the large scale general circulation of the Mediterranean Sea from observations and modelling: a review. Palaeogeography, Palaeoclimatology, Palaeoecology, 158:153–173, 2000. [8] M.B. Cita, C. Vergnaud-Grazzini, C. Robert, H. Chamley, N. Ciaranfi, and S. D’Onofrio. Paleoclimatic record of a long deep sea core from the eastern Mediterranean. Quaternary Research, 8:205–235, 1977. [9] R.C. Thunell, D.F. Williams, and J.P. Kennett. Late Quaternary paleoclimatology, stratigraphy and sapropel history in eastern Mediterranean deep-sea sediments. Marine Micropaleontology, 2:371–388, 1977. [10] C. Vergnaud-Grazzini, W.B.F. Ryan, and M.B. Cita. Stable isotope fractionation, climate change and episodic stagnation in the eastern Mediterranean Pleistocene records. Marine Micropaleontology, 10:35–69, 1977. 665
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[11] A.E.S. Kemp, R.B. Pearce, J. Pike, and J.E.A. Marshall. Microfabric and microcompositional studies of Pliocene and Quaternary sapropels from the eastern Mediterranean. Proc. Ocean Drill. Program Sci. Results, 160:333–348, 1998. [12] A.E.S. Kemp, R.B. Pearce, J. Pike, I. Koizumi, and S.J. Rance. The role of matforming diatoms in the formation of Mediterranean sapropels. Nature, 398:57–61, 1999. [13] M. Rossignol-Strick andW. Nesterof, P. Olive, and C. Vergnaud-Grazzini. Mediterranean stagnation and sapropel formation. Nature, 295:105–110, 1982. [14] H. Schrader and A. Matherne. Sapropel formation in the eastern Mediterranean Sea: Evidence from preserved opal assemblages. Micropaleontology, 27:191–203, 1981. [15] G.J. De Lange and H.L. Ten Haven. Recent sapropel formation in the eastern Mediterranean. Nature, 305:797–798, 1983. [16] S.E. Calvert. Geochemistry of Pleistocene sapropels and associated sediments from the Eastern Mediterranean. Oceanologica Acta, 6:255–267, 1983. [17] E.J. Rohling and W.C. Gieskes. Late Quaternary changes in Mediterranean intermediate water density and formation rate. Paleoceanography, 4:531 – 545, 1989. [18] T.F. Pedersen and S.E. Calvert. Anoxia vs. productivity: What controls the formation of organic-carbon rich sediments and sedimentary rocks? AAPG Bulletin, 74(4):454–466, 1990. [19] S.E. Calvert and T. Pederson. Organic carbon accumulation and preservation in marine sediments: how important is anoxia? In: Whelan, J.K., Farrington, J.W. (Eds.) Productivity, Accumulation and Preservation of Organic Matter in Recent and Ancient Sediments. pages 231–263, 1992. [20] H.L. Ten Haven, J.W. De Leeuw, P.A. Schenk, and G.T. Klaver. Geochemistry of Mediterranean sediments. Bromine/organic carbon and uranium/organic carbon ratios as indicators for different sources of input and post-depositional oxidation, respectively. Organic Geochemistry, 13:255–261, 1987. [21] D. Mercone, J. Thomson, I.W. Croudace, G. Siani, M. Paterne, and S. Troelstra. Duration of S1, the most recent sapropel in the eastern Mediterranean Sea, as indicated by accelerator mass spectrometry radiocarbon and geochemical evidence. Paleoceanography, 15:336–347, 2000. [22] P.J.M. Van Santvoort, G.J. De Lange, C.G. Langereis, M.J. Dekkers, and M. Paterne. Geochemical and paleomagnetic evidence for the occurrence of “missing” sapropels in eastern Mediterranean sediments. Paleoceanography, 12:773–786, 1997.
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[23] A.P. Roberts, J.S. Stoner, and C. Richter. Diagenetic magnetic enhancement of sapropels from the eastern Meditteranean Sea. Marine Geology, 153:103–116, 1999. [24] L. Vigliotti, L. Capotondi, and M. Torii. Magnetic properties of sediments deposited in suboxic-anoxic environments: relationships with biological and geochemical proxies. Paleomagnetism and Diagenesis in Sediments, 151:71–83, 1999. [25] J.C. Larrasoa˜na, A.P. Roberts, J.S. Stoner, C. Richter, and R. Wehausen. A new proxy for bottom-water ventilation in the eastern Mediterranean based on diagenetically controlled magnetic properties of sapropel-bearing sediments. Palaeogeography, Palaeoclimatology, Palaeoecology, 190:221–242, 2003. [26] L. Capotondi and L. Vigliotti. Magnetic and microfaunistical characterization of late Quaternary sediments in the Western Mediterranean (ODP Leg 161). Inference on sapropel formation and paleoceanographic evolution. Proceedings of the Ocean Drilling Program, Scientific Results, 161:505–518, 1999. [27] F.J. Jorissen. Benthic foraminiferal successions across Late Quaternary Mediterranean sapropels. Marine Geology, 153:91–101, 1999. [28] A. Murat. Pliocene Pleistocene occurrence of sapropels in the western Mediterranean Sea and their relation to eastern Mediterranean sapropel S1. Proceedings of the Ocean Drilling Program, Scientific Results, 161:519–527, 1999. [29] A. Murat. Enr´egistrement s´edimentaire des pal´eoenvironements Quaternaires en M´editerran´ee Orientale. Ph.D. dissertation, 1991. [30] A. Piva, A. Asioli, R.R. Schneider, F. Trincardi, N. Andersen, E. ColmeneroHidalgo, B. Dennielou, J.-A. Flores, and L. Vigliotti. Climatic cycles as expressed in sediments of the PROMESS1 borehole PRAD1-2, Central Adriatic, for the last 370 ka, part 1: integrated stratigraphy. Geochemistry, Geophysics, Geosystems DOI:10.1029/2009/2007GC001713, 9(1), 2008. [31] R. Karlin and S. Levi. Geochemical and sedimentological control of the magnetic properties of hemipelagic sediments. J. Geophys. Res., 90:10373–1039, 1985. [32] B.W. Leslie, D.E. Hammond, W.M. Burleson, and S.P. Lund. Diagenesis in anoxic sediments from the california continental borderland and its influence on iron, sulfur, and magnetite behavior. J. Geophys. Res., 95:4453–4470, 1990. [33] L. Vigliotti. Magnetic properties of light and dark sediment layers from the Japan sea: Diagenetic and paleoclimatic implications. Quaternary Science Review, 16:1093–1114, 1997. [34] S. Giunta, A. Negri, C. Morigi, L. Capotondi, N. Combourieu-Nebout, K.C. Emeis, F. Sangiorgi, and L. Vigliotti. Coccolithophorid ecostratigraphy and multi-proxy paleoceanographic reconstruction in the Southern Adriatic Sea during the last 667
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deglacial time (Core AD91-17). Palaeogeography, Palaeoclimatology, Palaeoecology, 190:39–59, 2003. [35] H.F. Passier, G.J. de Lange, and M.J. Dekkers. Magnetic properties and geochemistry of the active oxidation front and the youngest sapropel in the eastern Mediterranean Sea. Geophysical Journal International, 145(3):604–614, 2001. [36] E.J. Rohling, H.C. De Stigter, C. Vergnaud-Grazzini, and R. Zaalberg. Temporary repopulation by low-oxygen tolerant benthic foraminifera within an upper Pliocene sapropel: evidence for the role of oxygen depletion in the formation of sapropels. Marine Micropaleontology, 22:207–219, 1993. [37] A. Negri, L. Capotondi, and J. Keller. Calcareous nannofossils, planktonic foraminifera and oxygen isotopes in the late Quaternary sapropels of the Ionian Sea. Marine Geology, 157:89–103, 1999. [38] A.W.H. B´e and D.S. Tolderlund. Distribution and ecology of living planktonic Foraminifer in surface waters of the Atlantic and Indian oceans. The Micropaleontology of the oceans, pages 105–149, 1971. [39] R.G. Fairbanks, M. Sverdlove, R. Free, P.H. Wiebe, and A.W.H. B´e. Vertical distribution of living planktonic foraminifera from the Panama basin. Nature, 298:841– 844, 1982. [40] B. Schmuker and R. Schiebel. Planktic foraminifers and hydrography of the eastern and northern Carribbean Sea. Marine Micropaleontology, 46:387–403, 2002. [41] B.K. Sen Gupta and M.L. Machain-Castillo. Benthic foraminifera in oxygen poor habitats. Marine Micropaleontology, 20:183–201, 1993. [42] G. Schmiedl, A. Mitschele, S. Beck, K-C. Emeis, C. Hemleben, H. Schulz, M. Sperling, and S. Weldeab. Benthic foraminiferal record of ecosystem variability in the eastern Mediterranean Sea during times of sapropel S5 and S6 deposition. Palaeogeography, Palaeoclimatology, Palaeoecology, 190:139–164, 2003. [43] L. Capotondi, M.S. Principato, C. Morigi, F. Sangiorgi, P. Maffioli, S. Giunta, A. Negri, and C. Corselli. Foraminiferal variations and stratigraphic implications to the deposition of sapropel S5 in the eastern Mediterranean. Palaeogeography, Palaeoclimatology, Palaeoecology, 235:48–65, 2006. [44] C. Morigi. Benthic environmental changes in the Eastern Mediterranean Sea during sapropel S5 deposition. Palaeogeography, Palaeoclimatology, Palaeoecology, 273:258–271, 2009. [45] J. Thomson, D. Mercone G.J. De Lange, and P.J.M. van Santvoort. Review of recent advances in the interpretation of eastern Mediterranean sapropel S1 from geochemical evidence. Marine Geology, 153:77–89, 1999.
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[46] G.J. De Lange, J. Thomson, A. Reitz, C.P. Slomp, M.S. Principato, E. Erba, and C. Corselli. Synchronous basin-wide formation and redox-controlled preservation of a Mediterranean sapropel. Nature Geoscience, 1:606–610, 2008. [47] K.C. Emeis, , and Leg 160 Shipboard Scientific Party. Paleoceanography and sapropel introduction. Proceedings of the Ocean Drilling Program, Initial Reports, 160:21–28, 1996. [48] E.J. Rohling, F.J. Jorissen, and H.C. De Stigter. 200 Year interruption of Holocene sapropel formation in the Adriatic Sea. Journal of Micropaleontology, 16:97–108, 1997. [49] C. Pujol and C. Vergnaud-Grazzini. Distribution patterns of live planktic foraminifers as related to regional hydrography and productive systems of the Mediterranean Sea. Marine Micropaleontology, 25:187–217, 1995.
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Water in Mid Ocean Ridge Basalts: Some Like it Hot, Some Like it Cold M. Ligi1 , E. Bonatti1,2 , D. Brunelli1,3 , A. Cipriani2,3 , L. Ottolini4 1, Institute of Marine Sciences, CNR, Bologna, Italy 2, Lamont-Doherty Earth Observatory, Columbia University, Palisades, New York, U.S.A. 3, Department of Earth Sciences, University of Modena, Modena, Italy 4, Institute of Geosciences and Earth Resources, CNR, Pavia, Italy marco.ligi@bo.ismar.cnr.it Abstract The presence in the Earth’s mantle of even small amounts of water and other volatiles has major effects: first, it lowers drastically mantle’s viscosity, thereby facilitating convection and plate tectonics; second, it lowers the melting temperature of the rising mantle affecting the formation of the oceanic crust. H2 O concentration in oceanic basalts stays below 0.2 wt% except for basalts sampled near “hot spots” that contain significantly more H2 O than normal MORB, implying that their mantle plume sources are unusually H2 O-rich. Basalts sampled in the Equatorial Atlantic close to the Romanche transform, a thermal minimum in the Ridge system, have a H2 O content that increases as the ridge is cooled approaching the transform offset. These basalts are Na-rich, being generated by low degrees of melting of the mantle, and contain unusually high ratios of light versus heavy rare earth elements implying the presence of garnet in the melting region. H2 O enrichment is due not to an unusually H2 O-rich mantle source, but to a low extent of melting of the upwelling mantle, confined to a deep wet melting region. Numerical models predict that this wet melting process takes place mostly in the mantle zone of stability of garnet. This prediction is verified by the geochemistry of our basalts showing that garnet must indeed have been present in their mantle source. Thus, oceanic basalts are H2 O-rich not only near “hot spots”, but also at “cold spots”.
1
Introduction
The distribution of water in the interior of our Planet, particularly in the Earth’s mantle, has become in the last few years an important objective of the Earth Sciences community. The presence in the mantle of even small amounts of water and other volatiles affects viscosity and partial melting of the mantle that rises below mid ocean ridges, lowering the peridotite solidus and permitting greater melt pro-
duction at lower mean extents of melting. Melting of the mantle that upwells beneath spreading centers induces significant density and viscosity changes, triggering dynamic upwelling by buoyancy of retained melt and by reduced density of depleted mantle. The H2 O content of the oceanic upper mantle can be estimated from the H2 O concentration in mid ocean ridge basalt (MORB) glasses correcting for the effects of degassing and magmatic differentiation.
Marine Geology
Figure 1: Distribution of Na8 and (H2 O)8 in MORB glasses along the axis of the MidAtlantic Ridge from Iceland to the Equator. Data are from Table 1, our unpublished results and the Petrological Database of the Ocean Floor (PETDB) of Lamont Doherty Earth Observatory. The H2 O content of normal MORB (NMORB) is generally below 0.2 wt% [1]. Given that H2 O is about as incompatible as Ce, and assuming â&#x2C6;ź10% average degree of melting of the mantle upwelling below mid ocean ridges (MOR), the mantle source of N-MORB is assumed to contain 0.01 to 0.02 wt% H2 O [2]. However, basalts from topographically swollen portions of MOR have H2 O concentrations higher than N-MORB (Figure 1). These swollen ridges are generally interpreted as being influenced by hot plumes rising from the transition zone or even deeper in the mantle. Thus, the H2 O content of the mantle source of plume-type oceanic basalts is probably significantly higher than that of the N-MORB source region. For example,
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the mantle source of the Icelandic [3] and Azores platform [4] crust contain between 620 and 920 ppm, i.e., several times higher than the N-MORB source. Concerning offridge hot-spots, a 405 Âą 190 ppm H2 O content has been estimated for the mantle source of Hawaiian basalts [5] supporting the hypothesis that plume-type mantle is H2 O-rich relative to the N-MORB mantle source. High water and volatile contents lower the mantle solidus, so that the mantle melts deeper and to a higher degree during its ascent below MOR. We report here that the H2 O content of basaltic glasses from the equatorial Mid Atlantic Ridge (MAR) is significantly higher than that of N-MORB [6]. However, these H2 O-rich basalts are associated
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not with a “hot” portion of MOR, but with the opposite, i.e. a thermal minimum in the ridge system. We will discuss a model that explains why the H2 O content of oceanic basalt is high not only at “hot spots”, but also at “cold spots”.
2
Methods
Basaltic glasses were sampled by dredging at several sites along the ∼220 km long MAR segment (eastern Romanche Ridge segment or ERRS) that extends south of the 900 km Romanche transform (Figures 2 and 3). Glasses were selected for freshness and analyzed for Rare Earth element concentrations (REE), H2 O and for major elements. Major elements were determined using a JEOL JXA 8600 microprobe at IGG-CNR, in Florence. The acceleration voltage was 15 kV, the sample current was 10 nA. The counting times were 40 s for Na and Cl and 10 s for all other elements, the spot size was 10 µm. H2 O content was determined at IGG-CNR in Pavia by secondary ion mass spectrometry (SIMS) with a Cameca IMS 4f ion microprobe following a procedure which involves “energy filtered” secondary ions [10] with emission energies in the range 75-125 eV. Under these experimental conditions the H background, measured on a sample of quartz, is typically 0.009 wt% H2 O. The values for H2 O in the Table 1 are the average of 3 measurements. The accuracy of analysis is estimated to be 10% relative. REE were determined with the Pavia ion microprobe. An optimised energy filtering technique [11] was applied to remove complex molecular interferences in the secondary ion mass spectrum. Light REE-rich basalts were analysed applying a deconvolution filter to the secondary-ion REE mass spec-
trum in order to reduce residual oxide interferences (i.e., BaO on Eu, CeO, NdO on Gd, GdO on Yb, and EuO on Er). Precision of the measure is on the order of 10 % relative, for REE concentrations in the range 0.1-0.7 ppm. Below 0.1 ppm precision is mainly limited by (Poisson) counting statistics and falls to ∼30 % relative. Accuracy is on the same order of precision. The experimental conditions involved a 9.5 nA, 16O- primary ion beam accelerated through -12.5 kV and focused into a spot 10-15 µm in diameter, and energy-filtered (75-125 eV) positive secondary ions detected under an ion image field of 25 µm. .
3
The Melting Model
We carried out numerical experiments to estimate the extent to which the upper mantle is cooled by a long-offset, lowslip transform, such as the Romanche. The temperature field before the onset of melting has been calculated by the steady-state advection-diffusion equation (1), where k=mantle thermal diffusivity, 8.04 10−7 m2 · s−1 ; vs =matrix velocity vector; a=adiabatic temperature gradient, 0.0003 °C/m and z=unit vector along zaxis. Mantle temperatures have been computed through a 3D-domain of mantle flow calculations, by the over-relaxation upwind finite difference method described by Morgan and Forsyth [12]), using a variable grid spacing (512x256x101) with the highest grid resolution (1 km) in the proximity of the plate boundaries. Temperature solutions were found assuming constant temperature at the surface (0 °C) and different mantle potential temperatures at 150 km depth in order to evaluate the Equato673
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Figure 2: Passive flow model geometry. Base of rigid plates represents the upper boundary layer in our plate thickening passive mantle flow model. It was obtained iteratively solving each time the mantle temperature field, starting from a constant-thickness plateflow model. The computed thickness of the African plate lithosphere at the ERRS transform intersection is ∼50 km. rial MAR cold spot effect. Melt parameters of the model (crustal thickness and maximum degree of melting) were inferred from the chemistry of basalts sampled from the centre of the ERRS [8]. A 1330 °C mantle potential temperature has been assumed at a depth of 150 km beneath the ERRS, 50°C colder than ”normal” (temperature that produces 6 km of crust, assuming that all melt is extracted). Mantle flow velocities were estimated in equation (1) assuming steady-state plate thickening passive flow [13, 14]. We modeled the corner flow induced by seafloor spreading in a computational frame 2048x1024 km wide and 150 km deep (2x2 km spaced grid points for each 1 km depth increment) assuming an incompressible, homogeneous, isoviscous mantle 674
beneath a Romanche-like ridge-transformridge plate boundary geometry (offset of 900 km) with a spreading rate of 16 mm/yr. We solved for the steady-state three-dimensional passive mantle flow via a Fourier pseudo-spectral technique [15]. The base of rigid plates, assumed to correspond to the depth of 700 °C isotherm, was obtained iteratively solving each time the mantle temperature field, starting from a constant-thickness plate-flow model. A 512 km long ridge segment, longer than the ERRS, was chosen in order to evaluate how far the transform effect extends along axis, avoiding numerical edge effects. Water in the upper mantle plays an important role in governing melt generation beneath spreading centres [16]. The amount of water present in the oceanic upper man-
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Table 1: Average chemical parameters obtained from basaltic glasses sampled along the ERRS. Samples obtained during cruises S-13, S-16 (RV Akademik N. Strakhov) and G96 (RV Gelendzik), (this work), and from cruise RC2086 (RV Conrad, Schilling et al. [8]*, trace element compositions from Hannigan et al. [9]). tle is sufficient to deepen the peridotite solidus [17] and cause partial melting in a region wider and deeper than that expected for an anhydrous mantle [18, 19]. We modeled melt generation, including the effect of water on the peridotite solidus using a modified form of a recent parameterization of experimental data developed by Katz et al. [20], adding a pressuredependent water bulk distribution coefficient and near-fractional melting regimes. The total amount of melting F, as defined in equation (2), can be expressed as a function of pressure P, temperature T and weight fraction of water dissolved in the melt, where Ts and Tl are the temperatures of the lherzolite solidus and liquidus, respectively. Equation (3) defines the decrease in the solidus temperature caused by a water content XH2 O in the melt, given an initial concentration in the source of XH2 O bulk and β, γ and κ are experimentally constrained parameters with values of 1.5, 0.75
and 43 °C wt%−g , respectively. We assume that peridotite major phases (such as cpx) are never exhausted from the residue, given the low predicted maximum degree of melting (<20%). Release of latent heat of fusion by freezing of melt and hydrotermal cooling have been neglected. Batch and near-fractional melting are assumed and simulated by mapping the melting interval from the batch melting experiments; water is treated as an incompatible component with a bulk distribution coefficient DH2 O that varies with pressure [16]. Thus, the instantaneous dissolved water fraction in the melt is estimated in the case of batch melting and of near-fractional melting model by equations (4a) and (4b), respectively; where equation (5) is the effective bulk partition coefficient with f0 melt retained. Because ∆TH2 O depends on the melt fraction F, in order to solve the equation (2) we use the Newton-Raphson iterative method, i.e. equation (6).
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Figure 3: Multibeam topography of the eastern Romanche ridge-transform intersection and predicted melt production beneath the ERRS. a) Shaded relief image based on multibeam data. Depth ranges from 7800 m (dark blue) to 1000 m (light grey). Spreading direction and small circle path (thick red solid line) have been computed using the AfricaSouth America Eulerian vector of NUVEL-1A model [7]. Sample locations (Table 1) are indicated by red circles (this work) and black triangles. b) Fraction of melt generated along the ERRS axis, including the effect of water on peridotite solidus. White thick dashed line marks the region of dry melting. Solid thick purple line marks the upper boundary of the region of melt production, i.e. where production rate is positive. Red thick dashed lines indicate boundaries between garnet and spinel stability fields. Isotherms are indicated by thin red lines. c) and d) Across axis sections showing fraction of melt generated at the ERRS centre (120 km) and in the proximity (50 km) of the ridge-transform intersection, respectively.
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Figure 4: Isobaric melting curves with different bulk water contents and for different melting regimes. We assumed a modal cpx concentration of 17% in the unmelted solid. The discontinuity in melt productivity at high degrees of melting is due to cpx exhaustion. a) Batch melting; b) near-fractional with a residual porosity f=1% and c) pure-fractional at pressures of 1 GPa.; d) Batch melting; e) near-fractional (f=1%) and f) pure-fractional at pressures of 2 GPa. Note, when hydrous melting is included, melting regime affects significantly melt productivity.
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Figure 5: REE patterns, from two different along axis locations, for the near-fractional melting model with different residual porosities. Solid lines connecting small diamonds are partial aggregated melts predicted for vertical increments dz = 5 km. The thick solid red lines connecting big orange diamonds are the mean compositions after complete mixing. Yellow square brackets indicate REE patterns from melt generated in the garnet stability field.
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The value of F can be obtained by equation (7), starting with the trial solution F(0) = 0, where j indicates the j-th iteration and where the derivative is defined in the case of batch and near-fractional melting models by equations (8a) and (8b), respectively. The convergence of such approximation to the desired solution F can be evaluated by the Banach’s contraction mapping theorem: the iteration is terminated when the absolute value of [(Fj+1 - F j )/Fj+1 ] is sufficiently small. Figure 4 shows isobaric melting curves obtained from batch and fractional melting models, at pressures of 1 GPa and 2 GPa, with a constant bulk water distribution coefficient (DH2 O =0.01) and different bulk water contents. Notice that adding water greatly depresses the solidus and produces a prominent “low-F tail” [19]. When water content (0.3 wt% melting curve) exceeds saturation in the melt, which is mostly a function of pressure [20, 21, 22] the melting function sharply increases just above the solidus due to the overabundance of water acting as an additional phase. When major phases, such as cpx, are exhausted from the residue, the productivity decreases discontinuously and then rises again (cpx-out criterium), because melting reactions start to consume principally opx [20]. Note that during batch melting the solid retains incompatible elements, such as water, up to high degrees of melting, thus water affects significantly the maximum extent of melting. In fact, addition of water to a peridotite system increases monotonically the degree of melting at constant temperature and pressure. In contrast, fractional melting determines rapid depletion of water in the residual solid. Melt productivity is low, and only a small percent of the melt fraction is produced before water’s complete exhaustion, when
melting proceeds above a ”dry solidus” with higher production rates [16, 18] reaching the values of dry peridotite [4]. A bulk partition coefficient for water between melt and residue, that decreases during progressive decompression melting because of the drop in pressure and in the modal abundance of pyroxenes [16] yields a sharper wet-to-dry transition than would a constant value of 0.01 of the partition coefficient. We calculated crustal thickness, mean pressure of melting, mean degree of melting, and mean composition of the aggregate melt, at any locations along axis from the centre toward the tip of the ridge segment, for each of the following melting models: wet and dry, batch, near-fractional and pure-fractional. We assumed as mantle mineral assemblages for garnet, spinel, and plagioclase peridotite those of McKenzie and O’Nions [20] and mineral proportions, in the transition zone between 85 and 60 km, varying linearly from pure garnet peridotite to pure spinel peridotite. REE distribution coefficients and source contents are from Hellebrand et al. [23]. The melt production rate at any place (x,y,z) beneath the ridge is given by equation (9). The total volume of melt production, per unit time per unit length of the ridge , at an along axis location y0 , is given by equation (10) and was computed by integrating equation (9), over the cross-sectional area R in which melting occurs (i.e. over the region where production rate is positive). The crustal thickness Hc is obtained from equation (11), where U0 is the sea-floor half-spreading rate, rm and rc are mantle and crustal densities, respectively. The mean pressure of melting P was found by integrating the product of the depth and the melt production rate over the melting region R divided by the total melt production rate as in equation (12). The mean de679
Marine Geology
gree of melting, in equation (13), was calculated by a similar method, following the definition given by Forsyth [24] and adopting the nomenclature proposed by Plank et al. [25]. The composition of the aggregated melts in equation (14), was estimated by integrating the instantaneous composition of the liquids produced at each degree of melting weighted by the melt production rate, where Ci is the mean concentration of aggregated melt ci and it is the instantaneous concentration in the liquid for the i-th element. Assuming constant solid phase proportions entering in the melt, the bulk distribution coefficient Di , between liquid and residual solid can be evaluated by equation (15), where Xj is the fraction of the j-th mineral and dij is the partition coefficient for the i-th element between the j-th phase and liquid. Hence the instantaneous concentration of thei-th element, during near-fractional melting, is given by equation (16a), where Di * is the effective bulk partition coefficient with Ď&#x2020;0 melt retained and during batch melting, assuming that melt and solid move together vertically upward, by equation (16b), where Fmax is the maximum extent of melting at the top of the melting column located at (x, y0 ). Figure 5 shows modeled REE patterns from near-instantaneous melts predicted for vertical increments of 5 km at two different along axis locations. The increasing influence of garnet in the aggregate melt, as the ridge-transform intersection is approached, is clearly displayed.
4
Water-rich basalts
The H2 O content of ERRS glasses ranges from 0.25 to 1.10 wt% (Table 1). Their depth of eruption is greater than 2100 m below sea level; thus, they must be un680
dersaturated in H2 O at these depths [21] with little or no H2 O loss during eruption. The absence of vesicles in the glasses supports this conclusion. Their Cl content is <0.11 wt% and mostly below the limit of detection (0.04 wt%), suggesting no contamination by sea water. In order to correct for the effects of differentiation, we calculated (H2 O)8 , i.e., H2 O normalized to 8 wt% MgO [27, 28], assuming olivine-plagioclase-clinopyroxene fractionation. Fractionation of incompatibles such as H2 O, is better described by power-law equations. We used the equation of Taylor and Martinez [28] to correct H2 O at a common 8 wt% MgO. To limit possible errors introduced by this correction, we used only analyses within the range of 5.5-8.5 wt% MgO. The correction lowers somewhat the H2 O values but does not affect relative trends (Table 1). We used the equation of Plank and Langmuir [26] to correct Na2 O for all the samples, facilitating comparisons of our ERRS data with MORB data. Figure 6 shows FeO versus MgO for all ERRS samples. The data do not show clear fractionation trends, suggesting more than one parental magma or liquid-line-of-descent (LLD). In calculating the corrected FeO values we considered the different quality of the fractionation trends for different sites to account for varying liquid-lines of descent and varying compositions of fractionating FeO. We calculated patterns of LLD for all sites where more than ten analyses are available (Figure 6b), assuming olivine-clinopyroxeneplagioclase fractionation and using the program PETROLOG by Danyushevsky [1], that includes the effects of H2 O on olivine liquidus temperature. From an inspection of the computed LLD patterns, we derived an expression similar to that of Plank and Langmuir [26] for Na2 O, where the frac-
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tionation slope increaseas as Fe8 increases. In evaluating the effectiveness of the estimated local correction, we applied to the data fractionation corrections taken from the literature [27, 28]. Although average values and standard deviation are greatly affected by the fractionation correction we used, the along axis trend is approximately the same. We decide to adopt the expression of Taylor and Martinez [28] due to the small differences with our estimated local correction and to allow comparisons with other water-rich basalts. (H2 O)8 and Na8 plotted versus latitude along the MAR axis (Figure 1) reveal variations of basalt water content, with maxima in regions where the MAR is affected by mantle plumes, such as at 62º-70º N (Iceland), 35º-45º N (Azores), and at 15º20’ N. Maxima in H2 O content are generally mirrored by minima in Na8 (Figure 1), consistent with the idea whereby plumerelated high degree of melting and waterrich plume mantle source go together [29, 30]. Glasses from the equatorial MAR are an exception to this pattern, in so far as they are H2 O-rich while Na8 is also high. High Na8 basalts are consistent with a low degree of melting in this region [31, 8]. Peridotite mineral composition also suggests that the mantle in the equatorial MAR underwent exceptionally low (<5%) degrees of melting [32], probably due to the combined effect of a regional equatorial Atlantic thermal minimum [31, 29], and of a strong “transform cold-edge effect” [33], that cools the ridge as it approaches old/thick/cold lithosphere across transform offsets. We carried out numerical experiments to estimate the extent to which the upper mantle is cooled by a long-offset, lowslip transform, such as the Romanche, and how partial melting and H2 O distribution
are affected. We assumed a 900 km long transform offset, a half spreading rate of 16 mm/yr and the base of rigid plates determined by the 700 °C isotherm (Figure 2). The ridge segment impacting on the transform was assumed to be 512 km long, longer than the real ERRS, in order to evaluate how far the transform effect extends along axis, avoiding numerical edge effects. Assuming a ∼175 ppm H2 O content in the upper mantle [16] the peridotite solidus is lowered causing partial melting in a subridge mantle region wider and deeper than would be if the mantle were dry [18, 19]. We modeled melt generation, including the effect of H2 O on the peridotite solidus, batch and nearfractional melting are assumed and simulated by mapping the melting interval from the batch melting experiments, water is treated as an incompatible component with a bulk distribution coefficient DH2 O that varies with pressure [16]. Our numerical results (Figure 3) show a strong decrease of ”crustal production” as the ridge approaches the transform, and no melting at all in a 20-40 km wide strip close to the fracture zone (Figure 3b), in agreement with the observation that the basaltic crust is nearly absent in that strip [32]. A cross-ridge sub-triangular melting region, ∼600 km wide at its base, is predicted beneath the centre of the segment, where the maximum extent of melting is 16.5% (Figure 3c). The melting region becomes smaller and asymmetric moving toward the RTI, with the maximum degree of melting decreasing rapidly (∼8% 50 km from the RTI), and the initial depth of melting varying greatly across axis (Figure 3d). Water addition deepens the onset of melting to 85 km beneath the centre of the segment, too cold for the anhydrous solidus to encounter garnet peridotite. Thus, water 681
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addition allows a significant melt fraction to be generated in the presence of residual garnet (Figure 3c). The initial depth of melting shoals in the proximity of the RTI, where the partially molten region is mostly due to hydrous melting. The numerical model predicts that basalts sampled close to the Romanche fracture zone are generated exclusively in the sub-ridge “wet melting” mantle interval, i.e., between ∼80 and 60 km deep, within the region of stability of garnet. We would thus expect a significant “garnet signature” in the chemical composition of our basalts, since they are undiluted by melts produced in the “dry melting interval”, above ∼60 km depth within the spinel stability field. REE partition coefficients during melting are different across the 80 to 60 km depth boundary between garnet stability below, and spinel stability above: the heavy REE are compatible with garnet but not with spinel. Thus, melting in the garnet stability field produces liquids depleted of HREE relative to LREE, and with (Sm/Yb)n ratios well over one [34, 35, 36, 37] and increasing with the proportion of melt generated in the garnet stability field. The deeper the level where the ascending mantle stops melting, the higher the proportion of melt generated in the presence of garnet [38]. The concentration of highly incompatible elements in the aggregated liquid should be inversely proportional to the mean degree of melting. Therefore, incompatible elements should be increasingly enriched moving along axis towards the RTI. We calculated crustal thickness, mean pressure of melting, mean degree of melting, and mean composition of the aggregate melt, at any location along axis from the centre toward the tip of the ERRS, for each of the following melting models: wet and dry, batch, near-fractional and pure-fractional. 682
Basalt Na8 , (Sm/Yb)n and H2 O contents increase along axis toward the RTI, as predicted by the numerical model (Figures 7 and 8). Note that when hydrous melting is included, the selected melting regime (batch, pure or near-fractional) affects melting parameter predictions, due to the pressure release melt parameterization adopted. The observed along-axis average patterns of melting parameter chemical indicators, such as Na8 , Fe8 and REE concentrations (Figures 7 and 8), suggest that a pure-fractional or near-fractional melting model with a very low residual porosity (<0.5 %) fits the data best.
5
Conclusions
We conclude that the H2 O content of the oceanic basaltic crust peaks not only close to “hot spots”, but also at “cold spots” along MOR. However, while “hot spot” H2 O maxima are caused by high degrees of melting of their H2 O-rich mantle plume sources, the “cold spots” H2 O enrichment is due to low degrees of melting occurring mostly within the “wet melting” depth interval below the ridge, largely within the garnet stability field, with minor dilution from shallower “dry” melts. Our results are consistent with the ubiquitous presence in the deeper part of the sub ridge melting column of volatiles and enriched components, that are tapped preferentially during incipient “wet” melting, but are normally diluted by more abundant “dry” melts, generated in the shallower part of the melting column. Extrapolating from these results, we expect relatively high H2 O content in basaltic crust generated at other “thermal minima” along the mid Ocean Ridge system, as at the Australian/Antarctic Discordance [39] and at the Gakkel Ridge [40].
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6
Acknowledgements
ported by the Italian Consiglio Nazionale Ricerche, Fondazione Cassa di Risparmio Research sponsored by EURO- di Modena through CARBRIDGE, and CORE/EUROMARGINS Programme Fondazione Onlus Rita Levi-Montalcini. (Project 01-LEC-EMA21F). Work sup-
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Figure 6: FeO? (FeO total) versus MgO. a) Data from the ERRS basaltic glasess are scattered from linear trends, suggesting more than one parental magma or liquid line of descent (LLD). b) Blue diamonds mark data from site S16-26. Orange triangles and solid red line outline computed LLD for compositions from site S16-26. Solid blue line, slope used to calculate Fe8 .
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Figure 7: Along axis relationships between different definitions of mean extent of melting (FV , Forsyth [24] FB , Plank and Langmuir [26]) and the observables (volume and composition of oceanic crust) for the melting model in Figure 3. a) Average Na8 of ERRS basaltic glasses and predicted Na2 O concentrations in the aggregated melt assuming source concentration C0 =0.3%, distribution coefficient D=0.03 [26], and different melting regimes. Near fractional melting curve with residual porosity f<0.5% fits the data best. Error bars indicate standard deviation. b) Maximum degree of melting (FM ax ) and crustal thickness (Hc ). c) Mean extent of melting curves: FB , black and FV , red lines. 685
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Figure 8: Relationships between melt parameters predicted for mid-ocean ridge melting regimes and values obtained from the basalts sampled along the ERRS axis (Table 1). a) Crustal thickness. b) Average degree of melting and Na8 . c) Mean pressure of melting and Fe8 . Note that left and right vertical axes in b and c are different and independent from each other. d) Chondrite normalized ratio (Sm/Yb)n . e) Models of water content in the aggregated melt and observed H2 O concentrations in basaltic glasses. Error bars indicate standard deviation.
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Figure 9: Melting model equations
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References [1] L.V. Danyushevsky. The effect of small amounts of H2 O crystallisation of midocean ridge and backarc basin magmas. J. Volcanol. Geotherm. Res., 110:265–280, 2001. [2] P.J. Michael. The concentration, behavior and storage of H2 O in the suboceanic upper mantle - Implications for mantle metasomatism. Geochim. Cosmochim. Acta, 52:555–566, 1988. [3] A.R.L. Nichols, M.R. Carroll, and A. Hoskuldsson. Is the Iceland hot spot also wet? Evidence from the water contents of undergassed submarine and subglacial pillow basalts. Earth Planet. Sci. Lett., 202:77–87, 2002. [4] P.D. Asimow, J.E. Dixon, and C.H. Langmuir. A hydrous melting and fractionation model for mid-ocean ridge basalts: Application to the Mid-Atlantic Ridge near the Azores. Geochem. Geophys. Geosyst., 5:Q01E16, 2004. [5] J.E. Dixon and D.A. Clague. Volatiles in basaltic glasses from Loihi Seamount, Hawaii; evidence for a relatively dry plume component. J. Petrology, 42:627–654, 2001. [6] M. Ligi, E. Bonatti, A. Cipriani, and L. Ottolini. Water-rich basalts at mid-oceanridge cold spots. Nature, 434:66–69, 2005. [7] C. DeMets, R.G. Gordon, D.F. Argus, and S. Stein. Effect of recent revisions to the geomagnetic reversal time scale on estimates of current plate motions. Geophys. Res. Lett., 21:2191–2194, 1994. [8] J.G. Schilling, C. Ruppel, A.N. Davis, B. McCully, et al. Thermal structure of the mantle beneath the Equatorial Mid-Atlantic ridge - Inferences from the spatial variation of dredged basalt glass compositions. J. Geophys. Res., 100:10057–1007, 1995. [9] R.E. Hannigan, A. Basu, and F. Teichmann. Mantle reservoir geochemistry from statistical analysis of ICP-MS trace element data of equatorial mid-Atlantic MORB glasses. Chemical Geology, 175:397–428, 2001. [10] L. Ottolini, P. Bottazzi, A. Zanetti, and R. Vannucci. Determination of hydrogen in silicates by secondary ion mass spectrometry. Analyst, 120:1309–1313, 1995. [11] P. Bottazzi, L. Ottolini, R. Vannucci, and A. Zanetti. An Accurate Procedure for the Quantification of Rare Earth Elements in Silicates. In SIMS IX Proceedings, edited by A. Benninghoven, Y. Nihei, R. Shimizu and H.W. Werner, John Wiley Sons, Chichester (England). pages 927–930, 1994. [12] J. Morgan Phipps and D.W. Forsyth. Three-dimensional flow and temperature perturbations due to a transform offset: Effects on oceanic crustal and upper mantle structure. J. Geophys. Res., 93:2955–2966, 1988. 688
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[13] Y. Shen and D.W. Forsyth. The effects of temperature and pressure dependent viscosity on three-dimensional passive flow of the mantle beneath a ridge-transform system. J. Geophys. Res., 97:19717–1972, 1992. [14] D.K. Blackman and D.W. Forsyth. Mantle Flow and Melt Generation at Mid-Ocean Ridges. Geophysical Monograph, 71:311–326, 1992. [15] M. Ligi, M. Cuffaro, F. Chierici, and A. Calafato. Three-dimensional passive mantle flow beneath mid-ocean ridges: an analytical approach. Geophys. J. Inter., 175(2):783–805, 2008. [16] G. Hirth and D.L. Kohlstedt. Water in the oceanic upper mantle: Implications for rheology, melt extraction and the evolution of the lithosphere. Earth Planet. Sci. Lett., 144:93–108, 1996. [17] D.R. Bell and G.R. Rossman. Water in the Earth’s mantle: The role of nominally anhydrous minerals. Science, 255:1391–1397, 1992. [18] M.G. Braun, G. Hirth, and E.M. Parmentier. The effect of deep damp melting on mantle flow and melt generation beneath mid-ocean ridges. Earth Planet. Sci. Lett., 176:339–356, 2000. [19] P.D. Asimow and C.H. Langmuir. The importance of water to oceanic mantle melting regimes. Nature, 421:815–820, 2003. [20] R.F. Katz, M. Spiegelman, and C.H. Langmuir. A new parameterization of hydrous mantle melting. Geochem. Geophys. Geosyst., 4:1073, 2003. [21] J.E. Dixon and E.M. Stolper. An experimental study of water and carbon dioxide solubilities in mid-ocean ridge basaltic liquids. 2) Applications to degassing. J. Petrology, 36:1633–1646, 1995. [22] B. Mysen and K. Wheleer. Solubility behavior of water in haploandesitic melts at high pressure and temperature. Am. Mineral., 85:1128–1142, 2000. [23] E. Hellebrand, J.E. Snow, P. Hoppe, and A.W. Hofmann. Garnet-field melting and Late-stage Refertilization in ”Residual” Abyssal Peridotites from the Central Indian Ridge. J. Petrology, 43:2305–2338, 2002. [24] D.W. Forsyth. Crustal thickness and the average depth and degree of melting in fractional melting models of passive flow beneath mid-ocean ridges. J. Geophys. Res., 98:16073–1607, 1993. [25] T. Plank, M. Spiegelman, C.H. Langmuir, and D.W. Forsyth. The meaning of “mean F“: claryfying the mean extent of melting at ocean ridges. J. Geophys. Res., 100:15045–1505, 1993. [26] T. Plank and C.H. Langmuir. Effects of the melting regime on the composition of oceanic crust. J. Geophys. Res., 97:19749–1977, 1992. 689
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[27] E.M. Klein and C.H. Langmuir. Global correlations of ocean ridge basalt chemistry with axial depth and crustal thickness. J. Geophys. Res., 92:8089–8115, 1987. [28] B. Taylor and F. Martinez. Back-arc basin basalt systematics. Earth Planet. Sci. Lett., 210:481–497, 2003. [29] J.G. Schilling. Azores mantle blob: The rare-earth evidence. Earth Planet. Sci. Lett., 24:103–105, 1975. [30] E. Bonatti. Not so hot “Hot Spots” in the oceanic mantle. Science, 250:107–111, 1990. [31] E. Bonatti, M. Seyler, and N. Sushevskaya. A cold suboceanic mantle belt at the Earth’s Equator. Science, 261:315–320, 1993. [32] E. Bonatti, D. Brunelli, P. Fabretti, M. Ligi, R.A. Portaro, and M. Seyler. Steadystate creation of crust-free lithosphere at cold spots in mid-ocean ridges. Geology, 29:979–982, 2001. [33] P.J. Fox and D. Gallo. The tectonics of ridge transform intersections. Tectonophysics, 104:204–242, 1984. [34] P. Gast. Trace element fractionations and the origin of tholeitic and alkaline magma types. Geochim. Cosmochim. Acta, 32:1057–1086, 1968. [35] E. Anders and N. Grevesse. Abundances of the elements: Meteoritic and solar. Geochim. Cosmochim. Acta, 53:197–214, 1989. [36] Y. Shen and D.W. Forsyth. Geochemical constraints on initial and final depths of melting beneath mid-ocean ridges. J. Geophys. Res., 100:2211–2237, 1995. [37] E. Hellebrand, J.E. Snow, H.J.B. Dick, and A.W. Hofmann. Coupled major and trace elements as indicators of the extent of melting in mid-ocean-ridge peridotites. Nature, 410:677–681, 2001. [38] R.M. Ellam. Lithospheric thickness as a control on basalt geochemistry. Geology, 20:153–156, 1992. [39] D.M. Christie, B.P. West, D.G. Pyle, and B.B. Hanan. Chaotic topography, mantle flow and mantle migration in the Australian-Antarctic discordance. Nature, 394:637–644, 1998. [40] P.J. Michael, C.H. Langmuir, H.J.B. Dick, J.E. Snow, et al. Magmatic and amagmatic seafloor generation at the ultraslow-spreading Gakkel ridge, Arctic Ocean. Nature, 423:956–961, 2003.
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New High-Resolution Seismic Data off the Campi Flegrei: Insights into the Evolution of the Neapolitan Yellow Tuff (NYT) Caldera, Eastern Tyrrhenian Margin M. Sacchi1 , V. Di Fiore1 , E. Esposito1 , N. Fekete2 , J. Metzen2 , F. Molisso1 , S. Porfido1 , V. Spiess2 , C. Violante1 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, Marum, University of Bremen, Bremen, Germany marco.sacchi@iamc.cnr.it Abstract The Campi Flegrei is an active volcanic district that develops along the Tyrrhenian coast of SW Italy. The area has been active at least since the last 60 ka BP and is structurally dominated by a caldera, 6 km in diameter, associated with the eruption of the Neapolitan Yellow Tuff (NYT), a large ignimbrite unit dated at ca 15 ka BP that covers the area now occupied by the city of Naples, the Campi Flegrei and a large part of the continental shelf off the Pozzuoli Bay. Recent research on the Campi Flegrei has shown that the offshore stratigraphic framework and the volcanic structures off the Pozzuoli Bay are still largely unknown. Even the ages and the basic geometries of the individual structures that compose the offshore caldera system are poorly constrained. In this study we present the preliminary results of the interpretation of a high resolution seismic dataset recently acquired in the Napoli Bay. The seismic grid consists in ca 800 km of multichannel reflection profiles with average distance of ca 150 m between navigation routes in the Pozzuoli Bay. The results of the research include the recognition of the main offshore segments of the ring fault system associated with the development of the NYT caldera, the detection of recent sub-surficial magmatic intrusions, and the evidence of dramatic deformation and uplift of sub-seafloor strata offshore Pozzuoli as an expression of the late stage inner caldera resurgence that occurred over the last 6 kyrs.
1
Introduction
Volcanism at rifted continental margins is among the major geodynamic processes of the Earthâ&#x20AC;&#x2122;s crust. In back-arc basin settings, where rifted margins are often associated with thinned lithosphere, shallow magma reservoirs, and active tecton-
ics, there is a potential for highly explosive activity to occur, often accompanied by large caldera and/or fissure ignimbrite eruptions. The continental margin of Campania is presently one of highest-risk volcanic areas of the world. Growing attention has been paid during the last decades to the study of the volcanological and petrolog-
Marine Geology
ical evolution and the stratigraphic reconstruction of eruptive events associated with the Late Quaternary activity of the Campania volcanic district. The Campi Flegrei is an active caldera located on the coastal zone of SW Italy, close to the city of Naples, that has been characterized by explosive activity and unrest throughout the Late Quaternary [1, 2, 3, 4, 5]. Recent research at Campi Flegrei has shown that a significant part of the offshore volcaniclastic products and structures are still poorly known. The age and geometry of the offshore substructure of the caldera system is poorly constrained, and even the caldera-like diagnostic characters of the district, for long reported in the literature, have been recently questioned [6, 7]. Commonly, large caldera structures experience episodes of unrest characterised by which are vertical ground movements and an increase in shallow seismicity. Possible explanations from various models reveal a remarkable likeliness of a coupled mechanical and thermodynamic effect of heatflow from a relatively shallow magma chamber (a few km below the surface) and the uppermost aquifers. This could strongly modify and amplify ground deformations at calderas [8, 9]. At Campi Flegrei ground level changes could be reconstructed for the last 2.000 (e.g. [10]). Over this period, a series of deformation phases occurred with a total uplift in the order of +15 m. Most authors attribute these events to intra-caldera resurgence dynamics (e.g. [8, 9, 11, 12].
tween the western flank of southern Italy and the eastern Tyrrhenian margin. This area is the result of large-scale Late Neogene to Pleistocene lithospheric extension that accompanied the eastward accretion of the Apenninic fold and thrust belt during the roll-back of the subducting Adria foreland plate [15, 16, 17]. The structural pattern of the Campania continental margin displays Late Quaternary NE- trending normal faults, NW-trending transtensional faults, E-W trending left-lateral faults, with evidence of a general transtensional regime along major E-W left-lateral fault zones [18, 17, 19]. The resulting structural style at the scale of the continental margin is hence characterized by Quaternary half-graben systems bounded by major WSW-ENE listric normal fault with significant sinistral component that tend to merge into detachment surfaces at depth of 5-6 km to 10-12 km [20, 18]. Post-700 ka B.P. tectonic activity has been suggested for NE-SW normal faults within the Campania coastal zone, both onshore [21]) and offshore [18, 19]. Quaternary extension over the Campania continental margin has caused the onset of intense volcanism. Major volcanic centres are the Somma-Vesuvius, Ischia Island and the district of the Campi Flegrei, with its numerous vents both onshore and offshore the Naples Bay. The basin-fill architecture of the Naples Bay is made up by a Middle Pleistocene transgressive-regressive sedimentary cycle which includes interbedded volcanic deposits (likely ignimbrite units), subvolcanic features (e.g. domes, laccoliths), and submarine relics of volcanic edifices at the seafloor (e.g. tuff cones, tuff 2 Geological setting rings). (e.g. [22, 13]). The Campania Plain and the Naples Bay The Campi Flegrei is an active volcanic are integral components of a large Qua- area defined by a quasi-circular depression ternary extensional basin belt located be- that covers some 200 km2 of the coastal 692
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Figure 1: Geologic sketch-map of the Tyrrhenian margin of the Campania Apennines with location of the study area. zone of SW Italy, a large part of which develops off the Naples (Pozzuoli) Bay (Figure 1). The area has been active at least since 60 ka BP [23, 2], and is structurally dominated by a caldera, 6 km in diameter, associated with the eruption of the Neapolitan Yellow Tuff (NYT), a 40 km3 Dense Rock Equivalent (DRE) ignimbrite dated at ca 15 ka BP [24, 25, 26], that covered the district now occupied by the city of Naples, the Campi Flegrei and a large area of the continental shelf off the Pozzuoli Bay. Most authors also recognise an older caldera collapse in the same area that probably occurred after the eruption of the 150
km3 (DRE) Campanian Ignimbrite (CI), 39,000 ka BP. and, consequently, describe the Campi Flegrei as a typical example of a nested caldera system. However the link between Campi Flegrei and this eruption remains largely controversial [2, 6, 7] and the only clear geophysical evidence for a caldera in the Campi Flegrei district is the collapsed area of about 6 km in diameter associated with the NYT unit. Following the NYT event, the evolution of the Campi Flegrei has been dominated by hydromagmatic activity with occasional plinian phases and minor effusive activity forming lava domes [2]. Many large to
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Figure 2: Grid of multichannel reflection seismics and sub-bottom Chirp profiles acquired during oceanographic cruise CAFE 07, Leg 3 (modified after [13]). medium-scale eruptions occurred around 12-10 ky BP and in the past 6,500 years. The last event was the Monte Nuovo eruption in 1538 A.D, which occurred after some 100 years of intra-caldera uplift. Recent magma-related activity clustered in the centre of the caldera is testified by extensive hydrothermalism, accompanied by two very recent episodes (1970-71 and 1982-84) of shallow seismicity and ground/seafloor deformation originating uplift up to 3.5 m in 15 years, with maximum rates of 100 cm/year in the period 1983-1984 [10, 27, 9, 28].
3
Data and methods
This research work relies on the interpretation of multichannel reflection seismic profiles for the understanding of the sequence stratigraphic general architecture and the 694
spatial distribution of volcaniclastic units off the Pozzuoli Bay. Seismic-stratigraphic analysis was mostly conducted on a grid of seismic profiles acquired in 2008 during an oceanographic cruise in the Naples Bay [29, 13] (Figure 2). The grid consisted of ca 150 km of very high-resolution multichannel reflection seismic profiles acquired simultaneously by two systems with different energy source (a 1,7 l GI-Gun Sodera and a 0,4 l mini GI-Gun Sodera) and different hydrophone cables (a 16-channel, 100 m long cable and a 48-channel, 50 m long cable). The sampling rate was set to 0.250 ms, i.e. sampling frequency of 4 kHz. Ship positioning during navigation was determined by a differential GPS. Processing and visualization of the selected seismic data was done with the commercial software packages VISTA 2D/3D Seismic Processing, by Seismic Image Software (GEDCO) and KINGDOM suite.
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Figure 3: Location of multichannel seismic profiles illustrated in this study. Very high-resolution single-channel (subbottom Chirp) data were also used in order to complement seismic stratigraphic investigation on the Late Quaternary sequence, with a focus on the central part of the Pozzuoli Bay. Seismic-stratigraphic interpretation was calibrated on the basis of the sedimentological analysis of gravity cores samples acquired in the area by the IAMCCNR over the last ten years. The sequence stratigraphic method and nomenclature adopted for seismic interpretation is after [14].
4
Interpretation of seismic profiles
The interpreted seismic profiles (Figure 3) reveal a complex stratigraphic and structural setting, dominated by the occurrence of volcanic bodies and siliciclastic depositional units, mostly deriving from the dismantling of the adjacent vents and associated deposits [22, 30, 13]. Volcaniclastic units typically are interbedded with siliciclastic units of the Late Quaternary depositional sequence which represent in terms of chronostratigraphy, the deposits that formed between the onset of the post120 ka BP sea-level fall and the present day. Based on the methods and procedures of sequence stratigraphy [31, 14], 695
Marine Geology
Figure 4: Profile GeoB08-002 (pre-stack) and its interpretation. The profile shows Penta Palummo and Miseno banks that are among the oldest volcanic structures of the area. Sequence stratigraphic nomenclature is after Hunt and Tucker (1992, [14]). TWTT: twoway travel time (ms); Offset: horizontal distance (m). we could identify five main seismic stratigraphic units that from bottom to top are: - Penta Palummo and Miseno Banks unit (> 120 ka BP). This unit, often characterized by the lack of internal reflections occurs in the southern part of the Pozzuoli Bay, in the area of Penta Palummo Bank and is represented by a series of sub-units that can be grouped by stratigraphic position [13]. The unit is covered by the following marine successions of the Forced Regression Wedge Systems Tract (FRWST) and the Lowstand Systems Tract (LST) and has an age older than 120 ka BP (Figures 4, 5 and 6). - Forced Regression Wedge Systems Tract (FRWST) and Lowstand Systems Tract (LST) marine siliciclastic unit (120-18 ka BP). This unit is made up by forced regressive (FRWST) and lowstand (LST) marine deposits of the Late Quaternary depositional sequence (Figures 4, 5 and 6). It is characterized by irregular re696
flections, with evidence of tectonic deformation, shallow magmatic intrusions and intercalation with thick volcaniclastic units, like the Campania Ignimbrite (ca. 40 ka BP). The upper boundary of this unit is represented by an erosional surface often associated with unconformity that truncates the underlying strata and is correlated with the sea-level fall maximum (ca. -120 m) that occurred during the latest glacial Pleistocene (ca. 18 ka BP). - Nisida Bank unit (18-6 ka BP). This unit can be subdivided in a number of sub-units and is mostly represented by the Nisida Bank tuff-cone and associated volcaniclastic deposits, along with sisliciclastic units deriving from the dismantling of the volcanic centres in subaqueous environment [13] (Figure 5). The unit lies on the erosional surface associated with the last glacial maximum and is covered in turn by Late Quaternary
Marine research at CNR
Figure 5: Profile GeoB08-025 (pre-stack) and its interpretation. The profile illustrates a series of volcanic banks in the Pozzuoli Bay along with the Montagna bank (off Posillipo), represented by a volcaniclastic deposit that forms small-scale diapirs within pumiceous material [13]). Sequence stratigraphic nomenclature is after Hunt and Tucker (1992, [14]). TWTT: two-way travel time (ms); Offset: horizontal distance (m). Holocene deposits (18-6 ka BP). - Transgressive Systems Tract (TST) and Highstand Systems Tract (HST) marine siliciclastic units (<18 ka BP) The unit is represented by relatively thin siliciclastic deposits that formed during the rapid rise and highstand of the sea level following the last glacial maximum. The upper part of this unit is clearly involved in deformation and local folding of layers off Punta Pennata as well as off the Pozzuoli coastline (La Starza) (Figures 6 and 7). The lower part of this unit includes the NYT Formation (ca. 15 ka BP) [24, 25, 26] and a volcaniclastic unit, characterized by shallow diapirism, cropping out at the Montagna bank [32, 13]. - Nisida and Cape Miseno unit (<6 k a BP). This unit is represented by the youngest volcanic centres, like the Nisida complex and cape Miseno tuffcone (Refs). The volcaniclastic products
of Nisida overlie the deposits of Nisida Bank and hence are younger than 6 ka BP. They form a complex of sub-units progressively younger towards the North and display mound-like geometries and well-layered deposits on the flanks (Figure 8). The unit also includes shallow magmatic intrusions, like the M. DolcePampano Bank structure (Figures 5 and 9) as well as other volcanic edifices like the Cape Miseno tuff cone and Porto Miseno tuff ring that are characterized by a recent activity, as shown by [33]. The seismic grid also allowed for a detailed interpretation of some deformation structures that occur within the Pozzuoli Bay. Particularly, the seismic profiles clearly illustrate, for the first time, the shallow geometry of the ring fault system of the NYT caldera, and a dramatic deformation of the inner caldera fill that postdates 6 ka BP and can be interpreted as a late-stage inner 697
Marine Geology
Figure 6: Profile GeoB08-033 (pre-stack) and its interpretation. The profile shows the recent deformation and uplift of subseafloor layers associated with the NYT inner-caldera resurgence. Sequence stratigraphic nomenclature is after Hunt and Tucker (1992, [14]). TWTT: two-way travel time (ms); Offset: horizontal distance (m). caldera resurgence structure. Other minor deformation of sub-seafloor layers, like in the area of Punta Pennata represent a local effect due to differential deformation across the rim of the NYT caldera.
fault system associated with the development of the NYT caldera; c) the recognition of a dramatic deformation and uplift of sub-seafloor strata offshore Pozzuoli as an expression of the late stage inner caldera resurgence that occurred over the last 6 kyrs; d) the detection of very recent shallow magmatic intrusions in the M. Dolce5 Conclusion Pampano Bank area; e) the recognition of In the Pozzuoli Bay, the interpreted seis- Nisida as a very young (<6 k a BP) volmic profiles reveal a complex stratigraphic canic unit. and structural setting, dominated by the occurrence of volcanic bodies and siliciclastic depositional units, mostly deriving from the dismantling of the adjacent vents and volcaniclastic units. The most relevant results of this research include: a) a very high-resolution imaging of the Late Quaternaty stratigraphic units of the Pozzuoli bay; b) a detailed description of the geometry of the ring 698
Marine research at CNR
Figure 7: Profile GeoB08-065 (pre-stack) and its interpretation. The profile shows the recent deformation and uplift of the subseafloor strata due to the NYT inner-caldera resurgence. Details of the ring fault system, with associated shallow intrusions and the crest fault system of the caldera resurgence are also shown. Sequence stratigraphic nomenclature is after Hunt and Tucker (1992, [14]). TWTT: two-way travel time (ms); Offset: horizontal distance (m).
Figure 8: Profile GeoB08-090 (pre-stack) and its interpretation. The profile illustrates the volcanic structures of Penta Palummo and Nisida banks, as well as a series of volcaniclastic units (here referred to as Nisida complex), located between Nisida bank and Nisida Island. Sequence stratigraphic nomenclature is after Hunt and Tucker (1992, [14]). TWTT: two-way travel time (ms); Offset: horizontal distance (m).
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Figure 9: Profile GeoB08-108 (pre-stack) and its interpretation. The profile shows a recent shallow magmatic intrusion (M.Dolce-Pampano structure) South of the Ammontatura channel. NYT: Neapolitan Yellow Tuff; TWTT: two-way travel time (ms); Offset: horizontal distance (m).
6
Acknowledgements
We thank Capitan Emanuele Gentile, the officers and the crew of the R/V Urania for their skilled help during the CAFE-07 cruise. We have also appreciated at various stages of the research work the constructive comments by and useful suggestions of
Giovanni De Alteriis. Special thanks are due to Francesco Paolo Buonocunto and Paolo Scotto di Vettimo for their technical and logistic assistance. This research was partly carried out with financial support from ICDP Germany (V. Spiess) and the Research Project “Legge 5” Regione Campania (M. Sacchi).
References [1] P. Di Girolamo, M.R. Ghiara, L. Lirer, R. Munno, G. Rolandi, and D. Stanzione. Vulcanologia e petrologia dei Campi Flegrei. Boll. Soc. Geol. It., 103:349–413, 1984. [2] M. Rosi and A. Sbrana. Phlegraean Fields. 114(9):175, 1987. [3] G. Orsi, S. de Vita, and M. Di Vito.
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nested caldera (Italy): constraints on its evolution and configuration. J. Volcanol. Geotherm. Res., 74:179–214, 1996. [4] G. Orsi, M.A. Di Vito, and R. Isaia. Volcanic hazard assessment at the restless Campi Flegrei caldera. Bull. Volcanol., 66:514–530, 2004. [5] M.A. Di Vito, R. Sulpizio, G. Zanchetta, and M. D’Orazio. The late Pleistocene pyroclastic deposits of the Campanian Plain: new insights into the explosive activity of Neapolitan Volcanoes. J. Volcanol. Geotherm. Res., 17:19–48, 2008. [6] B. De Vivo, G. Rolandi, P.B. Gans, A. Calvert, Bohrson andW.A., F.J. Spera, and H.E. Belkin. New constraints on the pyroclastic eruptive history of the Campanian volcanic Plain (Italy). Mineral. Petrol., 73:47–65, 2001. [7] G. Rolandi, F. Bellucci, M.T. Heizler, H.E. Belkin, and B. De Vivo. Tectonic controls on the genesis of ignimbrites from Campanian Volcanic Zones, southern Italy. Mineralogy and Petrology, 79:3–31, 2003. [8] G. De Natale, F. Pingue, P. Allard, and A. Zollo. Geophysical and geochemical modelling of the 1982-1984 unrest phenomena at Campi Flegrei Caldera, southern Italy. J. Volcanol. Geotherm. Res., 48:199–222, 1991. [9] G. De Natale, C. Troise, and F. Pingue. A mechanical fluiddynamical model for ground movements at Campi Flegrei caldera. Journal of Geodynamics, 32:487– 517, 2001. [10] G. Berrino, G. Corrado, G. Luongo, and B. Toro. Ground deformation and gravity change accompanying the 1982 Pozzuoli uplift. Bull. Volcanol., 47(2):187–200, 1984. [11] G. Orsi, L. Civetta, C. Del Gaudio, S. de Vita, M.A. Di Vito, R. Isaia, S.M. Petrazzuoli, G. Ricciardi, and C. Ricco. Short-term ground deformation and seismicity in the nested Campi Flegrei caldera (Italy): an example of active block-resurgence in a densely populated area. J. Volcanol. Geotherm. Res., 91:415–451, 1999. [12] J. Gottsmann, H. Rymer, and G. Berrino. Caldera unrest at the Campi Flegrei: A critical evaluation of source parameters from geodetic data inversion. Journal of Volcanology and Geothermal Research, 150:132–145, 2006. [13] Sacchi M., Alessio G., Aquino I., Esposito E., Molisso F., Nappi R., Porfido S., and Violante C. Risultati preliminari della campagna oceanografica CAFE 07 - Leg 3 nei Golfi di Napoli e Pozzuoli, Mar Tirreno Orientale. Quaderni di Geofisica, 6:3–26, 2009. [14] D. Hunt and M.E. Tucker. Stranded parasequences and the forced regressive wedge systems tract: deposition during base-level fall. Sedimentary Geology, 81:1–9, 1992.
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[15] A. Malinverno and W.B.F. Ryan. Extension in the Tyrrenhian Sea and shortening in the Apennines as result of arc migration driven by sinking of the lithosphere. Tectonics, pages 227–245, 186. [16] J.S. Oldow, B. D’Argenio, L. Ferranti, G. Pappone, Marsella, E., and M. Sacchi. Large-scale longitudinal extension in the southern Apennines contractional belt, Italy. Geology, 21:1123–1126, 1993. [17] L. Ferranti, J. S. Oldow, and M. Sacchi. Pre-Quaternary orogen-parallel extension in the Southern Apennine belt, Italy. Tectonophysics, 260:325–347, 1996. [18] M. Sacchi, S. Infuso, Marsella, and E. Late Pliocene-Early Pleistocene compressional tectonics in offshore Campania (Eastern Thyrrenian sea). Boll. Geof. Teor. App., 36(141-144):469–482, 1994. [19] A. Milia and M.M. Torrente. Tectonics and stratigraphic architecture of a periTyrrhenian half-graben (Bay of Naples, Italy). 315:301–318, 1999. [20] Mariani, M, and R. e Prato. I bacini neogenici costieri del margine tirrenico: approccio sismo-stratigrafico. Mem. Soc. Geol. Ital., 41:519–531, 1988. [21] G. Gars and M. Lippman. Nouvelle donne´es n´eotectonique dans l’Apennin campanien (Italie du Sud). 298(II-11):495–500, 1984. [22] A. Milia. Stratigrafia, strutture deformative e considerazioni sull’origine delle unit`a deposizionali oloceniche del Golfo di Pozzuoli (Napoli). Boll. Soc. Geol. It., 117:777–787, 1998. [23] C. Cassignol and P. Gillot. Range and effectiveness of unspiked potassium-argon dating: experimental ground work and application. pages 160–179, 1982. [24] C. Scarpati, P. Cole, and A. Perrotta. The Neapolitan Yellow Tuff – A large volume multiphase eruption from Campi Flegrei, Southern Italy. 55:343–356, 193. [25] D. Insinga. Tefrostratigrafia dei depositi tardo-quaternari della fascia costiera campana. page 202, 2003. [26] A.L. Deino, G. Orsi, S. de Vita, and M. Piochi. The age of the Neapolitan Yellow Tuff caldera-forming eruption (Campi Flegrei caldera, Italy) assessed by 40Ar/39Ar dating method. Volcanol. Geotherm. Res., 133:157–170, 2004. [27] J.J. Dvorak and G. Berrino. Recent ground movement and seismic activity in Campi Flegrei, Southern Italy: episodic growth of a resurgent dome. J. Geophys. Res., 96(B2):2309–2323, 1991. [28] M. Battaglia, C. Troise, F. Obrizzo, F. Pingue, and G. De Natale. Evidence for fluid migration as the source of deformation at Campi Flegrei caldera (Italy). Geophysical Research Letters, Vol. 33, 2006.
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[29] J.F. Metzen. Volcanic Features & Sedimentary Interplayin Pozzuoli Bay and offshore Herculaneum, Italy, investigated with high-resolution marine seismics. page 100, 2008. [30] A. Milia, M.M. Torrente, M. Russo, and A. Zuppetta. Late-Quaternary volcanism and transtensional tectonics in the Bay of Naples, Campanian continental margin, Italy. Mineralogy and Petrology, 79:49–65, 2003. [31] H. W. Posamentier and P. R. Vail. Eustatic controls on clastic deposition II – sequence and systems tract models. 42:125–154, 1988. [32] M. Sacchi, B. D’argenio, V. Morra, S. Petrazzuoli, G. Aiello, F. Budillon, G. Sarnacchiaro, and R. Tonielli. Pyroclastic lumps: quick diapiric structures off the Naples Bay. Geophysical Research, Abstracts, 2:2446, 2000. [33] D. Insinga, A.T. Calvert, M.A. Lanphere, V. Morra, A. Perrotta, M. Sacchi, C. Scarpati, J. Saburomaru, and L. Fedele. The Late-Holocene evolution of the Miseno area (south-western Campi Flegrei) as inferred by stratigraphy, petrochemistry and 40Ar/39Ar geochronology. pages 97–124, 2006.
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Geo-biology of Mediterranean Deep-Water Coral Ecosystems M. Taviani1 , L. Angeletti1 , B. Antolini2 , A. Ceregato1 , C. Froglia2 , M. L´opez Correa3 , P. Montagna4,1 , A. Remia1 , F. Trincardi 5 , A. Vertino1,6 1, Institute of Marine Sciences, CNR, Bologna, Italy 2, Institute of Marine Sciences, CNR, Ancona, Italy 3, GeoZentrum Nordbayern,Universit¨at Erlangen-N¨urnberg, Germany 4, Lamont-Doherty Earth Observatory, Columbia University, Palisades, New York, U.S.A. 5, Institute of Marine Sciences, CNR, Venezia, Italy 6, Department of Geology and Geotechnologies, University of Milano-Bicocca, Milano, Italy marco.taviani@bo.ismar.cnr.it Abstract Cold (or deep) water corals (CWC) are a spectacular and widespread component of the ocean biota. Highly structured CWC ecosystems serve as biodiversity hotspots suggesting the need for the proper management of these unique habitats. Among the most relevant actors in promoting the growth of conspicuous structures (deep water reefs and mounds) are branching stony corals like Lophelia pertusa and Madrepora oculata. These widespread frame-builders occur also in the Mediterranean basin and are the backbone of Recent CWC provinces identified thus far in the Ionian Sea (Santa Maria di Leuca), South Adriatic Sea, Strait of Sicily, Catalan-Provenc¸al canyons and Alboran Sea. CWC distribution seem strongly influenced by oceanographic factors coupled with proper substrata, such as steep (canyon heads, seamount walls etc.) or rugged (slumped blocks) topographies. Dead pre-Modern coral assemblages are present in the entire Mediterranean Sea. These submerged fossil assemblages have been mostly U/Th and 14 C dated at the latest Pleistocene documenting propitious basin-wide conditions for their successful growth during glacial periods.
1
Introduction
The existence of cold (or deep) water corals (CWC) is known since centuries (overview in [1]). Their capability to construct substantial bio-constructions (‘reefs’, bioherms, mounds, build-ups), however, has been properly focused mostly in the last decade or so through the implementation of novel technologies to explore the deep ocean and the concomitant launch of a series of international and national multidis-
ciplinary programmes [2, 3, 4, 5, 6, 7]. As evocatively summarized by Freiwald et al. [8], the outcome of such on-going integrated research is that cold-water coral reefs are now ‘out of sight - no longer out of mind’. CWC are generally intended as those deep water azooxanthellate colonial scleractinian (stony) corals more commonly distributed between ca. 200-1200 m in a temperature range of ca. 4-14 °C. CWC group taxa with a pronounced frame-building ability (Lophe-
Marine Geology
lia, Madrepora, Oculina, Goniocorella, Solenosmilia, Enallopsamia, Dendrophylu` lia), some of which, such as Lophelia pertusa, Madrepora oculata and Solenosmilia variabilis, presenting a quasi-cosmopolitan geographic distribution [8]. For instance, in the Eastern Atlantic CWC build-ups constitute one of the major features along the NW European continental margin covering, albeit discontinuously, a distance of over 2400 nautical miles from sub-polar latitudes off Scandinavia to temperate settings off Iberia. Highly structured CWC build-ups are clustered in specific regions where the proper mixture of suitable topographic and oceanographic situations do exist [9, 10]. The most efficient framebuilders (mainly Lophelia with a growth rate attaining 2.5 cm·year−1 : [11]) work in making spectacular carbonate mounds, some encompassing many hundred thousands to millions years of discontinuous coral accretion [9, 12, 13, 14, 4, 5]. CWC mounds and smaller build-ups are documented on both sides of the Atlantic (overview in [5]. Individual giant mounds may reach a few km in diameter and many hundred meters in height and include both active (e.g., Magellan, Darwin, Belgica, Hovland, Porcupine Bank Canyon mounds) and buried examples (e.g., Enya, Viking mounds). In face of their spatiotemporal distribution and volumetric extent, it is evident that CWC are important carbonate producers whose contribution is far from being negligible in the assessment of the carbonate budget [15, 16]. Discrete areas characterized by substantial CWC growth are conveniently recognized as coral provinces whose number is steadily increasing. While one or more such corals often co-act in establishing bio-constructions at depth, a variety of other sessile invertebrates as soli706
tary scleractinians (such as Desmophyllum and Caryophyllia), octocorals, antipatharians (black corals), sponges and even giant oysters (Neopycnodonte), efficiently concur in making CWC habitats true biodiversity hotspots [8, 17, 3, 18, 19, 20, 21]. These coral habitats are significant performers in the economy of the ocean [8, 22, 18] and therefore require proper management and governance [18, 23, 24]. CWC are in fact under aggression by a number of anthropogenic threats [8, 18], above all invasive fishing practices like bottom trawling [25, 26, 27]. The predicted change in seawater chemistry bringing about a progressive acidification of the ocean is also a major concern for the fate of CWC worldwide because of its negative impact on the corals’ calcification ability [28].
2
Cold-water corals of the Mediterranean Sea
In the last 15 years, the Italian National Research Council (CNR) of Bologna has conducted systematic research on CWC in the frame of various national (CNR, FIRBAplabes) and European (ESF Moundforce, EU Hermes and Hermione) projects resulting in a sequence of oceanographic missions covering a substantial part of the Mediterranean Sea, from Alboran to the Levantine basin (Figures 1, 2, 3). This interest for deep sea coral research, however, stems from long before since CWC were routinely encountered during marine geological investigations of the deep Mediterranean Sea carried out beginning in the late 60’s by the former Laboratory then Institute of Marine Geology (now ISMARBologna) resulting in various publications dealing with corals (e.g., [29, 30, 31, 32]).
Marine research at CNR
Figure 1: Map showing the location of CNR scientific cruises devoted to the exploration of cold-water corals in the Mediterranean Sea (1996 to present). Remarkably, CWC have first been reported from the Mediterranean Sea as fouling an electric cable laid between Sardinia and Algeria at depths > 2000 m [33]. Because of its geographic location, the semi-enclosed Mediterranean Sea is inhabited by CWC of strict Atlantic affinity [34, 35, 36]. Although CWC belonging to extant framebuilding genera (i.a., Lophelia, Madrepora, Desmophyllum) are known from this basin since the Miocene at least, CWC morphologically indistinguishable from Recent Lophelia pertusa, Madrepora Oculata and Desmophyllum dianthus (= D. cristagalli) seem to have continuously settled the Mediterranean Sea from the early Pleistocene onwards as documented by fossil assemblages in southern Italy (e.g., [37,
36]), Rhodes [38] and submerged situations (see [36], for a review). The active geodynamic history backing the recent evolution of this basin, results in the exposure of a number of Cenozoic outcrops [39, 40] containing at places deep-sea coral faunas that render the Mediterranean a privileged place where to study major CWC evolutionary and biogeographic patterns [36]. The abundance of still-submerged dead (subfossil) CWC scattered throughout the entire Mediterranean Sea has been considered as an indication that better conditions for the settlement and maintenance of Atlantic-type CWC did exist in the recent past of this basin [30, 35, 36]. Carbon14 and Uranium-series dating reveal that most subfossil CWC are of late Pleistocene
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Marine Geology
Figure 2: Recent Cold-Water Coral Provinces of the Mediterranean Sea. age [32, 41, 42, 43, 44], either predating or postdating the last glacial maximum for which no coral evidence has been produced yet. The picture emerging from outcrop and submerged occurrences is that oceanographic conditions during cold phases of the Quaternary are more propitious for CWC to thrive in the Mediterranean when compared to present times whose significant coral growth (especially Lophelia) seems highly reduced [35, 32]. In the last decade this vision has been re-dimensioned by the discovery of lush CWC communities containing live Lophelia, Madrepora and Desmophyllum due to the accidental findings of scientific fishery operations [45, 46, 47, 48] and the large-scale implementation of modern exploration of the deeper reaches of the Mediterranean through re-
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mote operating vehicles (ROV) and submersibles [17]. Sites of active coral growth (Figure 2) are mainly located in the Ionian Sea off Apulia (Santa Maria di Leuca: [45, 49, 46, 47, 17, 50, 51], the Southern Adriatic [17], the Strait of Sicily (south of Malta, offshore islands and banks: [48, 17], the Catalan-Provenc存al canyon system [52] and Alboran [53]. Limited live coral growth is known in the (still poorly investigated) eastern Mediterranean [54, 32, 44]. With respect to the adjacent Atlantic Ocean, the situation of Mediterranean CWC differs because of the lack of substantial coral carbonate mounds in the latter. Most occurrences refer to CWC fouling or fringing canyon heads, precipitous walls and other hardground substrates [36, 44, 55, 17, 52]. The really most de-
Marine research at CNR
Figure 3: Examples of scientific operations related to CWC research carried out onboard RV Urania. A) Large volume grab sample taken during Cruise CORTI (Tuscan Archipelago, Tyrrhenian Sea, winter 2003); the arrow points subfossil corals embedded in mud (ca -400 m) whose age was later proved to be latest Pleistocene. B) Gravity core operation during MEDCOR cruise (winter 2009) to assess the stratigraphy of CWC south of Malta (Strait of Sicily). C) CTD-Rosette recovery during MEDCOR in the Strait of Sicily to collect oceanographic data on the coral sites, and water samples for geochemical and ocean acidification studies. D) Experiments related to the effects of ocean acidification on live deep water corals conducted during MEDCOR cruise in the Strait of Sicily. veloped coral provinces, i.e., Santa Maria di Leuca (SML), Southern Adriatic and Strait of Sicily, exploit contrasting topographic scenarios. Vigorous coral growth at SML is favoured by a dramatic seascape generated by the emplacement of gravity mass transported sediment as proposed by Taviani et al. [47] and documented by [50]. A partly similar scenario is found in the Southern Adriatic where also some of the CWC grow on slumped sedimentary blocks [56, 57, 17]. The rich CWC grounds south of Malta [48] take advantage of a submarine escarpment possibly related to regional tectonics of this sector of the Strait of Sicily ([17], Taviani et al. work in progress). Whenever present Pleis-
tocene to Recent coral mounds are small structures, rarely metrical, often representing a single or few episodes of coral growth (e.g., [58, 59]). A case in point are the last glacial buried CWC mounds of the eastern Tyrrhenian described by Remia & Taviani [58]. These small metrical mounds located on tops of a seaward-facing complex relief [59] are made up by Lophelia, Madrepora and Desmophyllum corals. These last glacial mounds were eventually oversilted and buried at the end of the Younger Dryas [58, 43]. A similar situation has been recently identified in the Middle Adriatic, near the edge of the Pomo/Jabuka Pit/ Depression where lush Lophelia, Madrepora and Dendrophyllia
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Figure 4: Examples of deep-water coral occurences in the Mediterranean Sea. A) Living Desmophyllum dianthus and Madrepora oculata from Santa Maria di Leuca coral province, Ionian Sea (CORSARO cruise, Station CR73, ca -670 m). B) Nylon fishing shreds trawled in the Strait of Sicily, offshore Gela (MEDCOR Cruise, Station MEDCOR 74, ca -800 m) fouled by living colonies of Lophelia pertusa (the biggest colony on the top-right is up to 5 cm) and the solitary coral D. dianthus. C) Fe-Mn-coated hardground collected in the Strait of Sicily at ca -600 m during MARCOS Cruise (Station MS 75), note the occurrence of Corallium rubrum, one of the deepest occurrence known so far (up to -600 m); scale bar is 5 cm. D) Very large and thick morphotype of Lophelia pertusa from the Southern Adriatic coral province (SETE06 Cruise, Station SE06-13, ca -300 m); scale bar is 1 cm. E) Radiograph of a gravity core obtained from CWC grounds in the Strait of Sicily during MARCOS Cruise (Station MS12, ca -600 m); scale bar is 5 cm. mounds active up to medieval times have been draped by an important mud influx [55]. Both situations document the importance of fine particle flow, in turn likely controlled by climatic factors [58, 43], in regulating the presence and ultimate fate of deep coral occurrences [31]. Only a few comprehensive studies have been so far devoted to the study of the ecology, biodiversity and functioning of CWC sites in the Mediterranean. By far, SML is at present the CWC province for which most information is available having been the target of many oceanographic missions 710
e.g. RV Urania cruises CORSARO in 2006; RV Universitatis cruises APLABES 1-3 in 2003-2006; RV Meteor cruise M701 in 2006 (ROV Quest); RV Pourquoipas? cruise MEDECO in 2007 (ROV Victor 6000). Prosperous coral growth here is best between 500-700 m [17, 60, 51] and is controlled by oceanographic factors, above all the influx of Adriatic Deep Water [47, 17, 61]. The area is a site of active coral growth since the latest Pleistocene [47, 42]. Its biodiversity is relatively high [62, 49, 63], with references therein) but somewhat lower than counter-
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parts in the Eastern Atlantic. The trophic web has been recently discussed by Carlier et al. [64] and the beneficial function exerted by CWC on meiofaunal diversity underlined by Bongiorni et al. [65]. Other CWC sites are scrutinized for understanding functioning and main characteristics, among which worth mentioning are the Strait of Sicily [17] and the CatalanProvenc¸al canyons [52]. In the Strait of Sicily, increasing information is paid to the coral grounds south of Malta [48]. This site is in fact located in a critical sector of the Mediterranean Sea under the dominant influx of two major water masses, the superficial inflow of Modified Atlantic Water and the deep counterflow of the Levantine Intermediate Water, the latter directly impinging onto the area of active coral growth centred at ca. 450-600 m. A complete swath bathymetric mapping of the area covered by active coral growth mostly limited to a submarine escarpment [17] has been recently completed through CNR cruises CORAL (2002), MARCOS (2007) and MEDCOR (2009). The south Malta site in the Strait of Sicily is particularly interesting since, besides the usual frame building ‘white corals’ Lophelia and Madrepora, a variety of species co-occur to its biodiversity including solitary scleractinians (Desmophyllum, Caryophyllia), yellow corals (Dendrophyllia), antipatharians (Leiopathes), octocorals (the calcified gorgonacean Corallium rubrum, here in the Strait of Sicily at the known deepest range), bivalves, serpulids, sponges, cirripeds (the large barnacle Pachylasma giganteum) and coral-predatory gastropods [66, 60, 67, 68]. Overall, coral habitats support or share the environment with other deep-sea macrofaunal elements besides those mentioned above, among which worth-mentioning are large limids (Acesta
excavata: [69]) and giant oysters (Neopycnodonte zibrowii: [70, 20, 62]).
3
Conclusive remarks
As for the past exploratory research should be continued since many sectors of the Mediterranean Sea are little known or totally unexplored for potential CWC occurrences. This action necessarily implies that the public research should have availability or certain access to opportune technologies such as ROV (Remote Operating Vehicle), AUV(Automatic Underwater Vehicle), gliders and landers what is not the case at present. Confidential industry information indicates that many living coral stocks do exist offshore Sicily, particularly along the Ionian margin and in the Strait of Sicily. The southern side of the Mediterranean Sea (entire North African margin and Levant) is equally highly promising as being colonized by CWC ecosystems. Further exploration should also be conducted in the Ligurian Sea, an area characterized by winter deep-water production, possibly housing prosperous coral growth [71], like the Catalan-Provenc¸al canyons. The newly discovered CWC site in the Marmara Sea is also a strong indication that further research should be carried out there [44]. A search for CWC should be definitely done in the eastern side of the central Adriatic Sea. Chances are that such Adriatic CWC stocks are prominent in controlling coral dispersal and successful settlement in the Southern Adriatic, Ionian Sea and Strait of Sicily following hydrologic pathways linked to North Adriatic Deep Water production and cascading events (e.g. [72]). With few exceptions, most Mediterranean CWC sites are still little known for their biodiversity and functional aspects, requir711
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ing additional studies on invertebrate and microbial communities, trophic webs and CWC impact on nekton. Regrettably, habitat mapping of main Mediterranean CWC provinces is still in its infancy, with preliminary work mostly limited at Cap de Creus [52] and SML [51]. The connectivity of disjointed CWC sites is also a major target for research to be pursued at the short term. A number of live-collected scleractinians (Lophelia, Madrepora and Desmophyllum) have been collected to be scrutinized for genetics (including degraded DNA on subfossil corals) to help unveiling their links [70]. In fact, no evidence yet exists regarding intra-Mediterranean source and dispersal pathways of CWC larval flows what should be really considered a major gap in understanding AtlanticMediterranean cold-water coral biogeography and evolution. Relevance should be paid to the usefulness of CWC and associated invertebrates (e.g., Corallium, antipatharians, bivalves etc.) as climatic tools. Increasing evidence is produced about the relevance of extracting geochemical signals of valuable paleoceanographic (therefore paleoclimatic) significance out of the carbonate and organic skeletons (e.g., [73, 74], and this volume for a revision). As a conclusive remark, it must be brought to public and political awareness that CWC habitats in Mediterranean as elsewhere are threatened by natural and, more commonly, anthropogenic causes (fishing malpractices, dumping, lit-
tering, ocean pollution and acidification). There is an unprocrastinable urgence not only for conducting basic scientific studies but also for concomitant political action capable of ensuring a future to such complex and important marine habitats.
4
Acknowledgments
The authors of this short article are or have been part of the ISMAR-CNR team funded to carry out CWC research in the last 15 years. Thanks are due to Captains (Emanuele Gentile, Vincenzo Lubrano Lavadera, and the late Nicolangelo Lembo), Officers, Crew and Shipboard Staff Members of CNR cruises CS96, LM99, CORTI, CORAL, COBAS, GECO, MARCOS, ARCO, MEDCOR, devoted to the exploration of Mediterranean CWC onboard RV Urania from 1996 up to present. Proper funding to carry out such investigations was provided by CNR grants, ESF Eurocore/Euromargins Moundforce, MiurFIRB Aplabes, COMP, EU Hermes and EU Hermione programmes. Cruise M70-1 of RV Meteor (Chief Scientist A. Freiwald) was highly significant in discovering and documenting many important CWC occurrences in the Central Mediterranean Sea. Among the many people who provided information, inspiration and ideas on the issue of Mediterranean cold-water corals Andr´e Freiwald and Helmut Zibrowius deserve a special mention. This is ISMARBologna scientific contribution n. 1675.
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[2] J.M. Roberts and A. Freiwald. Integrated European Research into Cold-Water Coral Reefs. The Journal of Marine Education, 21(4):41–45, 2005. [3] J.M. Roberts, A.J. Wheeler, and A. Freiwald. Reefs of the Deep. The Biology and Geology of Cold-Water Coral Ecosystems. Science, 312:543–547, 2006. [4] A. J. Wheeler, A. Beyer, A. Freiwald, H. de Haas, V.A.I. Huvenne, M. Kozachenko, K. Olu-Le Roy, and J. Opderbecke. Morphology and environment of cold-water coral carbonate mounds on the NW European margin. International Journal of Earth Science (Geologische Rundschau), 96(1):37–56, 2007. [5] A. Foubert and J.-P. Henriet. Nature and Significance of the Recent Carbonate Mound Record. Lecture Notes in Earth Sciences, 126:318, 2009. [6] P.P.E. Weaver and V.Gunn. Introduction to the Special Issue: HERMES-Hotspot Ecosystem Researche on the Margin of European Seas. Oceanography, 22(1):12– 15, 2009. [7] P.P.E. Weaver, A. Boetius, R. Danovaro, A. Freiwald, V. Gunn, S. Heussner, T. Morato, I. Schewe, and S. van den Hove. The Future of Integrated Deep-Sea Research in Europe: The HERMIONE Project. Oceanography, 22(1):178–191, 2009. [8] A. Freiwald, J.H. Foss˚a, A. Grehan, T. Koslow, and J.M. Roberts. Cold-water coral reefs. Out of sight - no longer out of mind. pages 1–84, 2004. [9] A. Freiwald. Reef-Forming Cold-Water Corals. In: Ocean Margin Systems. Springer-Verlag, pages 365–385, 2002. [10] A. Freiwald, V. H¨uhnerbach, B. Lindberg, J.B. Wilson, and J. Campbell. The Sula Reef Complex, Norwegian Shelf. Facies, 47:179–200, 2002. [11] A. Freiwald, R. Henrich, and J. P¨atzold. Anatomy of a deep-water coral reef mound from. Stjernsund. SEPM Spec. Publ., 56:141–161, 1997. [12] B. De Mol, P. van Rensbergen, S. Pillen, K. van Herreweghe, D. van Rooij, A. McDonnell, V.A.I. Huvenne, M. Ivanov, R. Swennen, and J.-P. Henriet. Large deepwater coral banks in the Porcupine Basin, southwest of Ireland. Marine Geology, 188:19–31, 2002. [13] T.C.E. van Weering, C. Dullo, and J.-P. Henriet. An introduction to geospherebiosphere coupling cold seep related carbonate and mound formation and ecology. Marine Geology, 198:1–3, 2003. [14] D. van Rooij, B. De Mol, V.A.I. Huvenne, M. Ivanov, and J.-P. Henriet. Seismic evidences of current-controlled sedimentation in the Belgica mound province, upper Porcupine slope, southwest of Ireland. Marine Geology, 195:31–53, 2003.
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[27] F. Althaus, A. Williams, T. A. Schlacher, R. J. Kloser, M. A. Green, B. A. Barker, N. J. Bax, P. Brodie, and M. A. Schlacher-Hoenlinger. Impacts of bottom trawling on deep-coral ecosystems of seamounts are long-lasting. Marine Ecology Progress Series, 397:279–294, 2009. [28] J.M. Guinotte, J. Orr, S. Cairns, A. Freiwald, L. Morgan, and R. George. Will human-induced changes in seawater chemistry alter the distribution of deep-sea scleractinian corals? Frontiers in Ecology and the Environment, 4(3):141–146, 2006. [29] R. Selli. Ricerche geologiche preliminari nel Mar Tirreno. Giornale di Geologia, 37:1–249, 1970. [30] M. Taviani and P. Colantoni. Paleobiocoenoses profondes a scleractiniaires sur l’escarpement de Malte-Syracuse (Mer Mediterran´ee): leur structure, leur age et leur signification. Revue de l’Institut Francais du Petrole, 39(5):547–559, 1984. [31] R. Sartori. Factors affecting the distribution of ahermatypic corals on the Mediterranean seafloor: a probabilistic study. Deep-Sea Research, 27(A):655–663, 1980. [32] G. Delibrias and M. Taviani. Dating the death of Mediterranean deep-sea scleractinian corals. Marine Geology, 62:175–180, 1985. [33] A. Milne-Edwards. Observations sur l’existence de divers mollusques et zoophytes a` de tr`es grandes profondeurs dans la mer M´editerran´ee. V. Masson et fils, pages 149–157, 1861. [34] J. M. P´er`es and J. Picard. Nouveau Manuel de Bionomie Benthique de la Mer M´editerran´ee. – Recueille des Travaux de la Station Marine d’Endoume. 31(47):1– 137, 1964. [35] H. Zibrowius. Les Scl´eractiniaires de la M´editerran´ee et de l’Atlantique nordoriental. M´emoires de l’Institut Oc´eanographique, 11:1–284, 1980. [36] M. Taviani, C. Corselli, A. Freiwald, E. Malinverno, F. Mastrototaro, A. Remia, A. Savini, and A. Tursi. First geo-marine survey of living cold-water Lophelia reefs in the Ionian Sea (Mediterranean basin). Facies, 50:409–417, 2005. [37] I. Di Geronimo, C. Messina, A. Rosso, R. Sanfilippo, F. Sciuto, and A. Vertino. Enhanced biodiversity in the deep: Early Pleistocene coral communities from Southern Italy. pages 61–86, 2005. [38] J. Titschack and A. Freiwald. Growth, deposition, and facies of Pleistocene bathyal coral communities from Rhodes, Greece. pages 41–59, 2005. [39] M. Taviani. The Mediterranean benthos from Late Miocene up to Present: ten million years of dramatic climatic and geological vicissitudes. Biologia Marina Mediterranea, 9(1):445–463, 2002. 715
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[51] A. Vertino, A. Savini, A. Rosso, I. Di Geronimo, F. Mastrototaro, R. Sanfilippo, G. Gay, and G. Etiope. Benthic habitat characterization and distribution from two representative sites of the deep-water SML Coral Province (Mediterranean). DeepSea Research II, 57(5/6):380–396, 2010. [52] C. Orejas, A. Gori, C. Lo Iacono, P. Puig, J.-M. Gili, and M.R.T. Dale. Coldwater corals in the Cap de Creus canyon, northwestern Mediterranean: spatial distribution, density and anthropogenic impact. Marine Ecology Progress Series, 397:37–51, 2009. ´ [53] G. Alvarez P´erez, P. Busquets, B. De Mol, N.G. Sandoval, M. Canals, and J.L. Casamor. Deep-water coral occurrences in the Strait of Gibraltar. In: Cold-Water Corals and Ecosystems. pages 207–221, 2005. [54] D. Vafidis, A. Koukouras, and E. Voultsiadou-Koukoura. Actiniaria, Corallimorpharia and Scleractinia (Hexacorallia, Anthozoa) of the Aegean Sea, with a checklist of the Eastern Mediterranean and Black Sea species. Israel Journal of Zoology, 43:55–70, 1997. [55] M. Taviani, L. Angeletti, A. Ceregato, and T. Bakran-Petricioli. Was enhanced riverine input responsible for the demise of Central Adriatic cold water reefs in historical times? Hermione Annual Meeting 2010 (La Valletta, Malta 12-16 April). Abstract book, page 49, 2010. [56] F. Trincardi, A. Freiwald, M. Taviani, and G. Verdicchio. Slide (slump)-created topography enhancing deep-water coral growth. Hermes 2nd Annual Meeting Hotel Tivoli Almansor, Algarve 24-30 March 2007, Conference programme & abstracts., page 10, 2007. [57] F. Trincardi, M. Taviani, A. Freiwald, L. Angeletti, F. Foglini, D. Minisini, A. Piva, and G. Verdicchio. An actualistic scenario for olistostrome genesis and emplacement (Gondola slide, SW Adriatic margin). Rendiconti online della Societ`a Geologica Italiana, 3(2):762–763, 2008. [58] A. Remia and M. Taviani. Shallow-buried Pleistocene Madrepora-coral mounds on a muddy continental slope, Tuscan Archipelago, NE Tyrrhenian Sea. Facies, 50:419–425, 2005. [59] A. Remia, M. Taviani, and L. Gasperini. Imaging fossil deep-sea coral mounds by integrated geophysical techniques. Geophysical Research Abstracts, 7:07689, 2005. [60] M. Taviani, L. Angeletti, B. Antolini, A. Ceregato, J. Evans, C. Froglia, L. Gandolfi, V. Garilli, C. Giampieri, E. Hernandez Goldstein, E. Leidi, F. Lorenzini, C. Maier, P. Montagna, M. Naumann, R. Rodolfo-Metalpa, A. Rosso, P.P. Zammit, and H. Zibrowius. Hermione cruise MEDCOR of RV Urania supplies new significant information on deep-sea ecosystems in the Strait of Sicily (deep-sea corals, cold seeps, 717
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[71] L. Tunesi and G. Diviacco. Observation by submersible on the bottoms off shore Portofino Promontory (Ligurian Sea). Atti del 12°Congresso dell’Associazione Italiana di Oceanologia e Limnologia (Isola di Vulcano, 18-21 Settembre 1996), pages 61–74, 1997. [72] F. Trincardi, F. Foglini, G. Verdicchio, A. Asioli, A. Correggiari, D. Minisini, A. Piva, A. Remia, D. Ridente, and M. Taviani. The impact of cascading currents on the Bari Canyon System, SW-Adriatic Margin (Central Mediterranean). Marine Geology, 246:208–230, 2007b. [73] P. Montagna, M. McCulloch, M. Taviani, C. Mazzoli, and B. Vendrell. Phosphorus in cold-water corals as a proxy for seawater nutrient chemistry. Science, 312:1788– 1791, 2006. [74] P. Montagna, S. Silenzi, S. Devoti, C. Mazzoli, M. McCulloch, G. Scicchitano, F. Sparaini, and M. Taviani. High-resolution natural archives provide new tools for climatic reconstruction and monitoring the Mediterranean Sea. Rendiconti Lincei, 19:121–140, 2008.
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Stratigraphic and Morphological Evidences of the Last Cycle of Relative Sea-Level Change; Some Examples from the Late Pleistocene to Holocene Sedimentary Successions in the Sicilian Offshore M. Mancuso1 , M. Agate2 , L. Fallo2 , F. Vaccaro2 , R. Catalano2 1, Institute for Coastal Marine Environment, CNR, Mazara del Vallo (TP), Italy 2, Department of Geology and Geodesy, University of Palermo, Palermo, Italy maria.mancuso@iamc.cnr.it Abstract Seismo-stratigraphic studies of the Late Pleistocene to Holocene sedimentary successions, located along the Sicilian continental margin, represent an opportunity to interpret the evolution of a sedimentary basin in terms of interplay among global sea level variability, sedimentary ciclicity, oceanographic changes and tectonic deformation. The focus of this paper is to compare the geological evolution of distinct areas of the submerged extension of the Sicilian chain, from the Western-Northern sector to the Southern one, to define the relations between sedimentation and eustatic sea-level oscillations. These relations are commonly complicated by deformation occurring in both the source (on land) and depositional (on sea) areas. In the Sicilian offshore, the Late Pleistocene to Holocene Q.5.e type 1 depositional sequence records the last fourth - (∼ 100 kyr) and fifth - order (∼ 20 kyr) eustatic sea level fluctuations, after the previous highstand of isotope Stage 5e. In the analysed areas, the Q.5.e depositional sequence (80 kyr to 0 b.P. in age) is made up of four systems tracts: the Falling Stage, Lowstand, Trangressive and Highstand Systems Tracts. Despite of the uniformity of quadripartite organization of the Q.5.e depositional sequence, the four Systems Tracts show different sedimentary response to local physiography, tectonics (including subsidence/uplift rates), type and rate of the sediment supply, and oceanographic conditions.
1
Introduction
It was not until the late Eighty that the models of Sequence Stratigraphy have been applied to the interpretation of the stratigraphic architecture of the continental margins by means of high-resolution seismic profiles, providing wide documentation about the casual links between
long-term, quasi-periodic variations in the Earth’s orbital parameters (eccentricity, obliquity, and precession), global climatic and sea- level changes, and their influence on the sculpturing - during the upper Pleistocene and Holocene times - of the continental shelves all over the world. In the Mediterranean Plio-Pleistocene interval - and in particularly in the Sicilian
Marine Geology
sedimentary successions - both high frequency (21 - 41 kyr) and low - frequency fluctuations (100 kyr or multiple, generally 200 - 400 kyr) are present in the marine record [1, 2, 3, 4, 5, 6]. These oscillations have been related to Milankovitch effects forced by astronomical factors [7, 8, 9, 10]. Important ice sheets volume variations are implied by these climatic cycles. When the cycles operate in concert, a high amplitude, high - frequency 4th (100 - 125 kyr) strongly asymmetric glacioeustatic cycle has been formed, with a amplitude of approximately 120 m, and modulated by higher frequency cycles. The original assumptions of Vail and coauthors and their global models, based on analysis of Meso-Cenozoic successions have been revaluated and modified, with the introduction of further controlling factors on sedimentation like the variability in the oceanographic regime, the physiographic context, the type and dispersal pattern of sediment supply and, of course, the local tectonics [11, 12, 13, 14, 15]. It is likely, on the basis of the data examined, that the same factors have played an important role in the development of the late Quaternary depositional sequences Q.5.e, recognised along the Sicilian margin [16, 17, 6, 18]. In this paper, we bring together a great deal of observational data on the sequence stratigraphic architecture of upper Pleistocene to Holocene deposits, outcropping in the Sicily offshore, from the Egadi Islands to the Termini Imerese Basin (in the Western and Northern sectors) and in the Selinunte-Sciacca offshore (in the Southern one), to define the effects of cyclic variations of the relative sea levels on the sedimentary and morphological processes of these wave-dominated continental shelves, and to discuss the geological evolution of 722
the 4th order depositional sequence recognized.
2
General setting
The Sicilian continental margin (Figure 1) belongs to the Sicilian-Maghrebian thrust belt and its deformed foreland basin (Figure 2).
2.1
Geological framework
The present structural setting has been inherited from the continental collision between the Corsica - Sardinia microplate and the Sicilian continental margin [23, 16] and the subsequent (mostly Plio Pleistocene) extensional tectonics, linked to the Tyrrhenian Sea opening [24, 25, 25]. The chain, partly submerged, progressively overthrusted the foreland units eastand south-eastward [23, 26, 20]. In the western and northwestern sectors (Egadi Islands, Erice, Castellammare, Palermo and Termini Imerese Basins), syn and post-Messinian extensional growth faulting, linked to the Tyrrhenian Sea opening, created intraslope basinal areas. In the southern sector, the WNW - ESE trending foredeep basin [20] is represented by the narrow, weakly deformed depression (the Gela Basin), partially buried under the Gela Thrust System [26]. During the last 125 kyr, a moderate regional uplift (0,01 - 0,22 mm¡yrâ&#x2C6;&#x2019;1 ; [27]) affects the coastal area - from the Egadi Islands to Cape Playa - deforming the littoral deposits, pertaining to the Last Interglacial highstand (dated at 130 â&#x20AC;&#x201C; 118 kyr B.P., [28]). The lack of preserved MIS 5.5 marine terraces in the southern Sicily also points to the tectonic subsidence during upper Pleistocene to Holocene times [27].
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Figure 1: Main physiographic characters of the Sicilian offshore. The Sicilian continental margin, Late Neogene to Recent in age, extends from the northern coast of Western Sicily northwards to the Elimi Chain and the Ustica volcanic complex, westward to the Egadi Islands, eastwards to the Cefal`u Basin, while southwards to the Adventure Bank and the Gela Basin. Inset: studied areas. A) Egadi Islands, B) Gulf of Castellammare and Palermo Mountains offshore, C) Gulf of Termini Imerese, D) Selinunte - Sciacca offshore.
2.2
Marine morphology
The modern seascape of the Sicilian offshore displays different morphologies, as it is characterised by the presence of a narrow and discontinuous shelf, alternating to broad shallow banks, platforms, steep slopes or a series of basins and structural highs. 2.2.1
The Egadi Islands
The Egadi Islands platform (Figure 3) is articulated in a structurally controlled rhomboid platform (the Marettimo Island continental shelf), and an open pericontinental shelf (around Favignana and Levanzo Islands), separated by narrow seaways. A number of small shoals, rising up to a water depth of -6 m and -40 m, and gentle
seaward sloping wave-cut terraces are located at different depths ranging to 30 m/ 110 m. The break in slope at the shelf edge is at about -130 m /-150 m. The slope degrades with an angle of 4째 to 6째. Slides, debris flow lobes and gullies are mainly diffuse along the steep flanks of the Marettimo insular platform, while erosional scours are localised along the MesoCenozoic rocky outcrops. The NS elongated Marettimo Channel divides the Marettimo platform from the Sicilian continental shelf. The oceanographic setting of the Egadi Island is influenced by the wind-driven component of the Tyrrhenian Sea circulation [29] and by the thermoaline circulation of the Mediterranean Sea [30, 31]. The prevailing winds (from SE and NW)
723
Marine Geology
Figure 2: Structural map of the South-Central Mediterranean Sea. Legend: 1) oceanic basins; 2) Eastern Sardinia tectonic units; 3) Kabylian - Calabrian tectonic units; 4) Sicilian - Appenninic tectonic units; 5) Gela Embricated Thrust System; 6) Foreland and deformed foreland: a) Plio-Pleistocene sedimentary cover; b) Miocene clastic-carbonatic deposits; c) Lower Miocene to Triassic carbonate substratum; 7) Volcanics; 8) Main thrust fronts, 9) other thrusts; 10) strike - slip faults; 11) structural axes; 12) normal faults. Acronyms: STF: Sardinia Thrust Front; DTF: Drepano Thrust Front; ETF: Egadi Thrust Front; ATF: Adventure Thrust Front; GTF: Gela Thrust Front; MTF: Main thrust front of the Sicilian - Maghrebian Chain. From [19, 20, 21] modified. The seismicity (M â&#x2030;Ľ 3 earthquakes) occurred between 1973 and 2006 in the Southern â&#x20AC;&#x201C; Central Mediterranean is shown in the inset. Light grey dots indicate events shallower than 30 km, grey ipocentre depths between 30 and 100 km, and dark grey dots ipocentre depths larger than 100 km. Data from NEIC (modified from [22]). produce almost continual wave action on the west and east facing margins. Due to the lack of major rivers in western Sicily, the terrigenous input to the shelf is low and this favours the growth of lime-secreting organisms and the proliferation of extensive sea grass meadows (Posidonia oceanica).
724
2.2.2
The Castellammare Gulf and the Palermo Mountains offshore
Moving eastwards, the Gulf of Castellammare is the largest coastal embayment in North-western Sicily, bounded to the East by the Palermo Mountains and to the West by the peninsula of San Vito Lo Capo (Figure 4). The continental shelf is 8 km width on the southern and eastern sides of the gulf, but it is very narrow or absent along the peninsula of San Vito Lo Capo and, at the opposite side, along the Cape Rama
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Figure 3: Marine geological map of the Falling Sea Level, Lowstand, Transgressive and Highstand Systems Tracts of the Q.5.e depositional sequence in the Egadi Islands offshore. Highstand deposits have been mapped only when their seismic resolvable thickness was > 2 msec (t.w.t). Legend. Stratigraphy: 1) Highstand Systems Tract - α: channels and moats; 2) Transgressive Systems Tract; 3) Lowstand prograding complex (Lowstand Systems Tract); 4) Falling Stage Systems Tract; 5) middle Pleistocene deposits; 6) lower Pliocene deposits; 7) Meso-Cenozoic substrate. Marine morphological characteristics: a) bioherms; b) paleoshorelines; c) breaks in slope; d) shelf break; e) wave cut terraces; f) contourites and shelf drift axes; g) slides and submarine failures; h) retrogressive slide scars; i) isobaths in meters; tectonic features: l) folds; m) normal faults; n) thrusts and reverse faults. headland. The coastal San Vito Mountains sea cliffs, and the small pocket beaches contrast with the broad, low relief alluvial plain and sandy-gravelly beaches along the central coasts of the Castellammare Gulf. At the eastern side, bedrock crops out on the inner and middle shelf; bathymetry mirrors the onshore morphology, with the Cape Rama headland extending offshore as wave – cut terraces with a very thin or non – existent sediment cover, colonised by sea
grass meadows. The shelf is delimited by a break located at a variable depth of between -130 m and -190 m. The continental slope deepens to -1000 m, at inclinations of 4° - 9°, and has been incised by numerous gullies and major canyons. All the material supplied by the Freddo, Jato and Nocella Rivers is trapped in a NWSE canyon - the Castellammare Canyon - bordered on its western margin by the San Vito High and eastwards by the off-
725
Marine Geology
shore prolongation of the Palermo Mountains. These structural ridges confine the Castellammare turbidite system preventing westward and eastwards progradation. Fluid vents and pockmarks, associated to the main slides, are diffuse, suggesting a relation between the two phenomena. The Palermo Mountains offshore (Figure 5) comprises a narrow shelf characterised by extensive area of rocky bottom, which continuity is interrupted by the Barra High (the northeastern prolongation of the Cape Gallo headland), which extends along the shelf and rises close to -50 m of depth. The continental shelf forms a narrow platform, approximately 10 km wide, steeply sloping to the shelf break located to -140 m /1-60 m of depth. A paleoshoreline, between -100 m /-110 of depth, divides the shelf in a outer smooth sector and in an inner rough area, where paleoriver valleys occur. A N-S trending canyon is the main feature that cuts the upper slope; this feature, here named Carini Canyon, is a scarcely sinuous, broad elevated erosive channel with wide levees. The upper slope morphology of the Palermo Mountains results from a combination of slides, retrogressive scars, large gullies and rocky outcrops. Debris flow lobes are also developed downslope some of the upper slope scars, particularly on the western and eastern sides of the Barra High. 2.2.3
The Termini Imerese Gulf
The Gulf of Termini Imerese (Figure 6) displays a narrow and gently seaward sloping, wave-dominated continental shelf, fed by the fine-grained sedimentary supply of the Torto, Northern Imera, Roccella, and Poletto Rivers. The continental shelf is
726
mostly covered by a continuous sandymud blanket, composed of terrigenous fine sands, silty clays and clayey silts formed by the coalescing prodelta muds of the rivers. Sedimentary creep and acoustic masking phenomena affect the prodelta slopes. The inner shelf, with slopes between 0.15° and 0.28°, encompasses 0 - 50 m water depths, while the outer shelf having a slope of between 1° and 3°, is delimited by a break located at a variable depth of between -90 m and -130 m of depth. This break in slope is related to the presence of outcropping mud volcano diapiric ridges, which extend parallel to the shelf edge for about 3 km / 5 km. 2.2.4
The Selinunte – Sciacca offshore
The southwestern Sicilian shelf (Figure 7) displays a large, almost gently seaward sloping surface, up to 16 km wide, with a shelf edge located at water depths ranging to -83 to -185 m. The upper continental slope is irregular and incised by numerous canyons and scars associated with slumps and debris flows. The western flanks of the Selinunte – Sciacca shelf and slope areas represent the eastern sectors of the Adventure Bank. The narrow NNW-SSE trending incision, located off the Granitola Cape, is the southeastern prolongation of the Mazara del Vallo – Marettimo Channel. Southwards, shallow rocky and volcanic banks, as the Graham and the Nerita Banks, rise up to 30 m from the surrounding sea floors. The interaction between the seafloor topography and the northward flowing Levantine Intermediate Waters (LIW) is responsible for the creation of seafloor bedforms and erosional moats, which have developed below the shelf edge and -300 m isobaths.
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Figure 4: Shaded relief image of the Castellammare Gulf, compiled from multi-beam data. The main elements have been recognised by means of morphological and seismicstratigraphic analysis (seismic data set is shown up in the left corner), side scan sonar mosaics, and aerial imageries with ligth levels and contrasts modified to delineate sediment and rock outcrops along the coasts, at shallow water depths (< 5 m). In the inset, the geological map of the Falling Stage, Lowstand, Transgressive and Highstand Systems Tracts of the Q.5.e depositional sequence in the Gulf of Castellammmare is shown (from [18]).
727
Marine Geology
3
Data and methodologies
The seismic profiles analysed in this study, for a total length of approximate 3000 km, densely cover the studied areas of the Sicilian margin. The seismic profiles have been acquired with 1000 joule, 4.5 kJ and 16 kJ Sparker and with higher frequency 3.5 kHz and 16 kHz Chirp Sub Bottom echo sounders. The high-resolution mono - channel seismic surveys and multibeam data have been carried out by the Department of Geology and Geodesy of the University of Palermo, in collaboration with the IAMC – C.N.R. Institute, the E.N.E.A., the Italian National Geological Service, the Department of Geological Sciences of the University of Cagliari. The peculiarities of the seismic data set allow us the application of high-resolution sequence stratigraphy to depositional sequences of limited thickness and areal extension. The cores collected from the Sicilian margin provide valuable information on the lithology of the Systems Tracts, identified seismostratigraphically. Biostratigraphic data [32] and the recent high – resolution biostratigraphy of the ODP Hole 963B [33, 34] allow us to restrict the studied deposits to the late Pleistocene – Holocene interval, after the MIS 5e highstand.
4
Sequence stratigraphic analysis
In the analysed areas, the Q.5.e depositional sequence (Figure 8) is composed of four Systems Tracts: the Falling Stage (FSST, 80 kyr - 30 kyr b. P., in age), the Lowstand (LST, 30 k kyr – 18 k kyr b.P.), the Transgressive (TST, 15 k kyr – 6,0 728
kyr b.P.) and the Highstand Systems Tract (HST, 6 kyr – 0 kyr b.P.). Despite of the uniformity of quadripartite organization of the Q.5.e depositional sequence, the four Systems Tracts (STs) show different sedimentary response to the local physiography, tectonics, type and rate of the sediment supply, and oceanographic conditions [17, 18]. Within the Q.5.e sequence, the four Systems Tracts are separated by basin-wide key surfaces identified by lateral terminations of reflectors, showing different sedimentary and seismostratigraphic expression in each sectors of the basins.
4.1
Genesis and significance of the key surfaces
The sequence boundary is represented on the continental shelf area by an unconformity surface of subaerial exposure and fluvial down cut, and a correlative paraconformity on the upper slope. This surface has been interpreted as the type 1 upper Pleistocene-Holocene sequence boundary occurring beneath the forced regressions [11, 35]. Between -120 and -160 m in depth, this surface incises the deposits of the underlying Pleistocene sequences forming a ”V” shaped valleys. The top of the shelf margin units is a transgression erosional surface (ravinement surface, [36]) replaced in the upper slope by a drowning surface, without erosional features. On the outer shelf, between -50 and -100 m of water depth, three orders of submerged paleo-shorelines (formed during the Holocene stillstands) have been laterally traced for several kilometres, bounding wave-cut terraces. The maximum flooding surface marks the transition between TST and HST seismic
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Figure 5: Geologic Map of Palermo Mountains offshore of the Falling Stage, Lowstand, Transgressive and Highstand Systems Tracts of the Q.5.e depositional sequence. Legend: 1) Highstand Systems Tract; 2) Transgressive Systems Tract; 3) Lowstand Systems Tract; 4) Falling Stage Systems Tract; 5) Acoustic substratum; 6) marine terraces; 7) incised valleys; 8) Posidonia oceanica meadows; 9) bioherms and beach rocks; 10) channels; 11) debrites and scars; 12) canyons; 13) levees; 14) a: buried tectonics, b: surface tectonics; 15) isobaths in meters units.
4.2
Stratal patterns of the Systems Tracts
The STs consist of two or more depositional elements and are characterised by different seismic facies, which are defined by the nature of reflectors, their internal geometry, and external morphology. Forced regressions, pertained to the FSST, have been identified in the sectors of the Sicilian outer shelf, where they onlap the upper Pleistocene unconformity surface, at an average distance of 9 km /12 km from the present shoreline, and at depths of -90 to -120 m. The prevalent seismic facies is charac-
terised by complex sigmoid - oblique reflectors and in rare cases oblique parallel, with an average slope of about 2.5째. The amplitude and the lateral continuity of seismic reflectors tend to thin out seawards. A less common reflector facies is characterised by oblique tangential configurations, with depositional dips approaching 10째, terminating up dip by a gently seaward sloping terrace. The forced regressive units can be subdivided into stacked and laterally offset clinoform sets, bounded by regressive surfaces of marine erosion of local extent, and/or their equivalent downdip surfaces. In seismic profiles, they are imaged by a stack of two or three 5th order parasequences that pinch-out on the basal un729
Marine Geology
Figure 6: Geologic Map of the southern Termini Imerese Gulf. Legend: 1) Posidonia oceanica meadows; 2) Cymodocea nodosa meadows; 3) gas-charged sediments; 4) mud diapirs; 5) modern fluvial mouths; 6) gullies and drainage pathways; 7) paleoshorelines; 8) breaks in slope; 9) buried normal faults; 10) normal faults; 11) Highstand Systems Tract; 12) Falling stage and Lowstand Systems Tracts; 13) pre-Pleistocene acoustic substratum. conformity between -90 and -120 m of depth. Inside each wedge, the offlap break deepens seawards drawing a sigmoid trajectory; otherwise, onlap terminations have not been often recognized, due to the enhanced erosion underwent during the ensuing marine transgression. The Lowstand prograding complex (Lpc) units, pertaining to the LST, are situated near the present-day shelf edge, and in the upper slope. They shows a seismic facies with predominantly oblique, medium and high amplitude reflectors of good lateral continuity, with foreset gradients ranging from 4° to 10°. At the base of these shelf-perched units, chaotic seismic facies and small mounds have been identified in close association with gullies and canyons, cutting the outer shelf. In the high energy oceanographic regime, 730
as those affected the Egadi Islands and the southern Selinunte – Sciacca offshore, along the upper slope, the forced regressive and the lowstand shelf - margin units are partially eroded by shelf – edge parallel separated moats, basinward confined by thinly layered convex upward seismic units. These units, up to 15-20 m high and 2 km long, show a slight symmetric profile and aggradational internal geometry and they have been interpreted as shelf edge contourites. The depocenters of the FSST and Lpc lengthen parallel to the paleomargin of the continental shelves but the major thicknesses are identified near to the canyon heads. TSTs are bounded at the base by an irregular transgressive erosional surface, with wave-cut terraces and channel features, reaching 500 m in width and 10 m in depth. Its upper boundary coincides with the max-
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Figure 7: Geologic map of the Q.5.e. depositional sequence in the Selinunte – Sciacca offshore. Legend: 1) Highstand Systems Tract; 2) Trangressive Systems Tract - α: shelf edge mud drift; 3) Lowstand Systems Tract - α: slope mud drift: 4) Falling Stage Systems Tract - α: slope fan. Marine morphological characteristics: a) bioherms; b) volcanic banks; c) pockmarks; d) mud diapirs; e) paleoriver channels; f) erosive conduits; g) gullies; h) shelf break; i) gas-charged sediments; l) wave-cut terraces; m) drift axes; n) debrites and talus; o) retrogressive slump scars; p) isobaths in meters. In the inset, the used seismic data base is shown. imum flooding surface. In the areas characterised by high rate of sedimentary supply as the Castellammare and Termini Imerese Gulfs or the Sciacca offshore, the most striking feature of TST is the retrogradational pattern of the three parasequences separated by mixed (flooding and erosional) surfaces, extending laterally throughout the studied basins. The three sub-units display a progressive landward shift of the coastal onlaps, which encroach between -98 m to -90 m, -75 m to -63 m and -60 m to the present -45 m isobath. The oldest parasequence, also detected in the Egadi Islands and in the Carini Embayment, is located in the outer sector of the
shelf and rests on a widespread erosional terrace. The intermediate parasequence is formed by the stack of prograding bodies, with low-angle clinoforms (2° - 5°), truncated at their tops by sub-horizontal ravinement surfaces, and plain – parallel deposits above. The last parasequence of TST retrogrades up to the mouths of the rivers, in the modern inner shelves. In the Castellammare and the Termini Imerese gulfs, as well as the Sciacca offshore, the HSTs are marked by thick prograding wedges, which extend continuously from the inner as far as the middleouter shelf. In the inner shelf off the
731
Marine Geology
Figure 8: Comparison between the stratal architecture of the Q.5.e depositional sequence in the studied areas. From the upper: Egadi Islands area. Legend: 1) Highstand Systems Tract; 2) Transgressive Systems Tract - Îą: bioherms; 3) Lowstand prograding complex (Lowstand Systems Tract), c shelf edge drifts; 4) Falling Stage Systems Tract - Îą: debris flows, pe: paleosoils; 5) seismic reflectors; 6) Tyrrhenian littoral deposits (MIS 5e); 7) Plio-Pleistocene deposits; 8) Miocene deposits; 9) Meso-Cenozoic deposits. Palermo Mountains offshore. Legend: 1) Highstand Systems Tract, a) shelf and slope clastic deposits, b) biogenic sands and gravels, forming the modern littoral prism; 3) Lowstand Systems tract, a: coastal regressive units, b: upper slope deposits; 4) Falling Stage Systems Tract; 5) Plio-Pleistocene sedimentary successions; 6) Meso-Cenozoic units; 7) slide scars; 8) slump complex; 9) faults. Termini Imerese area. Legend: 1) Highstand Systems Tract; 2) Transgressive Systems Tract; 3) mud diapirs; 4) Lowstand prograding complex, a: slumps and debrites; 5) Falling Stage Systems Tract, b) mud slope fan deposits; 6) Plio-Pleistocene sedimentary successions. SB: sequence boundary; SB + tse: merged transgressive surface of erosion and sequence boundary; mfs: maximum flooding surface. mouths of the main rivers, acoustic masking affects this ST, suggesting the presence of biogenic gas. The HSTs have been subdivided into two parasequences, separated by erosional surfaces. The lower parasequence shows faintly prograding geometries, shingled and oblique parallel, while the upper is oblique tangential. The modern depositional surface truncates the topsets of the last parasequence, suggesting the position of the modern fair-weather base level at 20 m / -30 m of depth. 732
Sometimes, in the inner shelves, creep and slumping features deform HSTs. These features lie at 2 - 3 km seaward from the coastlines, within elongated band in subsurface depths from 75 to 25 ms (twoway travel time), where the thickness of the HSTs exceeds 25 m close to the river mouths. Mud volcanoes and pockmarks in the subsurface occur updip and they indicate the existence of gas-charged fluid expulsion (Figure 9).
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Figure 9: Gas-related expulsion features: fossil, with breccias rims (a) and actual (b) phenomena. Side scan sonar lines from the Castellammare Gulf.
5
The development of the Q.5.e depositional sequence
During the formation of the Q.5.e depositional sequence, there are four factors which control its genesis and overall stratigraphic architecture: glacio-eustatism and tectonics, hydrodynamic regime, and the sedimentary supply changes. Tectonics is a crucial issue since the progressive seaward tilting/ or the uplift creates /or withdraws accommodation space necessary for development of prograding wedges at the shelf margin. The regime of the sedimentary supply and the dispersion patterns on the sea floor influence the development of particular facies associations, and consequently the overall depositional styles.
5.1
The record of the relative sea – level fall and the Last Glacial Maximum
Synchronously with the global climatic cooling and increasing continental ice volumes [37, 38, 39], during the transition from interglacial Riss - W¨urm to glacial
W¨urm, the relative sea level in the Sicilian margin started to drop after isotope Stage 5e and reached its lowest stand at -120 m at the W¨urmian glacial maximum, at about 18.000 a b.P. [40]. The 4th order FSSTs have been deposited as sea level progressively dropped, from Tyrrhenian positions (+ 5 /+8 m above the present sea level), during the isotopic stages 3 and 2. The clinostratified lithosomes have been for the most part interpreted as deltaic units, connected to well-developed ancient single stream channels and dominated by the action of rivers. Another characteristic is the presence of syndepositional slide scours and associated widespread slides at the base of the clinostratifications. Where the presence of stream mouths has not been detected or where headlands and bathymetric highs formed barriers to the along shore transport of littoral sediments, a different pattern of coastal accretion, related to the development of beaches, has been detected. In both circumstances, the topset beds of the forced regressions appear to be greatly eroded and the intensification of the process is notable seawards. It is not possible
733
Marine Geology
to quantify the exact volume of the missing material and the depth to which this process reached. In relation to the present conditions along the Sicilian coasts, we believe that toplap erosion of the Late Quaternary - forced regressions was wave-induced and that the base level of erosion oscillated between -20 m and -15 m of water depth, a hypothesis supported by the modern base level depth. The Lpc related to the low sea level interval during the Last Glacial Maximum level (22 cal kyr to 19 cal kyr b.P.), is attached to the forced regression wedges. Steep seafloor gradients, and the lack of large inshore sediment storage, favoured beach erosion [41] and reduced the progradation potential [42]. So, the units, which show high-angle (more than 7°) prograding reflectors or acoustically transparent facies, can be interpreted as sand-grained infralittoral prisms, derived from coastal erosion by storm waves. Where the gradient of the shelves increased sharply, and the shelves are cut off by canyon-heads, which carry the materials towards the slope, imposing channelized turbiditic fans have been developed [17].
5.2
The development of transgressive parasequences
The TST was developed from the Last Glacial Maximum (about 18 kyr B.P.) to the present highstand (about 6 kyr b.P.). The initial phase of the transgression shifted the shoreline from the offlap break of the last lowstand deltas, at -130 m of water depth, and the first morphological step, at -110 m delimiting coastward the wavecut terrace at the top of FSST and Lpc. Correlations with several global and regional sea level curves relate this phase 734
with the Deglacial Onset during the Oldest Dryas, at about 15 kyr B.P. (18. 000 cal a B.P.). As relative sea level rose, submerging wide sectors of the outer continental shelves, the contemporaneous increase of the wave regime level led to the erasure of all previous positive morphology and to favour the accumulation of thin (15 m – 18 m) low-angle deposits immediately above the ravinement surface. Furthermore, the seaward movement of mud, which was deposited above the lowstand deltas, created wedge-shaped units, referred to as the healing phase. In the Egadi Islands and in the Selinunte – Sciacca offshore, under the influence of the paleo-flows of the Levantine Intermediate waters, shelf-edge contourite drifts were deposited. A second phase of the relative sea level rise shifted the shoreline from -75 m up to -60 m in depth, increasing the accommodation potential in the middle shelf. The second parasequence of TST has been developed during a short – term regressive phase (as confirmed by toplap/erosional surface and progradational patterns of the lower strata set) and a subsequent short – term transgressive phase. The direction of this progradation is always perpendicular to the coastline, while the depocenters follow a shore-parallel direction. Upper Pleistocene mid-shelf progradational bodies similar, in geometry and paleobathymetry [43, 44, 45, 46], to the intermediate transgressive parasequences described here, have been ascribed to the short-term Younger Dryas climatic event. Our hypothesis is confirmed by biostratigraphic data [32, 34], recovered in sediment cores reaching the deposits in the upper slope section of the Castellammare, Termini Imerese and Gela Basins.
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During the development of the upper TST parasequence in the early Holocene, the morphology of the studied areas affected wave refraction and the formation of longshore currents to transport coastal sands and influence the development of beaches and, landward, of associated wave-cut terraces.
forced regressions and lowstand prograding complex developed. Eustatic rise may have compensated for by subsidence and the shoreface retreat process occurred, creating thin transgressive lags on the wave-cut terraces and healing-phase deposits beyond the shelf break; 2. shelf to slope systems with an high degree of subsidence and high rate of sedi5.3 Role of fine â&#x20AC;&#x201C; grained sedment supply (as the southern Sicily mariment supply on the Highgin, Figure 7), where significant sedistand development mentary deposits occurred, both during the sea level falling and rising. LargeOver the last 6 kyr B.P., the areas fed by nuscale instabilities occurred, due to a merous mountainous rivers have been charcombination of high sedimentation rate acterized by the development of a thick during the glacial period, tectonic overprograding wedges, almost to the modern stepping of the margin and erosion at shelf edge, and by the progressive reducthe shelf edge by bottom currents; tion of the aggradation space for sediments 3. fault â&#x20AC;&#x201C; controlled systems (as the on the shelves. Palermo Mts. Offshore and the TerThe genesis of this unusually thick HST is mini Imerese Gulf, Figure 8), where sea tied to two factors: the high fluvial suplevel fall and lowstand promoted the inply of muds, which cover up to 80% of the cision of the Castellammare and Carini shelfal surfaces and a favourable sedimenCanyons head several kilometres into tary dispersion regime, in which wave acthe shelf through retrogressive masstion strongly influences river processes and wasting and favoured the development sediments are swept laterally away, from of mixed sand/mud turbiditic fans. the distributary mouths and deposited offshore, in NNE direction in the northern Sicilian margin and in SSE direction in the 6 Conclusions southern offshore. The high resolution sequence stratigraphic
5.4
Comparison between the analysis of the four areas of the Sicilian northern and southern margin has allowed documentation of upper Pleistocene and Holocene geological Sicily margin
The different shelf to slope systems described before allow us to differentiate three type of margin: 1. ramp systems with low rate of subsidence and sediment supply (as the Egadi offshore, Figure 8) where, during sea level fall and lowstand, attached
evolution, in particular on the stratigraphic and morphological effects of the Late Quaternary 4th and 5th order asymmetrical sea level cycles. The continental shelf and upper slope of the Sicily offshore are characterised by a 4th order type 1 depositional sequence, subdivided in a Falling Stage Systems 735
Marine Geology
Tract, a Lowstand Systems Tract, a Transgressive Systems Tract and a Highstand Systems Tract. Sedimentation takes place in a still tectonic active setting, in which regressions and transgressions occurred almost simultaneously with eustatic sea level falls and rises. During the FSST and the LST, paleophysiography, rate of the sedimentary supply and direction of the transport of the littoral sediments, controlled depositional systems differentiation and their asymmetrical distribution on the outer shelf. Prograding parasequences at the shelf margin, delimited at the base by the sequence boundary, represent the forced regression units, attached to the lowstand deltas. The TST, mainly developed in the middle and inner shelf, can be divided into three retrogradational parasequences, related to short - term episodes of cooling, such as the Younger Dryas, that modulated the post glacial sea level rise. The development of the TST was favoured by the high amount
of silt and mud, supplied by coastal rivers, and by the persistent wave-dominated dispersion regimes, which allowed deposition along the outer shelf. Upper Holocene HST prograded under the influence of river input and determined the development of depocenters up to 30 m thick. Gas accumulation, high rate of deposition and seismic events favoured the genesis of shelfal retrogressive failures, with limited downward displacement. The geological evolution of these areas reflects the variations in coastal environments during the different phases of the Late Quaternary cycle of relative sea level change. The depositional patterns, which characterise every systems tract, vary depending on sediment supply and dispersal on the shelf, the location of entry points and the variations in the accommodation space for the aggradation of shelf sediments. The last determines the site of the Systems Tracts in the different sectors of the shelf and upper slope.
References [1] F.J. Hilgen. Astronomical calibration of Gauss to Matuyama sapropels in the Mediterranean and implication for the geomagnetic polarity time scale. Earth and Planetary Science Letters, 104:226–244, 1991. [2] F.J. Hilgen. Extension of the astronomically calibrated (polarity) time scale to the Miocene/Pliocene boundary. Earth and Planetary Science Letters, 107:349–368, 1991. [3] C.G. Langereis and F.J. Hilgen. The Rossello composite: a Mediterranean and global reference section for the Early to early Late Pliocene. Earth Planet. Sci. Lett., 104:211–225, 1991. [4] L.J. Lourens, F.J. Hilgen, A. Antonarakou, A.A.M. Van Hoof, C. VergnaudGrazzini, and W.J. Zachariasse. Evaluation of the Pliocene to early Pleistocene astronomical time scale. Paleoceanography, 11:391–413, 1996.
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[5] F.J. Hilgen, W. Krijgsman, C.G. Langeris, L.J. Lourens, A. Santarelli, and W.J. Zachariasse. Breakthrough made in dating of the geological record. EOS, 78:285– 287, 1997. [6] R. Catalano, E. Di Stefano, S. Infuso, A. Sulli, P.R. Vail, and F.P. Vitale. Sequences and systems tracts calibrated by high-resolution bio-chronostratigraphy: the Central Mediterranean Plio-Pleistocene record. Society of Economic Paleontologists and Mineralogists Special Publication, (60):155–177, 1998. [7] J.D. Hays, J. Imbrie, and N.J. Shackleton. Variations in the earth orbit: pacemaker of the ice ages. Science, 194:1121–1131, 1976. [8] N.J. Shackleton and N.D. Opdyke. Oxygen isotope and paleomagnetic evidence for early Northern Hemisphere glaciation. Nature, 270:216–219, 1977. [9] L.D. Keigwin and R.C. Thunnel. Middle Pliocene climatic change in the Western Mediterranean from faunal and oxygen isotopic trends. Nature, 282:292–296, 1979. [10] W.F. Ruddiman, M.E. Raymo, and A. McInyre. Matuyama 41,000 –year cycles: North Atlantic Ocean and northern hemisphere ice sheets. Earth Planet. Sci. Lett., 80:117–129, 1986. [11] H.W. Posamentier, G.P. Allen, D.H. James, and M. Tesson. Forced regressions in a sequence stratigraphic framework: concepts, examples, and exploration significance. A.A.P.G. Bull., 76:1687–1709, 1992. [12] H.W. Posamentier and G.P. Allen. Variability of the sequence stratigraphic model, effects of local basin factors. Sediment. Geol., 86:91–109, 1993. [13] W. Helland-Hansen and J.G. Gjelberg. Conceptual basis and variability in sequence stratigraphy: a different perspective. Sediment. Geol., 92:31–52, 1994. [14] A.G. Plint and D. Nummedal. The falling stage systems tract: recognition and importance in sequence stratigraphic analysis. Geological Society Special Publication, 172:1–17, 2000. [15] B.D. Ritchie, R.L. Gawthorpe, and S. Hardy. Three – dimensional numerical modeling of deltaic depositional sequence 1: influence of the rate and magnitude of sea-level change. Journ. Sediment. Res., 74(2):203–220, 2004. [16] M. Agate, R. Catalano, S. Infuso, M. Lucido, L. Mirabile, and A. Sulli. Structural evolution of the northern Sicily continental margin during the Plio-Pleistocene. UNESCO Reports in Marine Science, 58:25–30, 1993. [17] M. Agate, A. D’Argenio, D. Di Maio, C. Lo Iacono, M. Lucido, M. Mancuso, and M. Scannavino. La dinamica sedimentaria dell’offshore della Sicilia Nordoccidentale durante il Tardo Quaternario. La Sicilia occidentale. Guida alle escursioni del 79° Congresso Nazionale della Societ`a Geologica Italiana, pages 157– 167, 1998. 737
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[18] M. Agate, M. Mancuso, and G. Lo Cicero. Late Quaternary sedimentary evolution of the Castellammare Gulf (North-western Sicily offshore). Boll. Soc. Geol. It., 124:21–40, 2005. [19] P. Tricart, L. Torelli, G. Brancolini, M. Croce, L. De Santis, D. Peis, and N. Zitellini. D´erive d’arcs insulaires et dynamique m´editerran´eenne suivant le transect Sardaigne – Afrique. C.R. Ac. Sc., Paris, II, 313:801–806, 1991. [20] M. Antonelli, R. Franciosi, G. Pezzi, A. Querci, G.P. Ronco, and F. Vezzani. Paleogeographic evolution and structural setting of the northern side of the Sicily Channel. Mem. Soc. Geol. It., 47:141–157, 1991. [21] R. Catalano, P. Di Stefano, A. Sulli, and F.P. Vitale. Paleogeography and structure of the Central Mediterranean: Sicily and its offshore area. Tectonophysics, 260:291– 323, 1996. [22] S. Pondrelli, S. Salimbeni, G. Ekstr¨om, A. Morelli, P. Gasperini, and G. Vannucci. The Italian CMT dataset from 1977 to the present. Phys. Earth Planet. Int., 159:286–303, 2006. [23] R. Catalano, B. D’Argenio, and L. Torelli. From Sardinia Channel to Sicily Strait. A geological section based on seismic and field data. Acc. Naz. dei Lincei, Atti dei Convegni Lincei, 80:109–127, 1989. [24] A. Fabbri, P. Gallignani, and N. Zitellini. Geological evolution of the perTyrrhenian sedimentary basins. Sedimentary basins of Mediterranean margins, pages 101–126, 1981. [25] R. Catalano and A. Milia. Late Pliocene – Early Pleistocene structural inversion in offshore Western Sicily. The Potential of Deep Seismic Profiling for Hydrocarbon Exploration, pages 445–449, 1990. [26] R. Catalano, P. Di Stefano, F. Nigro, and F.P. Vitale. Sicily mainland and its offshore: a structural comparison. UNESCO report in marine science, 58:19–24, 1993. [27] L. Ferranti, F. Antonioli, B. Mauz, A. Amorosi, G. Dai Pra, G. Mastronuzzi, C. Monaco, P. Orr`u, M. Pappalardo, U. Radtke, P. Renda, P. Romano, P. Sans`o, and V. Verrubbi. Markers of the last interglacial highstand along the coast of Italy: tectonic implications. Quaternary International, 145-146:30–54, 2006. [28] F. Antonioli, G. Cremona, F. Immordino, C. Puglisi, C. Romagnoli, S. Silenzi, E. Valpreda, and V. Verrubbi. New data on the Holocenic sea-level rise in NW Sicily (Central Mediterranean Sea). Global and Planetary Change, 34:121–140, 2002. [29] S. Pierini and A. Simioli. A wind-driven circulation model of the Tyrrhenian Sea area. Journ. Mar. Syst., 18:161–178, 1998.
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[30] A.R. Robinson and M. Golnaraghi. The physical and dynamical oceanography of the Mediterranean Sea. NATO – ASI Series C: Mathematical and Physical Scienses, 419:255–306, 1994. [31] E. Demirov and N. Pinardi. Simulation of the Mediterranean Sea circulation from 1979 to 1993: Part I. The interannual variability. Journ. Mar. Syst, 33-34:23–50, 2002. [32] G. Buccheri, O. Ferretti, M. Agate, M. Bertoldo, F. Immordino, and M. Lucido. Valutazioni stratigrafiche, sedimentologiche e paleoclimatiche sui sedimenti tardo pleistocenici-olocenici del golfo di Castellammare (Sicilia nord occidentale), indagini sulle carote PA-CA2 e PA-CA3. Boll. Soc. Geol. It., 117:219–248, 1998. [33] E. Di Stefano. Calcareus nannofossil quantitative biostratigraphy of holes 969E and 963B (Eastern Mediterranean). Proceedings of the Ocean Drilling Program, Scientif Results, 160:99–112, 1998. [34] E. Di Stefano and A. Incarbona. High-resolution paleoenvironmentale recostruction of ODP Hole 963D (Sicily Channel) during the last deglaciation based on calcareous nannofossils. Mar. Micropaleo., 52:241–254, 2004. [35] H. W. Posamentier and W. R. Morris. Aspects of the stratal architecture of forced regressive deposits. Geological Society, London, Spec. Publ., 172:19–46, 2000. [36] D. Nummedal and D.J.P. Swift. Transgressive stratigraphy at sequence-bounding unconformities, some principles derived from Holocene and Cretaceous examples. Society of Economic Paleontologists and Mineralogists Special Publication, 41:241–260, 1987. [37] W. Ruddiman and A. McIntyre. The North Atlantic Ocean during the last deglaciation. Palaeogeography, Palaeoclimatology, Palaeoecology, 35:145–214, 1981. [38] J. Imbrie, J.D. Hays, D. Martinson, A. McIntyre, A. Mix, J. Morley, N. Pisias, W. Prell, and N.J. Shackleton. The orbital theory of Pleistocene climate support from a recise chronology of the marine d18 O record. Milankovitch and climate, 1:269–305, 1984. [39] D.G. Martinson, N.G. Pisias, J.D. Hays, J. Imbrie, T.G. Jr. More, and N.J. Shackleton. Age dating and the orbital theory of the ice ages: development of a high resolution 0 to 300,000 year chronostratigraphy. Quaternary Research, 27:1–30, 1987. [40] R.G. Fairbanks. A 17,000 year glacio-eustatic sea level record: influence of glacial melting rates on the Younger Dryas event and deep ocean circulation. Nature, 342:637–642, 1989. [41] L.D. Wright, J.Chappell, B.G. Thom, M.P. Bradshaw, and P. Cowell. Morphodynamics of reflective and dissipative beach and inshore systems, South-western Australia. Mar. Geol., 32:105–140, 1979. 739
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[42] G.J. Orton and H.G. Reading. Variability of deltaic processes in terms of sediment supply, with particular emphasis on grain size. Sedimentology, 40:475–512, 1993. [43] F.J. Hern´adez-Molina, L. Somoza, J. Rey, and L. Pomar. Late Pleistocene – Holocene sediments on the Spanish continental shelves. Models for very highresolution sequence stratigraphy. Mar. Geol., 120:129–174, 1994. [44] A. Correggiari, M. Roveri, and F. Trincardi. Late Pleistocene and Holocene evolution on the north Adriatic sea. Il Quaternario, 9:697–704, 1996. [45] A. Cattaneo, F. Trincardi, and A. Asioli. Shelf sediments dispersal in the Late Quaternary transgressive record around the Tremiti High (Adriatic Sea). Giornale di Geologia, 1-2(59):217–244, 1997. [46] A. Cattaneo and F. Trincardi. The Late Quaternary transgressive record in the Adriatic epicontiental Sea: basin widening and facies partitioning. Society of Economic Paleontologists and Mineralogists Special Publication, (64):127–146, 1999.
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Paleomagnetism of Late Quaternary Sediment from the Eastern Tyrrhenian Sea as a Chronologic Tool for Marine Geology Investigations M. Iorio1 , J. Liddicoat2 , F. Budillon1 , R. Coe3 , A. Incoronato4 , D. Insinga1 , C. Lubritto5 , P. Tiano4 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, Department of Environmental Science, Barnard College, Columbia University, New York, USA 3, Earth Science Department, University of California, USA 4, Department of Earth Sciences, University of Napoli “Federico II”, Napoli, Italy 5, Department of Environmental Sciences, Second University of Napoli, SUN, Caserta, Italy marina.iorio@iamc.cnr.it Abstract Three piston cores (C1067, C1201, C1202) from the continental shelf and slope in the Salerno Gulf and Cilento offshore preserve small-scale, long-term change of Earth’s past magnetic field (Palaeo Secular Variation) between about 115 ka B.P and the present. The cores overlap as follows: C1201 spans the most recent 25 ka and overlaps C1067 and C1202, which respectively cover much of the last approximately 115 ka B.P. Cores C1067 and C1202 lack the record from about 20 ka to about 11 ka B.P., an absence that is due to slope erosional processes. The palaeomagnetic field behaviour, when correlated to curves of global relative palaeomagnetic field intensity and dated records of PSV for western Europe and Great Britain, coupled with discreet tephrochronology and geochronology dating (C14), allow us to compile a chronology for the combined cores. The resulting chronology for the first most recent 44 ka has application for placing time constraints on the marine geology and stratigraphy of the margin of the eastern Tyrrhenian Sea. So far, catastrophic events such as large-scale submarine slumps, volcanic eruptions, turbidite deposition, and abrupt changes in sedimentation rate have been documented. In particular, the changes in sedimentation rate seem to be linked to global rapid sea-level pulses and climatic events that induced concurrent reduction and/or abundance in the sediment supply from the adjacent coastal margin.
1
Introduction
cycle, are termed secular variation and are attributed to subsidiary features of the internal geomagnetic dynamo, generally exRecords of the geomagnetic field elements pressed by non-dipole terms of the total show that they vary on a wide range of field. Direct measurements of palaeomagtime scales. Variations longer than one netic secular variation (PSV) extend back year, excluding those related to sun-spot
Marine Geology
Figure 1: Morphobathimetry of Salerno and Palinuro Gulfs from Multibeam data sets (IAMC-CNR). Filled circles indicate the slope locations of C1201 [1], C1202 and C1067 (this study), PS88-19 [2]. All of the cores penetrated about 6 meters into the Holocene/Pleistocene deposits at about 300m below sea level. only for some 400 years, while palaeomagnetic studies provide the main source of information to extend the direct observations of the geomagnetic field variations to recent geological past. The detailed reconstruction of palaeomagnetic records retracing the Earthâ&#x20AC;&#x2122;s magnetic field variations, in addition to their implications for geomagnetism, have a valuable stratigraphic potential, since continuous PSV records may provide an original tool for high-resolution correlation and dating of recent sedimentary and volcanic sequences. So far high-resolution paleomagnetic records of PSV records have been obtained in sedimentary and volcanic sequence from several regions of the world (e.g. [3, 4, 5, 6, 7, 8, 9, 10, 11, 12, 13, 14, 15, 16, 17, 18, 19, 20, 1]). Also archaeological materials provide spot
742
readings of the past geomagnetic field, often of precise age, but they are discontinuous in both time and space. Still palaeomagnetic data from archeological material have been used to construct â&#x20AC;&#x2DC;master curvesâ&#x20AC;&#x2122; of regional palaeosecular variation such as those for Europe (e.g. [21, 10, 22]). The integration of different coeval PSV records from multiple sedimentary or volcanic sequences in the same region lead to the definition of regional PSV master curves (e.g. [4, 5, 8]) which are fundamental tools for the stratigraphic use of new individual PSV records. In this report we add to the data set for the Mediterranean region using a 6-m gravity core from about 250 m water depth on the Palinuro shelf in the Eastern Tyrrhenian Sea. In time and despite a sedimentary gap, the core spans all of the late Pleistocene
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and Holocene. The high-resolution investigation presented here combines paleomagnetic and petrophysical data with the objective of compiling a regional chronologic framework that has application in sedimentology, tephrochronology, and palaeoclimate research.
an important stratigraphic marker in marine cores in the region, and can be recognized on seismic data for the continental shelf and the upper continental slope [26, 31, 18, 1]. Finally recent intense tectonic deformation between Punta Licosa and Cape Palinuro steepened depositional slopes and, in combination with relative sea level variations, seismic activity and sporadic overload of volcanic deposits on 2 Marine geology the margin, caused slope failures and mass The study area encompasses the continen- transport on the margin [32, 2, 33]. tal margin between Cape Palinuro and the Sorrento Peninsula and contains two different physiographic settings that reflect two 3 Palaeomagnetic and distinct tectonic areas, one a structural desedimentological procepression that trends NE-SW bounded by normal faults and another that is a strucdures tural high controlled by the seaward extension of Cilento units (Figure 1). 3.1 Sampling In the studied area, recent tectonics, global sea-level oscillations, and variable sedi- Core C1067, C1201, and C1202 have a diment input have resulted in at least 2000ms ameter of 9 cm and were taken from the of Plio-Quaternary deposits in the Salerno upper continental slope in the Salerno Gulf Gulf and about 600ms in the Cilento off- and Punta Licosa margin at 40º 08.3898’ N, shore [24, 25]. The continental shelf is 14º 43.6212’ E, 40º 28.918’ N, 14º 42.236’ characterized by a late Pleistocene pro- E, and 40º 08.3393’ N, 14º 43.5672’ E, grading wedge that is unconformably over- respectively (Figure 1). The cores penelain by a postglacial drape [26, 2, 27]. The trated 4.5, 6.4, and 6 m of late Pleistoceneprograding stacking pattern in which beds Holocene sediment between about 250 to dip steeply to the south and west is the 300 m of water; the penetration was vertiresult of the gradual drop of sea level of cal as indicated by horizontal layering near about 120 m with respect to present sea the core base. At the Marine Coastal Envilevel [28, 29]. Thus, an erosional uncon- romental Instistute (IAMC) of Naples, Naformity separates the regressive unit from tional Research Council (CNR), Italy, the an overlying one with backstepping ge- cores, which were not oriented in azimuth, ometries that formed between about 18000 were cut into 1-m segments, split lengthand 5000 years ago when sea-level rose to wise, and stored horizontally at 4°C. The nearly the present position. The deposi- working half of cores C1202 and C1201 tional unit that formed during the later half were sub-sampled continuously throughout of the Holocene is wedge-shaped and de- their lengths for palaeomagnetic analysis creases in thickness toward the shelf edge. and put in PVC boxes 2 cm on a side; the A distinctive pumice layer deposited af- separation between samples was 1 cm or ter the 79 A.D. Plinian eruption [30] is less and all the laboratory analyses were 743
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completed within 15 months of core recovery. However in the uppermost 50 cm of core C1202 two intervals of 18 and 12 cm, characterized by marine mud with a high water content and tephra layer respectively, did not allowed undisturbed discrete palaeomagnetic sampling. The remaining (archive) half of the cores was used for continuous high-resolution sedimentological and physical properties analyses at 2cm intervals, respectively. Physical properties were measured almost immediately after coring to avoid drying or oxidation that might cause an alteration in the geochemical and mineralogical properties of the sediment.
3.2
Sedimentology
The lithostratigraphic observations on all cores were performed at the cm scale (Figure 2A, column I to III). The sediment is undisturbed hemipelagic mud in the upper part of the cores, which passes to a fine-grained sand characterised by an alternation of intervals with more or less muddy matrix in the remainer of the cores. Considered as a whole, the coarse fraction intervals probably record the Last Glacial Maximum (LGM), during which the energy of environmental dynamics increased. Several ash layers are intercalated in the three cores at different depths, and they were sourced mostly by the SommaVesuvius complex and volcanoes on Ischia. The most important volcanic layer is characterized by a 0.60-m coarse-grained pumice in C1201, attributed to the 79 A.D. Somma-Vesuvius eruption [1] and reference therein). Other than the disruption of the deposition of mud and silt on the continental shelf caused by the 79 A.D. eruption of Somma-Vesuvius, there are no marked erosional disconformities in C1201, while 744
in cores C1202 and C1067 a lithological discontinuity is present at about 0.75 and 0.60 mbsf, marking a stratigraphic hiatus. At the INNOVA CIRCE Laboratory of Second University of Naples (Caserta), 14 C (AMS) dates were obtained for foraminifera and mollusc shells that give calibrated [34] dates of cal. 24,903 ±463 ka B.P. in C1201, and cal. 6560 ±30 ka B.P and 43,900±300ka B.P for C1202.
3.3
Physical properties analyses
The physical properties analyses (Gamma Ray Density, P-wave velocity, Color Reflectance, not presented here, and Volume Magnetic Susceptibility, VMS) of the archive half of all cores were measured in a fully automated GEOTEK Multi-Sensor Core Logger (MSCL) at the IAMC-CNR Physical Properties Laboratory. A Bartington MS2E Point sensor was used to measure the low-field Volume Magnetic Susceptibility. The physical properties of the sediment were measured on the cm scale and the VMS for all cores is plotted in Figure 2B (columns V to VII). Mean VMS values is about 40 SI x10−5 for C1201 and C1202, and 30 SI x10−5 for core C1067, with the exception of several high peak values (the highest reaching 720 SI x 10−5 in C1201). In C1201, three of the peaks (at depths 0.10, 0.45,and 1.8 m) correspond to tephra layers V0, V1, and V3, (SommaVesuvius eruptions) [1] respectively. In C1202 and C1067, the two peaks at 0.2 correspond at tephra layer V2 (79A.D. Vesuvius eruption), while the peak at 4.9 mbsf in C1202 corresponds to tephra layer V6 (correlated to Ischia activity) (Figure 2B, Columns V toVII).
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Figure 2: A) Sedimentological core logs, 14 C datings and correlation for (I) C1201 [1], (II) C1202 and C1067 (this study), and (III) PS88-19 [2]. Depth scale: metres below sea floor (mbsf), legend as indicated in the figures. V0, V1, V2, V3 and V6 are tephra layers indentified by VMS and geochemical data [1], and references therein and this study). Labels on right side of PS88-19 indicate seismic units interpretation as distinguished by Trincardi et al., [2]. B) Continuous records of Volume magnetic susceptibilities for C1201 [1], C1202, C1067 and PS88-19 [2]. Where the cores overlap, dotted lines indicate the correlated parameter behaviours; correlated peaks are numbered 1 to 5 (see text for explanation). Depth is in metres below sea floor (mbsf). VMS values are plotted on a logarithmic scale. 745
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Figure 3: Representative IRM acquisition and NRM demagnetization behaviour plots for selected specimens. (A) Zijderveld diagram for a.f. demagnetization to 100 mT for samples 1201 SE 39 [23]. The horizontal and vertical components are indicated as square and triangle symbols, respectively. The demagnetization step is for peak applied field (mT); intensity scale is indicated as a bar. (B) Isothermal Remanent Magnetization (IRM) acquisition curve for sample 1202 SA32.
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3.4
Palaeomagnetism and rock of about 500 years. The time window was chosen knowing that in the sedimenmagnetism analysis
The palaeomagnetic directions (inclination and relative declination) and remanence of more than 400 discrete samples from core C1201 and C1202 were measured at the Palaeomagnetism Laboratory of the University of California, Santa Cruz, in a 2G superconducting magnetometer that has a noise level of 1 x 10−8 A/m. The samples were analysed for Anhysteretic Remanent Magnetization (ARM) and Isothermal Remanent Magnetization (IRM) in the Palaeomagnetism Laboratory of the University of Naples using a Molspin AF demagnetizer and Minispin magnetometer. The average intensity of the natural remanent magnetization (NRM) for cores C1201 and C1202 is 22.9 and 12.3 mA/m, respectively. It was found that for all the pilot samples there is high magnetic stability and the dominance of a single component remanence identified using the principal component analysis of Kirschvink [37] (Figure 3A). Saturation of the Isothermal Remanent Magnetization (SIRM) was attained by low fields (less than 0.25 T) in representative samples, indicating that the main magnetic carrier is titano-magnetite (Figure 3 B). On the basis of the demagnetization profiles for the pilot samples, the Characteristic Remanent Magnetization (ChRM) was considered to be represented at 30 mT (Figure 4, columns I and III for C1201, from Iorio et al., [1]). In particular, the mean inclination record averages of 57.4° and 64.9° for C1201 [1] and C1202 (interval 0-44 ka, this study) respectively, which is close to the inclination of an axial dipole field (59.9°). In order to reduce the noise of the palaeomagnetic directions, all the data were statistically filtered with a smoothing degree
tation environment where C1201 was collected, the time lag in the palaeomagnetic lock-in depth has a maximum value of about 500 years [38]. Knowing the magnetic homogeneity along the sedimentary sequence, relative palaeointensity curves (RPI) were obtained by normalizing the NRM intensity after 30mT demagnetization (NRM30mT) by the intensity of the ARM at the same demagnetization step (NRM30m/ARM30mT) [39], (Figure 3B, C1202 this study. Figure 4, Column V, C1201 from Iorio et al [1]).
4 4.1
Discussion and interpretation Lithologic Correlation
The sedimentological and physical properties data for C1201, C1202 and C1067 were visually compared (Figures 2A and 2B) with similar published data from the Punta Licosa upper continental slope (Core PS88-19 from Trincardi et al., [2]) and are described below. Core PS88-19 is also from the upper continental slope of Punta Licosa and is believed to be within 500 m of C1202, taking into account the navigation precision that was used to fix the localities. The core is mainly a muddy, fine-grained sand and contains the 79 A.D. Somma-Vesuvius pumice layer at the same stratigraphic position as in core C1202 and C1067. It is important to note, however, that PS88-19, as well as C1067, do not contain the complete record of tephra layer V6. So far tephra layer V2, attributed to the 79 A.D. eruption [26, 2, 40, 1] and reference therein, Budillon et al., in prep.) is 747
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Figure 4: C1201 Age model. Smoothed curves of relative paleomagnetic declinations and inclinations for C1201 versus time (Columns I and III) and correlation (dotted lines and marked features 10 to1, l to b and v to b ) with the Britain type curves ([4, 5] recalibrated and renamed according to Vigliotti [19]) and Lac du Bouchet PSV curves [8] (columns II and IV). Correlation between C1201 smoothed relative paleointensity curve (Column V) and relative paleointensity NRM/ARM curve from Lac du Bouchet [35] curves (Column VI), which is available for approximate ages greater than c. 11000 years B.P. Correlated paleointensity features 1 to 10 (Modified from Iorio et al.,[1]). present in all cores. Note in Figure 2A that for C1201 there are 2.5 m between tephra layer V2 and the start of coarse fraction at about 4 m, whereas there are only 0.50 m in all three cores. This finding is consistent with a stratigaphic hiatus due to sediment failure [2]. Continuing down in the cores there is a coarse fraction interval between 4 to 6 mbsf for C1201 and 0.65 to 2 mbsf for cores C1202, C1067 and PS8819 as determined by means of high resolution VMS data (Figure 2B, Peaks 1 to 5) such an interval is ascribed mainly to the LGM by Trincardi et al. [2] and Buccheri et al. [26]. Continuing down the Punta Licosa cores, the sediments are intercalated by three ash layers (Figure 2A columns II), the most important constituted by a decimetric layer (V6), characterized by light brown pumice, which has been correlated
748
to Ischia activity on the basis of geochemical features (Budillon et al., in prep.).
4.2
Age model, comparision with PSV master curves and Slide Time Gap
The C1201 age model is presented in Iorio et al. [1] with the same methodology described below for core C1202 and is not discussed here. For C1202, knowing that its top (0.2 mbsf) was assigned an age of about 2000 years B.P., as suggested by the presence of V2 (Budillon et al., in prep.) and that there was a sedimentary gap as suggested by seismic evidences and previous study in the area [2] and using two chronologic tie points (radiocarbon dates), it was possible to place calibration points (indicated by break line arrows in Figure
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5) along the record for C1202 to obtain a depthâ&#x20AC;&#x201C;age transformation by linear interpolation; the transformation is done assuming, as a first approximation, that there was a constant sedimentation rate between the sequential pairs of calibrated points. The average sedimentation rate in the two intervals below and above the sedimentary gap was of 13 and 7 cm/ka respectively, the different sedimentation rate obtained, is in agreement with a slope deposition during the respective low stand and transgressive sea level stages. Subsequently, the smoothed palaeomagnetic parameters were plotted versus the computed time and then compared for this time interval with the declination and inclination curves for France and Italy [8, 13]. When examined visually, the large-scale fluctuations of paleomagnetic directions in C1202 are similar to those in the reference curves. Subsequently, correlated directional features between our data and the reference curves were used to further refine the C1202 age model, assigning to each paleomagnetic feature recognised in our core the relative referring curve ages (Iorio et al., in prep.). By knowing the time interval spanned by the sediment in C1201 and C1202 , the relative paleomagnetic intensity curve for each core, and the Pleistocene palaeointensity master curve for the North Atlantic Ocean (NAPIS, [36]) can be compared (Figure 5). The comparison shows a good match, with distinct features in the interval 20-44 ka, that do not differ in age by more than Âą500 years for C1202 with respect to the NAPIS master curve. On the base of paleointensity curves and VMS data correlations, 14 C datings, coherence in sedimentation rates, and seismostratigraphic evidences [2], the time gap in the C1202 record is suggested to span from about 20 to 11 ka B.P.
Moreover, as the portion of the C1202 paleointensity record, which overlaps with the NAPIS curve, shows quite good agreement, it seems that these data, if confirmed by further analysis, document the relative palaeointensity for the late Pleistocene in the Eastern Tyrrhenian region.
4.3
Global control on slope sedimentation
The sedimentation rate established by the PSV dating for the most continuous core C1201 [1] is presented in Figure 6. Most of the well-marked changes in sedimentation rate seem to be synchronous with palaeoclimatic events and sea level Melting Water Pulses (MWP and MWP1A) of late Pleistocene, that were recognised in previous studies of sediments from the Atlantic Ocean and Mediterranean Sea [41, 42, 43] and through a palaeoclimatic study of Salerno Gulf slope sediment by Buccheri et al. [26]. In Figure 6, the occurrence of the late Pleistocene/Holocene climatic episodes and sea level oscillations reported by the above authors are plotted against the sedimentation rate curve established for core C1201 by Iorio et al. [1]. In particular, the sedimentation trends in the late Pleistocene record show remarkably low sedimentation rates of 8.0, and 11.0 cm/ka in the intervals between 2010019000 and 14800-12700 years B.P., which were interpreted by Iorio et al. [1] mainly as driven by the MWP and MWP1A sea level rises. In fact, in the Salerno Gulf, a rise of about 15m ([42] could correspond to a retrogressive shift of the coastline of about 1.5 km [29], with a consequent relative low sediment supply and low sedimentation rate. Late Pleistocene small but abrupt pulses 749
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of high sedimentation (Figure 6) were observed at 17200, 16200 and 15000 years B.P. that might correlate with the oscillating features in Heinrich Event 1 [1]. A longer period of sediment rate increase is observed between 12700â&#x20AC;&#x201C;11700 years B.P., reaching values of about 30 cm/ka; this time interval is coincident with the Younger Dryas cold period (12600-11500 years B.P. [44]), when the rate of global sea level rise decrease to about 3mm/yr [43]. At the transition from the late Pleistocene to the Holocene, an abrupt increase in the sedimentation rate to about 300 cm/ka also occurred in C1201 as evidenced by Iorio et al. [1], that is followed by an interval of slow sedimentation rate (mean value 12 cm/ka) that marks the beginning of the Holocene and spans up to about 4000 years B.P. (Figure 6). Such a low sedimentation rate remains fairly constant also during the time interval coincident with the deposition of sapropel S1 (9800-6500 years B.P.) [45] that occurred during a well-known warmer period in the Mediterranean region [46], suggesting that the sedimentation on the slope was not affected by this climatic event. Finally, from about 4000 years B.P. to 2000 years B.P., a further high increment in the sedimentation rate, with an average value of 53 cm/ka occurs in the record.
5
Conclusions
The sedimentologic, palaeomagnetic, 14 C data and physical properties for C1202
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and C1067, combined with similar records for C1201 and PS88-19, provide a PSV geochronology of the studied area that applies to geological and palaeoclimatic studies. In particular, in the offshore of Punta Licosa, it was possible to refine the dating of the large sediment removals that are due to a submarine slide along the prograding surfaces [2]. As well, shortterm sedimentation-rate fluctuations were detected by PSV chronology, and it was observed that in the late Pleistocene during periods of generally low sea level, pulses of melt water had a direct affect on the continental slope by reducing the rate of sediment accumulation, while alternating stadial and interstadial climatic periods in the same time interval also appear to affect the rate of deposition, with an increase in the sedimentation rate during cold phases. However in the Holocene, when the relative sea level increase has been continuous, a generally low sedimentation rate is measured for the continental slope, and the affect of warm periods, as during the S1 interval, is masked by the predominance of sea-level control, as is seen by the lack of variation in the sedimentation rate at the end of the S1 interval. Finally, it is believed that these high-resolution chronologies can be applied to future sedimentologic and paleoclimatic studies in the Southern Campania margin.
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Figure 5: C1202 Age Model and Slide Time Gap. (I and II) Relative paleointensity (RPI) smoothed curves for C1201 [1] and C1202 (this study) versus time and correlated paleointensity features (black line arrows) with the (III) NAPIS Master Curve [36]. Note that for C1202, the Age Model shown regards only the first 4.0 mbsf. From the correlation it is possible to date for core C1202 the time gap in the sedimentological record, which is due to an erosional hiatus originated by a large scale slide [2]. The slide gap spans from about 11000 to 20000 years B.P.
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Figure 6: Sedimentation rate changes for the continental slope core C1201 computed with PSV chronology (modified by Iorio et al., [1]). Numbers along the curve indicate sedimentation rates in cm/ka unit. In the C1201 PSV chronology curve, squares indicate the five calibration tie points used to anchor our PSV data to PSV European reference curves, diamonds indicate the 25 calibration points obtained after correlation with such reference curves. The ages of the main pulses and changes in the sedimentologic slope rate curve, that coincide with the ages of palaeoclimatic and sea level oscillation events noted in the text have been indicated: at 20100 years B.P. the Water Melt Pulse (WMP), between 17100 to 15000 (H1), the Cool Heinrich Event 1 oscillations and between 12600-11500 years B.P., the Cool Younger-Dryas (YD) event. At about 14600 years B.P., the Water Melt Pulse 1A (WMP1A), and between 14300 to 13200 years 752 B.P. The B¨olling-Aller¨od interstade (B-A), is marked. Finally, at 11700 years B.P. and between 9800 to 6500 years B.P., the Holocene/Pleistocene boundary (H/P) and the Mediterranean warmer interval (S1) are indicated, respectively.
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References [1] M. Iorio, J. Liddicoat, F. Budillon, P. Tiano, A. Incoronato, R. Coe, and E. Marsella. Palaeomagnetic Secular Variation Time Costrain on Late Neogene Geological Events in slope sediment from the Eastern Tyrrhenian Sea. SEPM (Society for Sedimentary Geology) Spec. Pub., 92:233–246, 2009. [2] F. Trincardi, A. Cattaneo, A. Correggiari, S. Mongardi, A. Breda, and A. Asioli. Submarine slides during relative sea level rise: two examples from the eastern Tyrrhenian margin. pages 469–478, 2003. [3] K.M. Creer, T.E. Hogg, P.W. Readman, and C. Reynaud. Palaeomagnetic secular variation curves extending back to 13.400 years B. P. recorded by sediments deposited in Lac du Joux, Switzerland. Journal of Geophysics, 48:139–147, 1980. [4] G.M. Turner and R. Thompson. Lake sediment record of the geomagnetic secular variation in Britain during the Holocene times. Geophysical Journal Rojal Astronomical Society, 65:703–725, 1981. [5] G.M. Turner and R. Thompson. The transformation of the British geomagnetic secular variation record for Holocene times. Geophysical Journal Rojal Astronomical Society, 70:789–792, 1982. [6] S.P. Lund and S.K. Banerjee. Late Quaternary paleomagnetic field secular variation from two Minnesota lakes. Journal of Geophysical Research, 90:803–825, 1985. [7] J.-C. Tanguy, I. Bucur, and J.F.C. Thompson. Geomagnetic secular variation in Sicily and revised ages of historic lavas from Mount Etna. Nature, 318:453–455, 1985. [8] N. Thouveny, K.M. Creer, , and I. Blunk. Extension of the Lac du Bouchet paleomagnetic record over the last 120,000 years. Earth and Planetary Science Letters, 97:140–161, 1990. [9] D.L Pair, E.H. Muller, and P.W. Plumley. Correlation of Late Pleistocene glaciolacustrine and marine deposits by means of geomagnetic secular variation, with examples from northern New York and southern Ontario. Quaternary Research, 42:277–287, 1994. [10] P. M`arton. Archeomagnetic directions: The Hungarian calibration curve. in “Palaeomagnetism and Tectonics of the Mediterranean Region”: Geological Society of London, Special Publication, 105:385–399, 1996. [11] J.C. Liddicoat and R.S. Coe. Paleomagnetic investigation of Lake Lahontan sediments and its application for dating pluvial events in the northwestern Great Basin. Quaternary Research, 47:45–53, 1997.
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[12] J.C. Liddicoat and R.S. Coe. Paleomagnetic investigation of the Bonneville Alloformation, Lake Bonneville, Utah. Quaternary Research, 50:214–220, 1998. [13] U. Brandt, N.R. Nowaczyk, A. Ramrath, A. Brauer, J. Mingram, S. Wulf, and J.F.W. Negendank. Palaeomagnetism of Holocene and Late Pleistocene sediments from Lago di Mezzano and Lago Grande di Monticchio (Italy): initial results. Quaternary Science Reviews, 18:961–976, 1999. [14] U. Frank, N.R. Nowaczyk, J.F.W. Negendank, and M. Melles. A paleomagnetic record from Lake Lama, northern Central Siberia. Phys. Earth Planet. Inter.,, 133:3–20, 2002. [15] J.T. Hangstrum and D.E. Champion. A Holocene paleosecular variation record from 14 C-dated volcanic rocks in western North America. J. Geophys. Res. doi: 10.29/2001JB000524, 107(B1)(2025), 2002. [16] A. Incoronato, A. Angelino, R. Romano, A. Ferrante, R. Sauna, G. Vanacore, and C. Vecchione. Retrieving geomagnetic secular variations from lava Flows: evidence from Mount Arso, Etna and Vesuvius (Southern Italy). Geophysical Journal International, 149:724–730, 2002. [17] G. St-Onge, J.S. Stoner, and C. Hillaire-Marcel. Holocene paleomagnetic records from the St. Lawrence Estuary, eastern Canada: centennial- to millennial-scale geomagnetic modulation of cosmogenic isotopes. Earth and Planetary Sci. Lett., 209:113–130, 2003. [18] M. Iorio, L. Sagnotti, A. Angelino, F. Budillon, B. D’Argenio, J. Dinar`es-Turell, P. Macr`ı, and E. Marsella. High-resolution physical properties and palaeomagnetic study of late-Holocene shelf sediments, Salerno Gulf, Tyrrhenian Sea. Holocene, 14:426–435, 2004. [19] L. Vigliotti. Secular variation record of the Earth’s magnetic field in Italy during the Holocene:constraints for the construction of a master curve. Geophysical Journal International, 165:414–429, 2006. [20] I. Snowball, L. Zill´en, A. Ojala, T. Saarinen, and P. Sandgren. FENNOSTACK and FENNORPIS: Varve dated Holocene palaeomagnetic secular variation and relative palaeointensity stacks for Fennoscandia. Earth and Planetary Sci. Lett., 255:106–116, 2007. [21] I. Bucur. The direction of the terrestrial magnetic. Field in France during the last 21 centuries, recent progress. Physics of the Earth and Planetary Interiors, 87:95–109, 1994. [22] U. Frank, N.R. Nowaczyk, A. Ramrath, A. Brauer, J. Mingram, S. Wulf, and J.F.W. Negendank. Palaeomagnetism of Holocene and Late Pleistocene sediments from Lago di Mezzano and Lago Grande di Monticchio (Italy): initial results. Quaternary ScienceR eviews, 18:961–76, 1999. 754
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[23] J.D.A. Zijderveld. A.C. demagnetization of rocks: analysis of results, in: Methods in paleomagnetism. pages 254–286, 1967. [24] R. Bartole, C. Savelli, M. Tramontana, and F.C. Wezel. Structural and sedimentary features in the Tyrrhenian margin off Campania, Southern Italy. Marine Geology, 55:163–180, 1984. [25] M. Sacchi, S. Infuso, and E. Marsella. Late Pliocene–early Pleistocene compressional tectonics in offshore Campania. Bollettino di Geofisica Teorica ed Applicata, 36:469–482, 1994. [26] G. Buccheri, G. Capretto, V. Di Donato, P. Esposito, G. Ferruzza, T. Pescatore, E. Russo Ermolli, M.R. Senatore, M. Sprovieri, M. Bertoldo, D. Carella, and G. Madonia. A high resolution record of the last deglaciation in the southern Tyrrhenian Sea environmental and climatic evolution. Marine Geology, 186:447–470, 2002. [27] ISPRA. Carta Geologica d’Italia 1:50000, Foce del Sele, n.° 486. 2009. [28] F. Trincardi and M. Field. Geometry lateral variation and preservation of downlapping regressive shelf deposits: Eastern Tyrrhenian Sea margin, Italy. . Journal of Sedimentary Petrol., 61(5):775–790, 1991. [29] F. Budillon, T. Pescatore, and M. R. Senatore. Cicli deposizionali del Pleistocene Superiore – Olocene sulla piattaforma continentale del Golfo di Salerno (Tirreno Meridionale). Bolettino Societ`a Geologica Italiana, 113:303–316, 1994. [30] L. Lirer, T. Pescatore, B. Booth, and G.P.L. Walzer. Two Plinian pumice-fall deposits from Somma–Vesuvius, Italy. Geological Society of American Bulletin, 84:759–772, 1973. [31] M. Sacchi, D. Insinga, A. Milia, F. Molisso, A. Raspini, M. Torrente, and A. Conforti. Stratigraphic signature of the Vesuvius 79 AD event off the Sarno prodelta system, Naples Bay. Marine Geology, 222-223:443–469, 2005. [32] F. Trincardi and M. Field. Collapse and flow of lowstand shelf–margin deposits: An example from the eastern Tyrrhenian sea, Italy. Marine Geology, 105:77–94, 1992. [33] A. Bellonia, F. Budillon, F. Trincardi, D. Insinga, M. Iorio, A. Asioli, and E. Marsella. Licosa and Acciaroli submarine slides, Eastern Tyrrhenian margin:characterisation of a possible common weak layer. Ext abs.Soc Geol. It., 2:142– 143, 2008. [34] M. Stuiver and P.J. Reimer. Radiocarbon Calibration program. Radiocarbon, 35:215–230, 1993.
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[35] N. Thouveny. Variations of the relative paleointensity of the geomagnetic field in western Europe in the interval 25–10 kyr BP as deduced from analyses of lake sediments. Royal Astronomical Society, Geophysical Journal, 91:123–142, 1987. [36] J.S. Stoner, C. Laj, J.E.T. Channell, and C. Kissel. South Atlantic and North Atlantic geomagnetic paleointensity stacks (0–80 ka): implications for interhemispheric correlation. Quaternary Science Reviews, 21:1141–1151, 2002. [37] J.L. Kirschvink. The least squares line and plane and the analysis of palaeomagnetic data. Geophysical Journal Royal Astronomical Society, 62:669–718, 1980. [38] L. Sagnotti, F. Budillon, J. Dinar`es-Turell, M. Iorio, and P. Macr`ı. Evidence for a variable paleomagnetic lock-in depth in the Holocene sequence from the Salerno Gulf (Italy): Implications for “high-resolution” paleomagnetic dating. Geochemistry, Geophysics, Geosystems. doi: 10.1029/20056C001043, 6(11):1–11, 2005. [39] S. Levi and S.K. Banerjee. On the possibility of obtaining palaeointensities from lake sediments. Earth Planet. Sci. Lett., 29:219 226, 1976. [40] D. Insinga, F. Molisso, C. Lubritto, M. Sacchi, I. Passariello, and V. Morra. The Proximal Marine Record of Somma-Vesuvius Volcanic Activity, in the Naples and Salerno Bays, Eastern Tyrrhenian Sea, during the last 3 Kyrs. Journal of Volcanology and Geothermal Research, 177:170–186, 2008. [41] I. Cacho, J.O. Grimalt, and M. Canals. Response of the Western Mediterranean Sea to rapid climatic variability during the last 50,000 years: a molecular biomarker approach. Journal of Marine Systems, 33-34:253–272, 2002. [42] P.U. Clark, A.M. McCabe, A.C. Mix, and A.J. Weaver. Rapid Rise of Sea Level 19,000 Years Ago and Its Global Implications. Science, 304:1141–1144, 2004. [43] A.J. Weaver, O.A. Saenko, P.U. Clark, and J.X. Mitrovica. Meltwater Pulse 1A from Antarctica as a Trigger of the Bølling –Allerød Warm Interval. Science, 299:1709– 1713, 2003. [44] A. Asioli, F. Trincardi, J.J. Lowe, D. Ariztegui, L. Langone, and F. Oldfield. Submillennial scale climatic oscillations in the central Adriatic during the Lateglacial. Paleoceanographic implications. Quarternary Science Reviews, 20:1201–1221, 2001. [45] J.S.L. Casford, E.J. Rohling, R. Abu-Zied, S. Cooke, C. Fontanier, M. Leng, and V. Lykousis. Circulation changes and nutrient concentrations in the late Quaternary Aegean Sea: A nonsteady state concept for sapropel formation. Paleoceanography doi: 10.1029/200PA00601, 17:1024, 2002. [46] D. Ariztegui, A. Asioli, J.J. Low, F. Trincardi, L. Vigliotti, F. Tamburini, C. Chondrogianni, C.A. Accorsi, A.M. Bandini Mazzanti, M. Mercuri, S. Van der Kaars, J.A. McKenzie, and F. Oldfield. ”Palaeoclimate and the formation of sapropel S1: 756
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inferences from Late Quaternary lacustrine and marine sequences in the central Mediterranean regionâ&#x20AC;?. Palaegeogr. Palaeoclim. Palaeoecol., 158:215â&#x20AC;&#x201C;240, 2000.
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Middle Mesozoic to Present Geological Evolution of the Central-Northern Sector of the Falkland/Malvinas Plateau as Inferred from Seismic Data E. Petruccione1 , A. Tassone2 , G. Nardi3 , E. Lodolo4 , M. Iorio1 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, Institute of Geophysics “Daniel A. Valencio”, Department of Geological Sciences, University of Buenos Aires, Buenos Aires, Argentina 3, Department of Earth Sciences, University of Napoli “Federico II”, Napoli, Italy 4, National Institute of Oceanography and Experimental Geophysics, Trieste, Italy marina.iorio@iamc.cnr.it Abstract The sequence stratigraphy is a methodology that, by means of the recognition of depositional sequences in stratigraphic architecture, reconstructs tectonic phases and relationships between tectonic, sedimentation and eustasy. A seismo-stratigraphic study of the northern sector of the Falkland Plateau (FMP) is derived from integrated analysis of two unpublished seismic reflection profiles combined with published seismic profiles, bathymetric and well data. Data analysis allowed the identification of an acoustic basement, four seismic sequences (A to D) and their bounding unconformities (r1, r2, r3 and r4), aged from middle Jurassic to Present. The precambric acoustic basement, bounded by r1 unconformity, overlain by the Middle-Late Jurassic syn-rift continental deposits of sequence A, which is bounded at the top by the r2 unconformity, testify a tectonic behaviour of uplift around the Jurassic/Cretaceous boundary. The overlying sequences B and C represent the post-rift phase. From the late Cretaceous up to Cretaceous/Paleogene boundary, a widespread erosion in the northern part of the FMP is suggested by the r4 unconformity cutting the underlying sequences C and B. The erosion is probably due to a southward tilting linked to underthrusting of the southern sector of the FMP. The lower and constant thickness of sequence D suggests, during the Cenozoic, a slow rate of subsidence, probably linked to the contemporaneous stop of the subduction of the FMP under the Scotia Plate.
1
Introduction
The sequence stratigraphy is based on the principle according to which the sedimentary successions are interpreted as the result of the interplay of three independent variables: eustatism, tectonic (uplift and
subsidence) and sediment supply. This interaction leads to cyclical variations in sea level, the stratigraphic evidences of which are sedimentary bodies or depositional sequences, bounded at the bottom and the top by unconformity and correlative conformity (sequence boundaries), considered
Marine Geology
Figure 1: Bathymetric map and localization of analyzed unpublished (continuous lines) and published (black dot lines) seismic profiles (References listed in [1]). The blue squares indicate the position of DSDP Leg 36 and Leg 71 wells. In red the portions of seismic profiles shown in Figures 2 and 3. AB: Argentine Basin, ACS: Argentine Continental Shelf, FE: Falkland Escarpment, FI: Falkland Islands, FPB: Falkland Plateau Basin, FMP: Falkland-Malvinas Plateau, FT: Falkland Trough, MEB: Maurice Ewing Bank, NSR: North Scotia Ridge. to be synchronous at global scale. These surfaces are easily recognizable in the seismic sections, where they appear as discrepancies between geometric surfaces and lateral terminations of seismic horizons. In offshore areas, using the principles of sequence stratigraphy for the interpretation of seismic reflection profiles, is possible, with the integration of well data, recognize and reconstruct the depositional sequences and depocenter and their migration in the time and space, allowing kinematic reconstruction of the tectonic structures, with the relative growth rates, and then establishing a chronology of geological events.
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For the study described in this paper, a comparative analysis, based on two unpublished seismic profiles, published seismic reflection and refraction profiles, bathymetric and well data, was carried out, aimed to identify seismic-stratigraphic sequences and main unconformities correlative in the central-northern Falkland Plateau Basin (FPB) in order to determine its evolution history during the main phases of the South Atlantic Ocean opening. The FPB is a portion of Falkland/Malvinas Plateau (FMP), which is morphologically a broad, V-shaped eastward extension of the Argentine continen-
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Figure 2: Portion of seismic reflection line 16 on the northern sector of the Falkland Plateau Basin and the southern part of Argentine Basin, and its interpretation (relative position in Figure 1). The deep depocenter and the faults affecting the basement in the Falkland Plateau Basin and escarpment are clearly visible. The syn-sedimentary faults in sequences A and the numerous diffraction hyberbolae characterizing the acoustic facies of the basement in the Falkland Escarpment are visible. The main unconformities are clearly recognizable: r1 unconformity is the basement top of Precambrian age; r2 is an unconformity which extends from the upper Late Jurassic to the lower Early Cretaceous; the dot line in B sequence represent an upper Early Cretaceous discontinuity; the r4 unconformity develops at the Cretaceous-Paleogene transition. Vertical scale in TWT seconds. For the description of the different A to D sequences, see text. tal shelf (ACS). The FMP is bounded to the north by the steep slope of the Falkland Escarpment (FE), which corresponds to the Agulhas-Falkland Fracture Zone (AFFZ), to the south by the Falkland Trough (FT), and to the east by the Maurice Ewing Bank (MEB) (Figure 1). In the last decades, the region has been the object of numerous investigations for scientific purposes, focused on the nature of the FMP basement. Several authors (e.g. [2] and references therein) affirm that it consists of a thinning continental crust (12 to 16 km), although the hypothesis of a basement formed by subaerial ocean spreading is not excluded ([3] and refer-
ences therein).
2
Regional geological setting
It is note that the FMP evolution history is closely linked to the Gondwana breakup and the South Atlantic Ocean opening. Ben-Avraham et al. [4, 5] recognize a â&#x20AC;&#x153;preriftâ&#x20AC;? phase (Permian-Late Triassic), during which buoyancy forces, produced by large mantle plumes, around the present day Bouvet triple junction, played an important role in the activation of strike-slip fracture systems accompanying to south761
Marine Geology
Figure 3: Portion of the seismic reflection line 15A on the northern sector of the Falkland Plateau Basin and the southern part of Argentine Basin, and its interpretation (relative position in Figure 1). It is possible to note the faults affecting the basement and forming small basins in proximity of the Falkland Escarpment. The main unconformities are clearly recognizable: r1 unconformity is the basement top of Precambrian age; r2 is an unconformity which extends from the upper Late Jurassic to the lower Early Cretaceous; the dot line in B represent an upper Early Cretaceous discontinuity; r3 is a Late Cretaceous unconformity; the r4 unconformity develops at the Cretaceous-Paleogene transition. Vertical scale in TWT seconds. For the description of the different A to D sequences, see text. ern Gondwana break-up. Others authors describe an “early syn-rift” phase (Late Triassic-Early Jurassic) in which evidence of crustal block rotations and subsequent basal volcanism, generated by strike slip faults systems which also caused the following AFFZ development, are found in the South Africa Kaapvaal Craton. The subsequent “syn-rift” phase, which lasted up to the Late Cretaceous, is generally subdivided in four steps ([6] and references therein). In the first one, which covers the upper Early Jurassic, a splitting up of Southern Gondawana took place with a consequent opening, in the Middle Juras-
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sic, of several basins all localised crosswise over the proto-Atlantic ocean margins [7]. At the Jurassic/Cretaceous boundary the new southern American and Southern African margins, with several shelf basins controlled by strike-slip fault systems, start to be outlined. Contemporaneously the FMP (South America) and the Agulhas Plateau (South Africa) take shape through the AFFZ activation [6]. During the Early Cretaceous the South Atlantic Ocean starts its formation throughout the southern midAtlantic Ridge activation. [8], At the Late Cretaceous, the FMP and Agulhas Plateau definitively separated ([8] and references
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therein). According to the evolution history of the South Atlantic Ocean, in the Falkland Plateau Basin (FPB), the marine sedimentation starts at the Late Jurassic and several studies have proposed that the FPB, filled with more than 7000 m of sediments, presents a variable depth and deepens from north to south (e.g [9] and references therein ). Its depocenter is located near the Falkland Trough (FT) (Figure 1), an E-W depression generated from the active leftlateral strike-slip fault, which marks the South America-Scotia plates boundary and separates the FPB from the North Scotia Ridge ([10] and references therein).
3
Material and methods
The two unpublished multichannel seismic reflection profiles here analysed (15A and 16), collected during an ArgentineUruguay Reconnaissance Seismic Survey (1977) in the northern area of the FPB, sum a total of about 490 km, and are oriented NW-SE and NE-SW respectively (Figure 1). The seismic profiles were available only on paper record. Some published seismic reflection and refraction profiles [11, 12, 13], collected within or very near to the study area (Figure 1, black dotted lines), were also utilized to complete the seismostratigraphic analysis, which was integrated with published well data reports. The wells were drilled on the MEB in the 1977 and 1980, during DSDP Legs 36 and 71 (sites 327A, 329, 330A, 330 and sites 511, 512A, 512 respectively, Figure 1: blue squares). All boreholes recover the sedimentary units, but only hole 330 reaches the acoustic basement, visible in the published seismic profiles [12]. A seismicstratigraphic interpretation was carried out
on the profiles using standard procedure. The seismic sequences, constituted by a relatively conformable succession of seismic reflectors, are distinguished on the basis of top and base bounding unconformities and their correlative conformities, and are interpreted on the base of their internal reflector geometries ([14] and references therein).
4
Seismo-stratigraphic interpretation
Based on the seismic profiles analysis, the stratigraphic succession is composed of an acoustic basement (BS) overlain by four seismic sequences (A to D), bounded by relative unconformities (r1, r2, r3 and r4). The BS, all recognized seismic sequences and their boundaries, were correlated in the studied FPB sector, by means of interpolated published seismic profiles (Figure 1). Although there are some uncertainties in our analysis, our interpretations are underpinned by the high continuity and reflectivity of the main discontinuities correlated all across the studied FPB area. Basement The top of the acoustic basement is defined by the r1 unconformity, detected on the unpublished seismic profiles, which represents an erosional surface (Figures 2 and 3). Numerous strong diffraction hyperbolas, testifying a strong contrast of acoustic impedance, mark this unit in certain sectors, in particular along the FE, where it crops, fractured and dislocated by high-angle normal faults, that define grabens and half-grabens (Figures 2 and 3). South of FE the basement deepens and the FPB depocentre is filled by the overlying seismic sequence A (Figures 2 and 3). The acoustic facies of this unit is 763
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characterized by a chaotic configuration. Sequence A Above the r1 unconformity, the basal seismic sequence A shows subparallel to divergent internal configurations (Figure 2). The vertical thickness of sequence A varies greatly along the studied profiles; it is thickest near the faults that control the depocenters of the depressions. However, it disappears or becomes rather thin on the structural highs (Figures 2 and 3). Some reflectors within sequence A are interrupted and dislocated by direct faults (Figure 2). Because these faults involve neither the more superficial reflectors of the sequence nor the overlying units, they are believed to be sinsedimentary faults. The upper boundary of sequence A is the r2 unconformity, which cuts the reflectors lying below and shows a very high amplitude, dipping steeply southward as for the r1 unconformity (Figure 2). Sequence B Seismic sequence B lies in angular unconformity above the r1 and r2. This sequence is characterized from a parallel internal configuration (Figure 3). Also the sequence B is absent or extremely thin on the structural highs of the basement (Figures 2 and 3). At the south this sequence is characterised by a slight dipping and thickening (Figure 2). The r3 unconformity represents the upper boundary of sequence B (Figure 3), and generally cuts sequence B reflectors, preserving, as the r2, a southward gently dip (Figure 3). Sequence C The seismic sequence C onlaps r2 and r3 unconformities and often shows an hummocky internal configuration (Figure 3). It has a spread and moderate thickness, presenting a slight southward inclination (Figure 3), whereas it is practically absent near the FE, most probably due to strong active erosion (Figures 2 and 3). The top of sequence C corresponds to the r4 unconformity, which is a very 764
marked reflector easily traceable throughout the area, where it tends to be subhorizontal, contrary to the boundary of the other sequences observed (Figures 2 and 3). This unconformity truncates the underlying sequences B and C, sometimes leading to their complete elision, as can be seen near the southern margin of the FE (Figures 2 and 3). Sequence D The overlying seismic sequence, indicated as D, extends from the r4 unconformity to the ocean floor (Figures 2 and 3). This sequence is characterized by a parallel internal configuration (Figure 2). The morphological evidences show an erosion of sequence D, this erosion some time reaches the r4 unconformity and suggests an active ocean floor erosion (Figures 2 and 3).
5
Discussion
The acoustic basement detected in the seismic profiles was interpreted from the interpolated DSDP site 330 hole data, as constitute of continental nature (crystalline basement), with a Pre-cambrian age [12, 11], and with the north FMP margin (FE in Figure 1) interpreted as a continental-oceanic contact by Rabinowitz and LaBrecque [15] and by Lorenzo and Wessel [16]. However Ludwig [17] and Barker [3] affirmed that in the central FMP the basement is not simply characterized by continental crust, as an extension of the ACS, but a thicker oceanic crust, representing a zone of transition between a continental scarp and an aborted ocean basin. Also the gravimetric and magnetic data observed in this area [1, 18], coupled with the observed oceanic crust samples (gabbros and olivine basalts) dragged at sea-floor along the south FE [13], suggest that the nature of
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the acoustic basement, in a restricted area south of FE could be different from a continental one. On the other hand, going south, despite the poor resolution of seismic data, the seismic facies analysis, together with gravimetric data analysis, suggest again a possible continental nature of the basement [1, 18]. These findings could be the expression of crustal fragments of different nature that could be split and dragged from proto-Atlantic margins during the activity of AFFZ [19]. The age and nature of sediments within seismic sequences and their boundaries were calibrated with borehole data. The top of BS, r1, interpreted as an erosional surface unconformity, could corresponds to pre-Cambrian U1 unconformity of Lorenzo and M¨utter [13]. The overlying sequence A can be correlated to continental and shallow marine water deposits (black shales) found at Site 330, on the western flank of the MEB [12, 11], varying in age from Middle Jurassic to upper Late Jurassic (unit D1, [13]). Sequence A fills the major depressions in the area, and it is thought that this sequence was deposited during the first phases of crustal extension due to the initial Gondwana break-up, prior to the opening of the Southern Atlantic Ocean, which led to the development of an accommodation space in the basin ([6]). The growth faults observed within this sequences (Figure 2) show a maintenance of accomodation space during its deposition. The top of sequence A is marked by the r2 unconformity, which has been related with the U2 unconformity of Lorenzo & M¨utter [13], varying in age from upper Late Jurassic to lower Early Cretaceous. So the r2 unconformity corresponds to an important erosional surface developed at the same time of the regional thermal uplift and activa-
tion of the AFFZ, which accompanied the Gondwana break-up [8, 6]. Sequence B, which lies in angular unconformity above the basement and sequence A, has been related to the D2 sequence, varying in age from middle Early Cretaceous to lower Late Cetaceous [13]. This sequence is characterized by shallow (black shales) to deep marine deposits [12, 11]. Finally the increasing thickness of sequence B in the FPB suggest the fast sinking of the studied area [12, 11, 20]. The marked reflector within the sequence B (dot line in B sequence, Figure 2), could be related with the change from shallow (black shales) to deep marine sedimentation evidenced in D2 sequence occurring at the late Early Cretaceous [13]. This change is thought to be related to the activity of new bottom currents within the deeper oceanic basin, developed in relation to the initial opening of the South Atlantic Ocean [11, 6]. The top of sequence B is delimited by the r3 unconformity (Figure 3), correlated to the U3 unconformity, dated from upper Early Cretaceous to middle Late Cretaceous and interpreted as a strong erosional contact linked to a lower Late Cretaceous uplift ([13]). The uplift is thought, by the same authors, to be induced from the input of heat from an oceanic ridge passing along the AFFZ. Above the r3, the seismic sequence C can be correlated with the Late Cretaceous pelagic sediments of D3 sequence of Lorenzo and M¨utter [13]. The sequence C, ending at the upper Late Cretaceous, is bounded in the upper part by the regional r4 unconformity (Figure 3). This unconformity has been correlated with the U4 unconformity, attributed to Southern Atlantic Ocean bottom currents (i.e.Del Ben and Mallardi [9] and references therein). However our seismic profiles show evidence 765
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(Figure 3) that at south of FE, the r4 unconformity cuts a deformed sequence B and r3 unconformity, suggesting that, at the Cretaceous/Paleogene boundary, the area experienced uplift, with consequent partial erosion along the FE area. At the same time an ongoing subsidence is still affecting at south as testified by the major thickness of sequence C and by the pelagic nature of correlated sediments (unit D3 of [13]). Subsidence and relative uplift at north of this area, is in agreement with the contemporaneous active transpressive system bounding the South America and Scotia plates, which produced the underthrusting, and consequent southward tilting, of the southern FMP ([9] and references therein). Sequence D, overlying the r4 unconformity, has been correlated to drift depositional sequences (D4 and D5 of [13]) with a sedimentation characterized, for the entire Cenozoic period, by marine shelf and pelagic sediments, due to the most recent phase of the South Atlantic spreading. Within sequence D, an irregular development of marked reflectors was found (Figure 2). Such surfaces, correlated to the U5 unconformity of Lorenzo and M¨utter [13], are probably linked to the activity of strong circum-polar currents activated after the Drake Passage opening at the EoceneOligocene transition (e.g. [21] and references therein). Finally, it is possible to notice that the sequence D, deposited in a time interval longer than sequence C, shows respect to it a lower and constant thickness cross over the studied area, except for the morphological highs. This finding suggests that the area, during the Cenozoic, underwent a decreasing of the subsidence rate, and supports the known change in the tectonic regime of the FMP occurred at about Late Miocene, when the northward migration of the North Scotia 766
Ridge ceases and the seafloor spreading stops in the Central Scotia Sea ([22] and references therein).
6
Conclusions
The aim of this work was to reconstruct the seismo-stratigraphic of the depositional history of the central-northern FPB sector, in order to analyze and compare its geological evolution during the main phases of the Africa-South America separation and the subsequent South Atlantic opening. Four seismo-stratigraphic sequences and four main unconformities have been recognised, interpreted and correlated throughout the studied area. During the Permian-Late Jurassic pre-rift tectonics, the acoustic basement, correlated throughout the r1 unconformity, undergone to a phase of erosion, probably due to the thermal uplift. Subsequently, during the Middle Jurassic, time in which the proto-Atlantic initiated its opening with consequent developing of several depressions positioned crosswise over the margins, the FMP basins formed, characterized by continental deposits, as testified by DSDP well data on the MEB. At the Jurassic/Cretaceous boundary, the r2 unconformity highlights a tectonic behaviour of uplift. Whereas the thickness of sequence B, characterized by marine sediments, suggests the geological evolution of basin from the beginning of Early up to the Late Cretaceous, under the continuing Gondwana break-up and developing of AFFZ phases. During the Late Cretaceous, as known from literature, this area underwent a strong southward tilting interpreted as a lower Late Cretaceous thermal uplift, linked to the migration of the oceanic ridge
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along the AFFZ, evidenced, from well data, by probable hiatus occurring in this time interval. The marked erosional unconformity r3 is probably an expression of this tilt. At the end of Late Cretaceous the studied sector is still continuing its subsidence, as testified by the thickness of sequence C and by its correlation with pelagic sediments from well data. Also if contemporaneously the FE area at north of the studied sector is still experiencing uplift, as seen by the thinning and planning of sequence C thickness on this area, with a consequent partially erosion, testified by the r4 unconformity cutting the r3 unconformity and the deformed sequence B. Subsequently, during the Cenozoic, this area undergone to positional behaviour characterized by marine sedimentation de-
posited at a slower rate of subsidence, as testified by the lower thickness of sequence D with respect to sequence C. This decrease of the accommodation space is attributed to the known change in the tectonic regime of the southern FMP margin, which passes from an oblique subduction to a clearly transform regime. Finally it was found that the basement, in a restricted area south of the FE, is not clearly characterized by continental crust. Future studies will be aimed to underpinning such hypothesis and investigate the origin of the basement in this area, which may be linked either to a spread of basaltic volcanism occurred during the early phase of the Gondwana break-up, or to a process of split and dragging of different crust fragments during the early activity of the Agulhas-Falkland Fracture Zone.
References [1] E. Petruccione. Struttura geologica e crostale del settore nord del plateau delle Malvinas e sue correlazioni tettono-stratigrafiche con la piattaforma continentale argentina. Ph.D. Thesis, University of Napoles “Federico II” (Italy), http://www.fedoa.unina.it/1114/, 2006. [2] G.S. Kimbell and P.C. Richards. The three-dimensional lithospheric structure of the Falkland Plateau region based on gravity modelling. Journal of Geological Society, 165(4):795–806, 2008. [3] P.F. Barker. Evidence for a volcanic rifted margin and oceanic crustal structure for the Falkland Plateu Basin. Journal of the Geological Society, 156:889–900, 1999. [4] Z. Ben-Avraham, C.J.H. Hartnady, and J.A. Malan. Early extension between the Agulhas bank and the Falkland Plateau due to rotation of the Lafonia microplate. Earth and Planetary Science - Letters, 117:43–58, 1993. [5] Z. Ben-Avraham, C.J.H. Hartnady, and K.A. Kitchin. Structure and tectonics of the Agulhas-Falkland fracture zone. Tectonophysics, 282:83–89, 1997. [6] D. Macdonald, I. Gomez-Perez, J. Franzese, L. Spalletti, et al. Mesozoic break-up of SW Gondwana: implications for regional hydrocarbon potential of the southern South Atlantic. Marine and Petroleum Geology, 20:287–308, 2003. 767
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[7] K. Burke. Development of graben associated with the initial ruptures of the Atlantic Ocean. Tectonophysics, 36:93–112, 1976. [8] C. Reeves. The geophysical mapping of Mesozoic dyke swarms in southern Africa and their origin in the disruption of Gondwana. Journal of African Earth Sciences, 30(3):499–513, 2000. [9] A. Del Ben and A. Mallardi. Interpretation and chronostratigraphic mapping of multichannel seismic reflection profile I95167, Eastern Falkland Plateau (South Atlantic). Marine Geology, 209:347–361, 2004. [10] E. Lodolo, R. Menichetti, R. Bartole, Z. Ben-Avraham, A. Tassone, and H. Lippai. Magallanes-Fagnano continental trasform fault (Tierra del Fuego, southernmost South America). Tectonics, doi: 10129/2003TC0901500, 22:1076, 2003. [11] P.F. Barker, I.W.D. Dalziel, G. Dinkelman, D.H. Elliot, et al. The evolution of the southwestern Atlantic ocean basin: results of Leg 36. In Barker P.F., Dalziel I.W.D. et al. (Eds.), Initial Reports of the Deep Sea Drilling Project, 36, U.S. Government Printing Office, Washington D.C. 35:993–1014, 1977. [12] P.F. Barker, I.W.D. Dalziel, G. Dinkelman, D.H. Elliot, et al. Site 330. in Barker P.F., Dalziel I.W.D. et al. (Eds.), Initial Reports of the Deep Sea Drilling Project, 36, U.S. Government Printing Office, Washington D.C., 36:207–227, 1977. [13] J.M. Lorenzo and J.C. M¨utter. Seismic Stratigraphy and Tectonic Evolution of the Falkland/Falkland Plateau. Revista Brasileira de Geociˆencias, 18(2):191–200, 1988. [14] D. Emery and K. Myers. Sequence Stratigraphy. Blackwell Science, Oxford, 1996. [15] P.D. Rabinowitz and J. LaBrecque. The Mesozoic South Atlantic Ocean and Evolution of its Continental Margins. Journal of Geophysical Research, 84(B11):5973– 6002, 1979. [16] J.M. Lorenzo and P. Wessel. Flexure across a continent-ocean fracture zone: the northern Falkland/Falkland Plateau, South Atlantic. Geo-Marine Letters,, 17:110– 118, 1997. [17] W.J. Ludwig. Geologic Framework of the Falkland Plateau. In Ludwig W.J., Krasheninnikov V.A. (Eds.), Initial Reports of the Deep Sea Drilling Project, 71, U.S. Government Printing Office. 71:281–293, 1983. [18] E. Petruccione, M. Iorio, E. Lodolo, G. Nardi, and A. Tassone. Structural setting of the westernmost Agulhas-Malvinas Fracture Zone at the intersection with the Argentinean Continental Shelf. Actas del XVI Congreso Geol`ogico Argentino, La Plata, 2005.
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[19] G. Uenzelmann-Neben and K. Gohl. The Agulhas Ridge, South Atlantic: the particular structure of a fracture zone. Marine Geophysical Researches, 25:305–319, 2004. [20] W. Harris, W.V. Sliter, P.F. Barker, I.W.D. Dalziel, et al. Evolution of the southwestern atlantic ocean basin: results of Leg 36, Deep Sea Drilling Project. In Barker P.F., Dalziel I.W.D. et al. (Eds.), Initial Reports of the Deep Sea Drilling Project, 36, U.S. Government Printing Office, Washington D.C. 36:993–1013, 1977. [21] E. Lodolo, F. Donda, and A. Tassone. Western Scotia Sea margins: Improved constraints on the opening of the Drake Passage. J. Geophys. Res. doi: 10.1029/2006JB004361, 111(B06101), 2006. [22] A.P. Cunningham and P.F. Barker. Evidence for westward-flowing Weddell Sea Deep Water in the Falkland Through, western South Atlantic. Deep-Sea Research,, 43:643–654, 1996.
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Holocene Sedimentary and Gravitative Processes in Highstand Prodelta Deposits on the Cuma Outer Shelf (Eastern Tyrrhenian Sea, Italy): an Integrated Approach E. Petruccione1 , G. Aiello1 , G. Capretto2 , M.R. Senatore2 , E. Marsella1 , M. Iorio1 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, Department of Geological and Environmental Studies, University of Sannio, Benevento, Italy marina.iorio@iamc.cnr.it Abstract High resolution seismic profiles interpretation and sedimentological and petrophysical correlation of piston cores has been carried out to define sedimentary and gravitative processes in a prodelta area of the Cuma offshore. This unit is constituted in the upper part by mud deposits, which is affected by a downward displacement cropping out at the sea bottom, related to creep-type process. The integrated approach (sedimentologic, petrophysic and seismostratigraphic analysis) allowed to establish that the sedimentological body interested by the deformation preserves an internal geometry testifying a slow sliding of the sedimentary body without reworking. Moreover the emplacement of the upper unit, characterized by mud deposits, is interpreted as controlled by high sediment supply, high water contents and shallow gas pockets which could represent a principal triggering factor of the submarine instabilities.
1
Introduction
In recent years, there has been increasing interest in Quaternary shallow-marine progradational deposits that developed on the Mediterranean shelf or at the shelf edge in response to active processes as rapid global sea-level fluctuations and local shifting of deposition sites, ([1, 2] among others). Where progradational deposits can be linked to a fluyial-drainage system, depositional environment is inferred to be deltaic and such deposits have been referred to as shelf-margin deltas. Well-
studied examples of shelf-margin deltas lie in the Tyrrhenian and Adriatic sea [3, 4] among others). Prodelta sediments are represented typically by fine-grained deposits delivered by the river, reflecting episodes of minor and major river supply. The morphology and sedimentary facies of marine deltas are controlled to a large degree by the sediment input and the hydrodynamic regime of the receiving marine basin (e.g [5]). The prodelta area are also disturbed commonly by soft sediment failure and, in some deltas, prodelta sediments are dominated by products of mass movements de-
Marine Geology
Figure 1: The study area and bathymetric map of the Cuma offshore showing the dataset of the present research. Black lines and labels indicate the selected chirp seismic profiles (four out of twenty-five) and related fix. Black circles are gravity cores. The dotted line marks the sea bottom area (grey) interested by creep phenomenon. riving from the delta front (e.g Coleman, [6] and references therein). The principal reasons for sediment-induced deformation include (a) the relatively high sedimentation rate on the delta front which causes undercompaction and high pore fluid pressures, leading in turn to loss of shear strength within the deposits: (b) biodegradation of organic debris and associated free methane gas which weakens the sediment stability; (c) shocking of accumulated sediment by storm wave action; (d) sediment instability induced by earthquakes. Among the factors that may have significant impact on delta construction the frequency of recurrence of exceptional events
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as river floods, mudflows and explosive eruptions (pyroclastic falls, surges and flows) from coastal volcanoes is to point out. These events can all induce the supply of large volumes of loose sediment into the delta system and over vast areas of the continental shelf. During the last decades, exploration by means of high-resolution swath bathymetry and sub-bottom profiling of Campanian margin (comprising several deltas systems), has provided much insight on environmental and climatic evolution, tephrostratigraphic deposits in marine settling, and widespread occurrence of submarine instabilities (including creeping, slides, slumps
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Figure 2: Transf 0 profile (location in Figure 1) and its line drawing. TS1+R1: postGlacial-Holocene transgressive surface; MFS: Holocene Maximum Flooding Surface; FST and LST: Late Pleistocene seismic units; TST: post-Glacial-Holocene seismic unit; HST: Holocene seismic unit. The upper HST unit shows submarine instability processes (creep); diffuse shallow gas pockets (arrows); active tributary channel (ch 1) of the Cuma canyon; gravity core C1158 and GA35N chirp profile projections are shown. characterized by a sequence supplied by the deltaic system of the Volturno river and evidenced by related sedimentary features, showing an internal organization with minor downlap surfaces and erosional truncations, separating different phases of progradation ([15], [16] and reference therein). The last progradation started at about 100 ka and ended at about 18 ka during the Last Glacial Maximum [17]. A marked 2 Geological Setting erosional surface separates the forced regression from the overlying deposits that The upper Pleistocene-Holocene stratiwere formed from about 18 to 5 ka, durgraphic architecture of the continental shelf ing the sea-level rise and from about 5 ka offshore the Volturno area (Figure 1), is and mass flows) [7, 8, 9, 10, 11, 12]. So far a new integrated approach, based on high resolution multiproxy analysis, was applied to define sedimentary and gravitative processes of highstand deposits in a prodelta area (Volturno River, Cuma offshore).
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Figure 3: M198A profile (location in Figure 1) and its line drawing. Abbreviations in Figure 2 legend. GA35N, GA37N and Transf 0 chirp profiles projections are shown. to present during the sea level highstand ([18, 7], [19] and references therein). The Upper Holocene depositional unit consists of a wedge showing high thickness decreasing toward the shelf edge. Fast aggradation episodes are largely due to inland volcanoclastic and alluvial contributions. The volcanic events are associated with Phlegrean Fields activity. The slope is characterized by uniform profile and is incised by several canyons. In particular the Cuma canyon, between Ischia and Ventotene islands, SW-NE oriented, is a recent complex drainage system which includes the northern submerged part of the Ischia island and is characterized by three main heads and several minor tributary channels [20].
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3
Materials and methods
Four Chirp profiles (penetration of 2550 meters below sea bottom), among the twenty-five profiles carried out during the oceanographic cruise IAMC-GMS02-01, are presented and discussed (Figure 1). The studied cores C1158 and C1162, (Figure 1), were drilled on the outer shelf, at about 10 and 12 km from the coast, respectively. The cores penetrated some 6 m into Upper Holocene deposits from about 100 to 150 meters below sea level. Core recovery was about 50% in cores C1158 and 65% in core C1162, consisting of 2.54 m and 3.54 m of sediment, respectively. Compaction is estimated to be roughly the 15% of the whole length. Cores were split in two halves, photographed and described for lithostratigraphic analysis and used for continuous high-resolution
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Figure 4: GA35N profile (location in Figure 1) and its line drawing. Abbreviations in Figure 2 legend. M198A and Transf 0 chirp profiles and C1158 projections are shown. petrophysical and sedimentological analysis. The lithostratigraphic analysis of sediment cores was based on the macroscopic characters of the sediment according to the standard methodology (e.g. [21]) and on standard grain size analysis. Representative stratigraphic logs listing these characters and the presence of gradual, sharp and erosional contacts, were developed (Figure 6). The physical properties, Volume Magnetic Susceptibility (VMS) and Gamma Ray Density (GRD), were measured at 2 cm intervals using a Multi-Sensor Core Logger (MSCL) at the CNR-IAMC petrophysical laboratory. The petrophysical measurements were undertaken as soon as possible after collection of data, in order to minimise any drying or oxidation effects on the magnetic, geochemical and mineralogical properties. Values of petrophysical properties were plotted in logs, then visually and mathematically compared in order to detect similar curve trends, and correlate to the lithological properties according to their physical properties (Figure 6). A numerical method to compare the petrophys-
ical data and process them is provided by using an interactive program (SPLICER) on Unix platform [22].
4
Grain size analysis
The grain size analysis of samples collected at 5 cm intervals along the cores was carried out and provided the data for texturally and environmentally characterizing the sediment. Particles coarser than 63Âľm were analysed by sieving. The ÂĄ 63Âľm was analysed by sedimentation method, following the drawing times of Belloni [23]. The statistical parameters of Folk and Ward [24] were calculated from cumulative curves and the percentages of all fractions were calculated for each sample. The sediment was classified using the Shepard diagram [14], and environmentally interpreted using the C-M diagram of Passega [13]. The percentages of all fractions and the calculated statistical parameters, were plotted versus stratigraphic depths of the studied cores (Figure 6). 775
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Figure 5: GA37N profile (location in Figure 1) and its line drawing. Abbreviations in Figure 2 legend The location of an active channel (ch 1, Figure 2) is shown controlled by a relic channel identified in the stratigraphic record. C1162 and M198A and Transf 0 chirp profiles are shown.
5 5.1
Results Seismic-stratigraphic interpretation
1째 of slope, determining a progradation of the platform (Figure 3). On the Transf 0 Chirp profile, the upper FST unit is characterized by clinoform reflectors which lie in downlap on different underlying reflectors (Figure 2). The wedge shaped LST unit is always present at the shelf break (Figures 2and 3). Its reflectors show low amplitude and continuity. The upper bounding surface of both the FST and LST is represented by TS1+R1, characterized by high amplitude and continuity, cutting the underlying reflectors (Figures 2 and 5).
A seismic-stratigraphic interpretation was carried out for all acquired profiles using standard procedure. The recognised seismic units, constituted by a relatively conformable succession of reflectors, are interpreted on the base of their internal reflector geometries ([25] and references therein). The unit boundaries have been correlated in the studied area. Data from gravity cores were used for the interpretation of the up5.1.2 Unit TST per HST unit (Figure 6). The unit is characterized by parallel and continuous seismic reflectors with moder5.1.1 Units FST and LST ate amplitude (Figures 2 and 5). Several The FST unit is constituted by reflectors flooding surfaces (f in Figure 2) are diswith moderate amplitude and continuity, tinguishable showing higher amplitude and which are in offlap towards SW with about continuity. The TST lower reflectors lies 776
Marine research at CNR
in onlap on the lower limit TS1+R1, while the upper limit is constituted by MFS. The thickness of this unit, which is laterally quite constant, is of about 10 m.
50 m to about 120 m (grey in Figure 1) and show a variable thickness, from about 5 to less then 1 m, decreasing towards the shelf break. On the base of seismic profiles the undulation phenomenon presents lateral continuity in the related shelf sector 5.1.3 Unit HST and is more developed in the southern part The unit is characterized by well strati- where it extends up to the shelf break (Figfied seismic reflectors with moderate to ure 1). high amplitude and good continuity (Figures 2 to 5), occasionally affected by shal5.2 Sedimentologic analysis low gas pockets, which hide the acoustic signal (Figures 2 to 5). However it is According with the Shepard diagram (Figpossible to notice a low reflectivity of the ure 6 A-a; B-a) the sediment is classified acoustic signal in the proximity of an ac- as silt, clayey silt and silty clay. Both cores tive channel (Figures 2 and 5). The lower are mainly composed of olive fine to very unit limit is MFS on which the HST reflec- fine silt, being the median diameter mainly tors lie in downlap. The thickness unit de- between 3 Âľm and 8 Âľm and with the sand creases from coast towards the shelf break and gravel fractions absent or less than 1%; with value ranging from about 10 to 5 m in the sediment is moderately to poorly sorted the distal area while in the proximal area, and shows mainly coarse skewed and lepdue to the presence of gas pockets obscur- tokurtic grain size distributions (Figure 6, I ing the signal, a precise estimate is not pos- to X). The Passega C-M pattern is typical sible (Figure 3). of an environment with quiet water (Figure 6 A-b; B-b), where very fine sediment set5.1.4 Upper HST seafloor undulations tle without being sorted. The sediment comprises volcanic and bioLocally in the upper part of HST unit, it clastic components. The volcanic comis possible to recognize several undulations ponent is mainly constituted by black ash characterized by chaotic or well-layered and/or millimetric grey pumice clasts. The conformable reflectors with low lateral bioclastic component is represented mainly continuity and variable thickness. More- by molluscs shell fragments, rarely parts of over it is possible to notice the conformable Echinoids can be recognized. nature of both the undulated layered facies Intercalated within this sediment there and the underlying (pre-existent) reflectors are several levels of finer (clayey silt (Figures 2 to 5). or silty clay) and coarser (medium silt) Artefacts and/or sea-surface wave effects poorly sorted sediment, showing platikurcausing the observed undulations are ex- tic curves. Some alignments of shell cluded since they would have affected fragments of Echinoids can be recognized sub-bottom profiles along all their length, within the coarser levels, testifying the ocwhereas parts of the profiles appear undu- currence of higher energy events. lation free. These data suggest a continuous and The deposits involved in the undulations undisturbed distal sedimentation, typical extend in water depths ranging from about of an aggrading/prograding continental 777
Marine Geology
shelf, also supplied by the contribution of volcanic deposits and storm events. The macroscopic sedimentological observations together with the grain size analysis allow us to subdivide both core logs in three sedimentological units (r1 and r limits Figure 6, C and D).
high-resolution (centimetric scale) correlation between the two cores. Thanks to this procedure, we were able to define 10 homologous points occurring at the same interval depths (Âą 5cm) on both cores in all petrophysical logs (Figure 6, 1 to 10).
5.4 5.3
Petrophysic analysis
To allow comparison among petrophysical (VMS and GRD) data and to correlate them to the 5 cm step grain size analysis, the data were first averaged out with a 5 cm window and plotted against depths (Figure 6, XI to XIV). The VMS and GRD data, are characterized from top to bottom core, by wavering pattern around similar mean values (58,6 and 61,8 SIx10-5 for VMS data and 2,01 and 1,99 gr/cc for GRD data in cores C1158 and C1162 respectively) with several distinct peaks mostly of high amplitude (Figure 6, XI to XIV). From a visual check of the logs it was noted a noticeable correspondence between C1158 and C1162. To compare petrophysical features between cores, the obtained data were processed using the SPLICER software [22], which statistically compares the different wiggle matching of the data set from one core with those from the others, until an optimum correlation, based on cross correlation functions, is established. So it was found that, in the depth intervals from 0.28 to 2.54 and 0.52 to 3.41 mbsf for C1158 and C1162 respectively (dark area in Figure 6, XI to XIV) the data show a very good correlation among the VMS values and good for the GRD values. Afterwards, we analyzed the data more in detail trying to pinpoint single prominent and highly informative peaks and troughs in the mathematically correlated portion of datasets. This process is aimed to obtain a 778
Sedimentologic and petrophysic correlation
The obtained petrophysical correlation for core C1158 and C1162 were than checked within the remaining grain size logs and it was found a good correspondence with the relative increase of the clay fraction percentage in the sedimentological records (Figure 6,I and VI) at the same stratigraphic depths of seven out of ten relative high (VMS) and low (GRD) correlated peaks (Figure 6, XI to XIV). So far it is possible to suggest that a higher relative concentrations of magnetic minerals (VMS higher values) are expected in clay contents, as well as low value peaks, for GRD properties, are to be expected, considering the high porosity and the high water content of the clay fraction (Weber et al., 1997). Finally around the sedimentological boundary units r and r1, peaks 4 and 10 correspond to a relative increase of the silt fraction percentage, in line with the marked lithological changes (Figure 6. I and VI, C and D). So far coherent correlation of sedimentological and petrophysical parameters from the south-western and north-eastern sectors of the studied area shows a high resolution pattern that is comparable within the first 2,5 m of sediment, and we may conclude that the correlated upper part of seismographic unit HST, which has been sampled by cores C1158 and C1162, is constituted by the same sedimentological layers (Sedimentologic Unit 1 and 2 in Figure 6 C and D).
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Figure 6: Logs of C1158 and C1162 respectively show: the variation versus depth of gravel, sand, silt and clay fractions concentration (I and VI); the variation versus depth of the Folk and Ward (1957) statistical parameters (II to V and VII to X). Legend for columns II to V and VII to X: Mz: Graphic Mean; ﾏナ: Inclusive Graphic Standard Deviation; Ski: Inclusive Graphic Skewness; Kg: Graphic Kurtosis. A) and B) Classification diagram of Shepard, 1954 (a) and C-M diagram of Passega, [13] (b) for C1158 and C1162 respectively. Legend for Shepard [14] diagram: S: sand; C-S: clayey sand; S-C: sandy clay; C: clay; Si-C: silty clay; C-Si: clayey silt; Si: silt; S-Si: sandy silt; Si-S silty sand. C) and D) GRD raw data (b) and macroscopically sedimentological logs (a) of C1158 and C1162 projected against the related seismic profiles showing the correlated sedimentation and density changes (r and r1: principal changes in I and VI and C and D). High resolution petrophysical properties correlations of C1158 and C1162 (XI to XIV): VMS (XI and XII) and GRD smoothed data (XII and XIV). The dark area in XI to XIV indicates correlation by means of cross correlation functions. See text for details. The numbers 1 to 10 indicate synchronous events of petrophysical (XI to XIV) and sedimentological (I and VI) variations occurring in the measured cores. Please note that the upper part of core C1158 is about 30 cm shorter than core C1162, probably due to the lost of sediment occurred during coring.
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6 6.1
Discussion Late Pleistoce-Holocene Stratigraphic architecture of continental shelf
Interpretation of the selected Chirp profiles allowed reconstruction of relationship between depositional elements of the Late Pleistocene and Holocene sedimentary sequence in the Cuma offshore.The seismostratigraphic analysis shows the occurrence of clinoforms in the FST and LST units (Figures 2 and 3), suggesting that, on the base of regional stratigraphic architecture of the continental shelf (i.e. [1, 16, 7] among others), they could represent prograding deposits emplaced during the Late Pleistocene sea regression when subaerial exposure associated with strong erosion took place, which originated a significant stratigraphic gap over the shelf, as suggested by the lack of the offlap breaks pertaining to the FST and LST units (Figures 2 and 3). The TS1+R1 erosional reflector (Figures 2 and 5), corresponding to the major regional surface topping the Late Pleistocene prograding units [26], can be followed readily all across the study area representing the end of the last regression at about 18 ka (Last Glacial Maximum ) when the sea reached -120 m with respect to the present sea level (i.e. [7, 19]). The TS1+R1 surface, associated with the time-transgressive landward shift of the fair-weather wavebase during the rapid sea-level rise that accompanied the Post Glacial-Holocene deglaciation, is coincident with both the ravinement surface (Figure 3) and the continental erosional surface never detected in the studied area. The LST, TST and HST units constitute the incomplete Late Pleistocene â&#x20AC;&#x201C; Holocene 780
sequence. Above TS1+R1, the seismic profiles show onlapping deposits (TST unit) representing the transgrassive system track in which the relatively continuous, parallel reflectors bounded by the more distinctive f surfaces representing coarse upward successions related to the episodic rising of the sea level during its rising up to the maximum flooding (5-6 ka B.P.). The expression of the maximum flooding sea event is represented by the MFS surface, which was recognised on seismic profiles by means of downlap geometry of the reflectors of the highstand systems tract HST unit.
6.2
Gravitative processes in low energy high stand deposits
In the upper part of the HST (Figures 2 to 4) the interpretation of the observed sub-seafloor undulations with internal layered seismic facies, is not straightforward. Two options about their genesis have to be considered in the light of similar studies (e.g. [27, 28, 29] among others). A plausible interpretation would relate the observed undulations to slow sediment deformation processes such as sediment creep folds, as suggested for similar seafloor features [30, 31, 4] produced by slow plastic deformation under a continuous load [32]. An alternate option is considering these undulations as sediment waves produced solely by depositional processes, for example by river-sourced storming events or hyperpycnal flows, similarly to the interpretations made in other areas (e.g. [33] and references therein) and bearing in mind that Campanian rivers are likely to develop hyperpycnal discharges [9, 34, 35]. However, unlike most undulations produced by turbidity currents and hyperpyc-
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nal flows, which are frequently associated to erosional features [36, 29] among others) the observed undulations are unrelated to erosional features. Moreover the deformed upper HST unit often shows internal slightly countersloping reflectors and mainly develops in correspondence to a higher gas pocket phenomena. Furthermore, in the deformed unit, both the undulated layered facies and the underlying (pre-existent) reflectors are in conformable position, indicating that the undulations were not generated by the roughness of the upper surface of the underlying reflectors [37]. All these differences do not allow an interpretation of the observed structures as sediment waves produced solely by depositional processes, even though this hypothesis, lacking multibeam data, cannot be definitively excluded. However a different evidence in favour of a very slow gravitative process may be furnished by the seismic interpretation, coupled with piston cores data (C1158, and C1162; Figures 2, 4 and6), which highlight that this part of the unit includes the same deformed mud layers rich in water content, (Figures 3 and 6A and B) and that such layers are contemporaneously characterized by undisturbed sedimentation and preservation of the internal geometry, and so demonstrating a lack of reworking. So far the emplacement of the undulation in the mud sediment, which is controlled both by high siliciclastic supply in correspondence to the Volturno river mouth and by the occurrence of shallow gas pockets (probably due to high contents of organic matter in the sediments), is likely due to a softsediment deformation
(creeping) with the occurrence of shallow gas pockets representing a triggering factor of the submarine instabilities.
7
Conclusions
High resolution seismic profiles interpretation evidence in the HST Holocene unit of Cuma offshore, a recent creep-type downward displacement, which crops out at the sea bottom. Seismic interpretation suggests also that gravity mass instability may be postulated for a large part of the continental shelf. In the sedimentological record of the upper sliding unit, ten homologous points in the correlated petrophysical logs have been recognized and at the same stratigraphic depths of the ten correlated petrophysical maximum values, a correspondence has been observed with a relative increase of the percentage clay fraction. Such sedimentological and petrophysical correspondence, the provenience of both correlated cores from the same deformed deposits (as evidenced by the seismic profiles) and the undisturbed sedimentation (as demonstrated by the Passega diagrams) constitute evidences that the sedimentological body interested by the deformation preserves an internal geometry testifying a slow sliding of the sedimentary body without reworking. The emplacement of the upper unit characterized by mud deposits is interpreted as controlled by high sediment supply, high water contents and shallow gas pockets which could represent a principal triggering factor of the submarine instabilities.
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References [1] F. Trincardi and M.E. Field. Geometry, lateral variation, and preservation of downlapping regressive shelf deposits: eastern Tyrrhenian Sea margin, Italy. Journal Sedim. Petrol., 61(5):775–790, 1991. [2] F.L. Chiocci, G. Ercilla, and J. Torres. Stratal architecture of Western Mediterranean margins as the result of the stacking of Quaternary lowstand deposits below glacioeustatic fluctuation base-leve. Sedimentary Geology, 112:195–217, 1997. [3] F.L. Chiocci and W.R. Normark. Effect of sea-level variation on upper-slope depositional processes offshore of Tiber Delta, Tyrrhenian Sea, Italy. 104:109–122, 1992. [4] F. Trincardi, A. Cattaneo, A. Correggiari, and D. Ridente. Evidence of soft sediment deformation, fluid escape, sediment failure and regional weak layers within the late Quaternary mud deposits of the Adriatic Sea. Mar.Geol.,, 213,:91–119, 2004. [5] D.J.P. Swift, S. Phillips, and J.A. Thorne. Sedimentation on continental margins: IV. lithofacies and depositional systems. In: Shelf Sand and Sandstone Bodies: Geometries, Facies and Sequence Stratigraphy. 14:89–152, 1991. [6] J.M. Coleman. Dynamic changes and processes in the Mississippi River Delta. Geological Society American Bulletin, 100:999–1015, 1988. [7] G. Buccheri, G. Capretto, V. Di Donato, P. Esposito, et al. A high resolution record of the last deglaciation in the southern Tyrrhenian Sea: environmental and climatic evolution. Mar.Geol.,, 186,:447–470., 2002. [8] M. Iorio, L. Sagnotti, A. Angelino, F. Budillon, et al. High resolution petrophysical and palaeomagnetic study of Late Holocene Shelf Sediments, Salerno Gulf, Tyrrhenian Sea. The Holocene, 14(3):426–425, 2004. [9] F. Budillon, C. Violante, A. Conforti, E. Esposito, et al. Event beds in the recent prodelta stratigraphic record of the small flood-prone Bonea Stream (Amalfi Coast, Southern Italy). Marine Geology, (222-223):419–441, 2005. [10] T. De Pippo, C. Donadio, M. Pennetta, C. Petrosino, et al. Coastal hazard assessment and mapping in Northern Campania, Italy. Geomorphology, 97:451–466, 2008. [11] M. Iorio, J. Liddicoat, F. Budillon, P. Tiano, et al. Palaeomagnetic Secular Variation Time Costrain on Late Neogene Geological Events in slope sediment from the Eastern Tyrrhenian Sea in ”External Controls on Deep-Water Depositional Systems”. 92:233–243, 2009. [12] C. Violante. Geohazard in Rocky Coastal Areas. page 322, 2009.
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[13] R. Passega. Grain size representation by CM patterns as a geological tool. Journ.Sedim. Petrol.,, 34(4):830–847, 1964. [14] F.P. Shepard. Nomenclature based on sand-silt-clay ratios. Journ. Sed. Petrol.,, 24(3):151–158, 1954. [15] R. Bartole, D. Savelli, M. Tramontana, and F.C. Wezel. Structural and sedimentary features in the Tyrrhenian margin off Campania, Southern Italy. Mar. Geol.,, 55,:163–180, 1984. [16] G. Aiello, E. Marsella, and M. Sacchi. Quaternary structural evolution of the Terracina and Gaeta basins (Eastern Tyrrhenian margin, Italy). Rend.Fis.Acc. Lincei, 11:41–58, 2000. [17] J. Chappell and N.J. Shackleton. Oxygen isotopes and sea level. Nature, 324:137– 140, 1986. [18] D. Hunt and M.E. Tucker. Stranded parasequence and the forced regressive wedge systems track: deposition during base-level fall. Sedim.Geol.,, (181):1–9, 1992. [19] G. Ciampo. Reconstruction of Late Pleistocene – Holocene palaeobathymetries from Ostracoda on the Tyrrhenian continental shelf. Geobios, 36,:1–11, 2003. [20] M. Pennetta, A. Valente, D. Abate, G. Budillon, et al. Influenza della morfologia costiera sulla circolazione e sedimentazione sulla piattaforma continentale campano-laziale tra Gaeta e Cuma (Italia meridionale). Boll.Soc.Geol.Ital., 117:281–295, 1998. [21] M. Tucker. Techniques in sedimentology. pages 1–394, 1988. [22] T. Hagelberg, M.J. Shackleton, N. Pisias, and Shipboard Scient. Party. Development of composite depth sections for Sites 844 trough 854. page 79–85, 1992. [23] S. Belloni. Una tabella universale per eseguire granulometrie col metodo della sedimentazione unica o col metodo del densimetro di Casagrande modificato. Geol.Tecn.,, 16,:1281–1289, 1969. [24] R.L. Folk and W.C. Ward. Brazos river bar: a study in the significance of grain size parameters. 27,:3–26, 1957. [25] D. Emery and K. Myers. Sequence Stratigraphy. 1996. [26] M.G. Coppa, L. Ferraro, M. Pennetta, B. Russo, et al. Sedimentology and micropaleontology of the core G39 – C27 (Gaeta bay, central Tyrrhenian Sea, Italy). Il Quaternario, 9(2):687–696, 1996. [27] A. Cattaneo, A. Correggiati, T. Marsset, T. Yannick, et al. Seafloor undulation pattern on the Adriatic shelf and comparison to deep-water sediments waves. Marine Geology, 324:137–140, 2004. 783
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[28] T. Marsset, B. Marsset, Y. Thomas, A. Cattaneo, et al. Analysis of Holocene sedimentary features on the Adriatic shelf from 3D very high resolution seismic data (Triad survey). Marine Geology, 213:73–89, 2004. [29] L.M. Fernandez-Salas, F.J. Lobo, J.L. Sanz, V. D`ıaz del R`ıo, et al. Morphometric analysis and genetic implications of pro-deltaic sea-floor undulations in the northern Alboran Sea margin, western Mediterranean Basin. Marine Geology, 243:31– 56, 2007. [30] A. Correggiari, F. Trincardi, L. Langone, and M. Roveri. Styles of failure in late Holocene highstand prodelta wedges on the Adriatic shelf. Journal Sedimentary Research, 71:218–236, 2001. [31] V. Lykousis, D. Sakellariou, and G. Rousakis. Prodelta slope stability and associated coastal hazards in tectonically active margins: Gulf of Corinth (NE Mediterranean). In: Submarine mass movements and their consequences. pages 433–440, 2003. [32] P.R. Hill, K.M. Moran, and S.M. Blasco. Creep deformation of slope sediments in the Canadian Beaufort Sea. Geo Marine Letters, 2:163–170, 1982. [33] B. Urgeles, C. DeMol, M. Liquete, M. Canals, et al. Sediments undulations on the Llobregat prodelta: Signs of early slope instability or sedimentary bedforms? Journal Geophysical Research, 112(B05102):12 pp, 2007. [34] F. Budillon, D. Vicinanza, V. Ferrante, and M. Iorio. Sediment transport and deposition during extreme sea storm events at the Salerno Bay (Tyrrhenian Sea) comparision of field data with numerical model results. Natural Hazards Earth System Sciences, 6:839–852, 2006. [35] M. Sacchi, F. Molisso, C. Violante, E. Esposito, et al. Sea seismic examples off the Amalfi cliffed coasts, eastern Tyrrhenian. Insights into flood-dominated fan-deltas: very high-resolution. In: Geohazard in Rocky Coastal Areas. 322:33–71, 2009. [36] G. Verdicchio and F. Trincardi. Short-distance variability in slope bed-forms along the southwestern Adriatic margin (central Mediterranean). Marine Geology, 234:271–292, 2006. [37] C. Berndt, A. Cattaneo, S. Magdalena, F. Trincardi, and D. Masson. Sedimentary structures offshore Ortona, Adriatic Sea – Deformation or sediment waves? Mar.Geol, 234,:261 – 270, 2006.
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Submarine Landslides: Mediterranean Sea
Case Studies in the
M. Rovere, G. Dalla Valle, F. Foglini, F. Gamberi, M. Marani, F. Trincardi Institute of Marine Sciences, CNR, Bologna, Italy marzia.rovere@bo.ismar.cnr.it Abstract Submarine landslides are the result of downslope transfer of sediment toward the deepest part of the ocean basins. They play an important role in the shaping and evolution of continental margins and have wide implications in defining the geological hazard along the coastal areas. This contribution is the showcase of recent studies carried out by ISMAR on submarine slides in the Mediterranean. In the SW Adriatic a slide has been emplaced during the MIS2-glacial interval and more recent coalescing slide scars document a diffuse instability. In the Sicily Channel recurrent sediment failures exploit specific stratigraphic surfaces as glide planes. In the Gioia Basin a mass-transport complex is the result of multiple events since the uplift of Calabria in the Pleistocene. The Cefal`u Basin hosts a buried frontally-emergent slide. In the Paola Basin shallow-seated slides are triggered by mud diapirism. Offshore SE Sardinia mass-transport deposits are associated with canyon systems. Offshore Calabria mass-wasting is concentrated along linear fault scarps. The database spans a range of geodynamic contexts and shows that mass-transport deposits often result in the stacking of repeated failures, can occur at any moment in a margin history, can be related to seismicity and regional uplift, can influence deep sea ecosystems, may act as potential hydrocarbon reservoirs and not least, that the study of their rheology and kinematics is fundamental in assessing the tsunamigenic potential.
1
Introduction
Gross-scale continental margin shaping is represented by mass-transport complexes that are the result of repeated sediment failures and the stacking of translated deposits. Both the failure events and the resulting mass-transport deposits involve highly different mechanisms of emplacement that are then reflected in the geometry and internal structure of the displaced masses. Terms such as slides, slumps or flows are based on the transport mechanism. For example, slides transfer on a planar glide surface, while slumps move on an upwards concave surface with a rotational movement
that causes the mass to be highly disrupted. Debris-flows are defined as a mass of larger clasts supported by a finer matrix associated with distal compressional structures. One type of landslide may evolve into another or trigger other slope movements. It is widely accepted that seismicity, slope oversteepening, climate change, gas hydrates dissociation, rapid accumulation of sediments, wave loading, tectonic activity, isostatic rebound are all causal factors capable of contributing to the instability of a submarine area. The occurrence of submarine landslides has been traditionally associated with low stand conditions based on
Marine Geology
Figure 1: bathymetry of the Central Mediterranean Sea and location of the submarine landslides (red areas) discussed in the text. sequence stratigraphy or statistical observations of enhanced instability during initiation of glaciation and the transition from glacial to interglacial periods. Notwithstanding, several studies have now demonstrated that submarine landslides can occur at any time in a continental margin history. New data generation, such as high resolution bathymetry data and 3D seismic techniques, have revealed the internal structure and detailed geometry of landslide deposits. All these elements have proved to be fundamental in characterizing the rheological behavior of submarine landslides. This paper deals with several examples in the Central Mediterranean (Figure 1) displaying how, through the integrated analysis of high resolution multi-
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beam bathymetry and seismic data, accurate biostratigraphy correlations and sequence stratigraphy modeling, is possible to reconstruct the evolution, to assign a time and to establish a kinematic model of submarine landslides.
2
SW Adriatic Margin
The Southwestern Adriatic margin (Figure 1) is characterized by Pleistocene progradational sequences that show to be repeatedly affected by mass-wasting events [1]. In this section, we concentrate on the Gondola and Vieste Slides, but evidence of slope instability is widespread north and south of the Dauno Seamount and within
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Figure 2: a) bathymetry of the SW Adriatic margin, b) Gondola slide blocks in multibeam and c) TOBI imagery, d) CHIRP profile across the Gondola slide blocks (showing preserved stratification). the Bari Canyon (Figure 2a). The Gondola Slide is the most prominent masstransport deposit within the area, with a 10km-wide headscarp and a 60-km-runout; it was emplaced during the MIS-2 glacial interval (30-25 ka [2]), hence during a glacial low stand of sea level. The headwall area is occupied by large blocks that have been swept by strong bottom currents since their emplacement, preventing the slide burial [2, 3], Figure 2b-d). The blocks are up to 3-km-long and 500 m wide, with a 40-m-thick preserved stratification that can both represent a poten-
tial hydrocarbons reservoir and positively influence the settlement of a variety of benthic hard-bottom organisms, including frame-building corals [4]. The Vieste Slide is a smaller feature located northward, it has a 30-m-high headscarp and 7-km-wide scar; sequence stratigraphy modeling and accurate biostratigraphic correlations indicate that the slide mobilizes contourite deposits emplaced after the LGM [5], during the sea level rise. To the north of the Gondola Slide, the outer shelf is characterized by N-S oriented scars and scours [3] entrenched in middle Pleistocene sed-
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iment that have been proved to consist of shelf-edge muddy contourite deposits rimmed landward by the present-day moat ([6] Figure 2a). These current-dominated deposits are intercalated with small scale mass-wasting episodes. In a similar way, downslope the Vieste slide, recent thinskinned mass failure affects the down slope flank of numerous sediment drifts ([5] Figure 2a). This evidence suggests that also muddy-bottom current deposits are susceptible of failure during or soon after their deposition [6] and can contribute to the margin instability.
3
Gela Basin
The continental slope of the Gela Basin (Figure 1) hosts diverse types of masstransport deposits dated back to LGM (Father slide, 24-18 ka [7]). After the emplacement of the first large mass-transport deposit, the occurrence of repeated mudflows and high sedimentation rates rapidly increased pore pressure and fluid expulsion, giving rise to a generalized margin instability ultimately resulting in the two exposed Twin Slides [7]. Although their apparent similarity, the Twin Slides show quite a different behavior: the northern Twin Slide represents a debris avalanche deposit overridden by a slump deposit with pressure ridges in the toe region. The southern Twin Slide consists of a debris avalanche deposit with broken blocks and erosional features at the base [8]. This marked differentiation reflects the distinct stratigraphic units affected by failure: the Northern Slide emanated from younger and less consolidated progradational units of the last glacial low stand, while Southern Slide involved older deposits, possibly more consolidated, from the previous 788
two glacial cycles [7]. The Twin Slides represent the youngest failures in the Gela Basin. Their head scarps clearly down cut late Holocene units, suggesting a very young age for these failures [7]. Anyhow, the Twin Slides occurred at least in two stages: the first stage mobilized the recently deposited drape of the post-glacial depositional sequence (at Sapropel S1, 8.5 cal. kyr BP); the second stage involved older more lithified materials and resulted in the formation of pressure ridges and slide blocks (Figure 3d, [8]). The Twin Slides exploited key stratigraphic surfaces as glide planes. The northern Twin Slide failed on the marine onlap of the progradational lowstand wedge and moved basinward atop a preexisting mass transport deposit [8]. The southern Twin Slide emploied the shear plane previously followed by the Father Slide [7]. Along with this major evidence, also some smaller-scale mass wasting processes are affecting the margin. The buried slide scarp of Father Slide strongly influence bottom currents, but it is along the basinward flank of the southern moat that thin-skinned mass wasting is occurring (Figure 3a [6]).
4
Gioia Basin
In the Gioia Basin (Figure 1), the large Villafranca mass-transport complex (VMTC) developed as a result of repeated failure events starting from Late Pleistocene. Notwithstanding the basin witnessed largescale mass-wasting processes since the Early Pleistocene [9]. The oldest event is the emplacement of the Nicotera basinwide slump (â&#x2C6;ź 700 km2 ) due to the beginning of regional tectonic uplift of Calabria (700-900 ka). After that, sheet turbidites from the southern Calabria and northeast-
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Figure 3: a) bathymetry of the Gela basin slope, b) Twin Slides in multibeam and c) TOBI imagery draped on the multibeam bathymetry, d) CHIRP-sonar profile across the Twin Slides deposits. ern Sicily margins buried the slump until the development of a channel-levee complex build up by the Niceto and Villafranca channels (Figure 4a). In Mid-Late Pleistocene, when a local differential uplift took place in NE Sicily, the 46 km3 Villafranca frontally-confined slide was discharged. The lithologic discontinuity at the base of the levee wedge is the main basal shear surface above which sediment translation occurred (Figure 4b, stage 2). Subsequent failures where located in the upper slope and resulted in the burial and remolding of the headwall area of the Villafranca
slide through a slab slide (Figure 4b, stage 3 and Figure 4c), a frontally-emergent debris-flow (Figure 4b, stage 4 and Figure 4c) and subsidiary processes, such as detached and downslope translated blocks (Figure 4c). In the distal part of the basin a buried chaotic body of 100 km2 overlies the Villafranca slide and may be the product of hyperpycnal flows from a channel system located further upslope (Figure 4a). Moreover, to the west of the VMTC, the channel-levee complex is becoming the site of incipient instability in the form of embryonic headwalls and proto-slump scars
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Figure 4: a) map of the VMTC deposits and the Nicotera basin-wide slump within the Gioia basin, b) 4 stages of VMTC evolution, c) 3D view from the headwall area of VMTC. (Figure 4a). The detailed chronology of mass-wasting achieved through the combination of high resolution bathymetry and seismic data[10] proves that the NE Sicily margin is highly susceptible to repeated and large-scale mass wasting.
5
Cefalu` Basin
The Cefal`u basin (Figure 1) is subdivided by the Solunto High into two smaller basins: the Palermo and Capo dâ&#x20AC;&#x2122;Orlando basins (Figure 5b). Active tectonics affect the area and NW-SE extensional faults dislocate the coastal basins and the nearby off790
shore area generating sedimentary troughs vs uplifted blocks. The Capo dâ&#x20AC;&#x2122;Orlando basin appear to be lowered in respect to the adjacent Palermo basin and it hosts the 200 km2 Orlando frontally emergent slide. The Orlando Slide is characterized by very well defined lateral margins with up to 50 m scarp heights. The dx margin (Figure 5d) is characterized by en-echelon geometry, while the sx (Figure 5c) margin coincides with an extensional fault that may be connected to mapped land structures (Figure 5b). The slide upper surface is covered by smaller-scale dip-slip collapses from the high steep lateral scarps and by rotated blocks in the proximal part, that are imaged
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Figure 5: a) sparker profile across the Orlando slide, b) bathymetry of the Capo dâ&#x20AC;&#x2122;Orlando basin, c) and d) lateral margins of the Orlando slide as imaged by side-scan sonar data, e) frontal ramp in seafloor reflectivity. in side-scan sonar data (Figure 5d). Whereas the confined portion (10-kmwide) of the landslide sits in the slope and is mainly exposed at the seafloor or is barely covered, the unconfined portion lies in the basin plain and is completely buried. The frontal ramp of the landslide spreads toward the NE forming two tongues overriding the paleo-seafloor (Figure 5e). The headwall region is located quite close to the coastline (15 km) at a significant water depth of 1,000 m. The continuous
shape of the headwall scarp with only minor salients is the evidence that the main phase of movement has taken place during a single phase and that retrogression is only a minor process in the slide evolution. This aspect coupled with the vicinity of the headwall area to the coastline, the narrow shelf-bathymetric jump and the high seismicity of the area represent a treat for the tsunami hazard along the lowered part of the northern Sicily coast and the Aeolian Islands [11].
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Marine Geology
Figure 6: Bathymetry of the Paola Ridge, D1-D3 are the mud diapirs.
6
Paola Basin
In the northern Calabrian margin (Figure 1), the Paola Ridge consists of circular or elongated ridges cored by a transparent seismic facies that are explained as sediment remobilization structures consisting of mud diapirs and mud volcanoes, associated with extensional faults dipping into sediment sealed by Messinian evaporites (Figure 6 [11]). Pockmarks fields and gas charged sediments are the evidence of degassing from inactive diapirs. Furthermore, two mud volcanoes are shown 792
by high backscatter mud flows (RMV and MMV in Figure 6). A combination of over steepened slope (diapiric topography) and gas-charged sediments can be envisaged for the activation of the shallow-seated landslides (3-4 km of lateral extent) located downslope from the diapirs and the intermediate slope pockmarks. Landward, a basin-wide mudflow (600 km2 ) is also documented (Figure 1) and it dates back to 1314 cal. kyr BP. It may be related to a relative sea level rise increase [12].
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Figure 7: bathymetry of the NE Sardinia margin. The headwall scars of the thin-skinned slides are indicated.
7
NE Sardinia margin and Capo Vaticano Ridge
In the NE Sardinia margin (Figure 1) there are examples of slope failures associated with the southern branch of the Tavolara canyon system, the Molara and Olbia canyons (Figure 7). The upper part of the canyon flanks are severely interested by thin-skinned slides testifying that also steeply highly erosional, hyperpycnal flows conduits can contribute into margin instability [13]. The Capo Vaticano peninsula (southern
Calabria) is a structural high located along the Tyrrhenian side of the Calabrian Arc (Figure 8) that was generated by longitudinal normal faults active since the Late Pliocene. Starting from the Middle Pleistocene, the rifting processes of the southern peri-Thyrrhenian basins were coupled with a strong regional uplift. The offshore prolongation of Capo Vaticano high (Capo Vaticano ridge) is now affected by mass wasting scars in its southern flank, that are probably caused by the still ongoing uplift of the area.
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Figure 8: bathymetry of the Capo Vaticano ridge. The headwall scars along the southern flank of Capo Vaticano ridge are indicated.
8
Concluding Remarks
The outcomes of these studies lead to some few important conclusions. 1) An accurate dating of the slides can be achieved through biostatigraphy and sequence stratigraphy correlations. Dating of a slide is not only important for addressing a geological problem, but is fundamental in establishing a recurrence timescale for an underwater territory. 2) Key stratigraphic layers, along which sediment translation is favored, can be recognized through seismic data and direct 794
sampling. This knowledge is essential to estimate the recurrence of landslides and future instability-related events that can affect a submarine area. 3) A detailed characterization of the internal make-up, basal shear surface and lithological composition of a landslide can reconstruct its rheology i.e. how a landslide behaves while moving downslope. A more advanced knowledge of landslides deformation will contribute to more accurate tsunami modeling that are still at their infancy. 4) Pre-conditioning factors are important
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and most of the times several factors may concur to cause a submarine mass movement. Notwithstanding, we demonstrated how several submarine landslides occur in areas of strong regional uplift, a phenomenon that can not be mitigated by human interventions. 5) Sedimentary processes, such as contourite deposits and mud diapirism, play an important role in triggering slope instability. A detailed knowledge of sedimentary
processes along continental margins is fundamental in assessing potentially vulnerable areas. 6) Submarine landslides may become fauna habitats and influence deep-sea ecosystems. Moreover can act as seals for gas and oil. Exploration of landslides can have wide implications in underwater exploitation planning for economic development and environmental quality strategies.
References [1] F. Trincardi, A. Cattaneo, A. Correggiari, and D. Ridente. Evidence of soft sediment deformation, fluid escape, sediment failure and regional weak layers within the late Quaternary mud deposits of the Adriatic Sea. Marine Geology, 213:91–120, 2004. [2] D. Minisini, F. Trincardi, and A. Asioli. Evidence of slope instability in the Southwestern Adriatic Margin. Natural Hazards and Earth System Sciences, 6(1):1–20, 2006. [3] G. Verdicchio and F. Trincardi. Short-distance variability in slope bed-forms along the southwestern Adriatic margin (central Mediterranean). Marine Geology, 234(14):271–292, 2006. [4] F. Trincardi, M. Taviani, A. Freiwald, F. Foglini, G. Verdicchio, and D. Minisini. A positive interaction between slump-induced seafloor topography and deep-water coral growth (SW Adriatic sea). 33rd International Geological Congress Book of Abstracts, 33:1351793, 2008. [5] G. Verdicchio and F. Trincardi. Sequence stratigraphy of Late Quaternary slope deposits in South Adriatic. Geo Acta, 1:97–116, 2008. [6] G. Verdicchio and F. Trincardi. Mediterranean shelf-edge muddy contourites: Example from Gela and South Adriatic basins. Geo Marine Letters, 28(3):137–151, 2008. [7] D. Minisini, F. Trincardi, A. Asioli, M. Canu, and F. Foglini. Morphologic variability of exposed mass-transport deposits on the eastern slope of Gela Basin (Sicily Channel). Basin Research, 19:217–40, 2007. [8] D. Minisini and F. Trincardi. Frequent failure of the continental slope: The Gela Basin (Sicily Channel). Journal of Geophysical Research, 114:F03014, 2009.
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[9] F. Gamberi and M. Marani. Hinterland geology and continental margin growth: the case of the Gioia Basin (southeastern Tyrrhenian Sea). Tectonics of the Western Mediterranean and North Africa, 262:349–363, 2006. [10] F. Gamberi, M. Rovere, and M. Marani. Mass-transport complex evolution in a tectonically active margin (Gioia Basin, Southeastern Tyrrhenian Sea). Marine Geology, 279(1-4):98–110, 2011. [11] F. Gamberi, M. Marani, and M. Rovere. Architecture and Causes of Submarine Landslides in the Southeastern Tyrrhenian Sea. 27th IAS Book of Abstracts, 27:177, 2009. [12] F. Trincardi, A. Cattaneo, A. Correggiari, S. Mongardi, A. Breda, and A. Asioli. Submarine slides during sea level rise: Two examples from the eastern Tyrrhenian margin. pages 469–478, 2003. [13] G. Dalla Valle, F. Gamberi, and M. Marani. Submarine Mass Wasting Processes along the Continental Slope of the Eastern Sardinian Margin. 27th IAS Book of Abstracts, 27:125, 2009.
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Environmental Changes and Human Impact in the Southern Tyrrhenian Sea During the Last 520 Years: Evidence from a High-Resolution Foraminiferal Record M. Vallefuoco1 , F. Lirer1 , L. Ferraro1 , M. Sprovieri2 , L. Capotondi3 , L. Bellucci3 , S. Albertazzi3 , S. Giuliani3 , A. Angelino1 , M. Iorio1 , M. Iavarone1 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, Institute for Coastal Marine Environment, CNR, Capo Granitola (TP), Italy 3, Institute of Marine Sciences, CNR, Bologna, Italy mattia.vallefuoco@iamc.cnr.it Abstract Climate has been acknowledged as one of the main driving factors behind human development and in turn human activity may cause strong changes on marine ecosystems. Coastal areas represent very useful marine sedimentary archives to investigate the potential effects of human impact on shelf marine ecosystems. The high-resolution investigation presented here of a southern Tyrrhenian marine record from the continental shelf of the Salerno Gulf provides an unprecedented opportunity to study the main natural and human impact that have occurred in this area during the last five centuries. Evidence of an important turnover between carnivorous and herbivorous opportunistic planktonic foraminfera species is recorded after the Maunder event. In addition, after 1940 AD, a strong increase in the distribution pattern of the Globigerina bulloides species combined with a sudden increase in the abundance pattern of the benthic foraminifer Bulimina aculeata, and evident enhancement of benthic and planktonic foraminifera (as number of specimens per gram of dry sediment) can be directly associated to the effects of the construction of the Sele River dam (1934 AD).
1
Introduction
At present, the study of historical records, aimed at a better understanding of the Earthâ&#x20AC;&#x2122;s climatic system and a more accurate prediction of its future evolution, represents one of the most important priorities of the scientific community, in such a way that it has become necessary to combine climatically sensitive proxy records with
data from historical chronicles. The Mediterranean Sea, due to its high sedimentation rates, offers a nice opportunity to investigate global and regional climate changes in its marine sedimentary archives. An increasing number of multiproxy studies (calcareous foraminifera, pollens, dinoflagellates, stable isotopes) document intense climatic oscillations during the last millennium: the Little Ice Age (LIA),
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the Medieval Warm Period (MWP), etc. These events were previously recognized in Mediterranean continental and marine archives by Mann et al. [1] and Esper et al. [2]. Particularly, the MWP (∼800 to ∼1300 AD) was a time of warm climate in Europe with temperatures allegedly comparable with the present ones [3]. Contrarily, the LIA (∼1450 to ∼1900 AD) appears to be characterized by an intense widespread cooling (on the order of 0.5 – 1.0°C) associated to an evident lowering of about 100 m of the equilibrium line altitude of mountain glaciers around the world (e.g. [4]). In the present contribution, combining planktonic and benthic foraminifera distribution patterns with historical chronicles, we present a preliminary high-resolution and well-dated dataset from a super-expanded shallow water marine sedimentary record of the eastern Tyrrhenian sea. This study aims to characterize the possible impact of the main climatic variations along with local human growth on the continental shelf marine environment for the last 520 years.
2
Geological setting and Material and Methods
The Salerno Bay is a peri-Tyrrhenian basin located along the Eastern Tyrrhenian sea [5]. It represents the offshore counterpart of the Sele coastal plain (Figure 1) which formed in response to a large scale Late Neogene-to-Pleistocene lithospheric extension that accompanied the eastward accretion and anticlockwise rotation of the Apenninic fold and thrust belt during the roll-back of the subducting Adria plate [6, 7, 8]. The physiography of the Salerno Bay reveals a significant control by the tec798
tonic setting of the coastal zone and a high sediment supply. The studied gravity core C90-1m (40°35.76’N; 14°42.48’E) has been collected with a SW104 corer. This system allows the recovery of undisturbed (devoid of sediment deformation) and very well preserved marine records, together with the water-sediment interface. Core C901m was retrieved close to the shelf break of the northern Salerno Bay, at a water depth of 103.4 m (Figure 1). The sedimentary succession consists of 106 cm thick hemi-pelagic marls (Figure 1). The high-resolution age model of the C901m core is based on 210 Pb and 137 Cs radiometric datings with an accurate definition of the sediment chronology for the upper 40 cm b.s.f.. The lower part of the core was dated by extrapolation from the overlying sediments. Analysis of planktonic and benthic foraminifera was performed on 104 and 43 samples, respectively. Sampling spacing was, for planktonic foraminifera, 1 cm from the top of the core (seafloor) down to the base, while for benthic ones the sampling spacing was 2 cm. All taxa are quantified as percentages of the total number of planktonic and benthic foraminifera. The concentration is reported as number of specimens per gram of dry sediment. Regarding the adopted taxonomy, some planktonic species or morphotypes are lumped together according to the scheme reported in Jorissen et al. and Capotondi et al. [9, 10]. The generic attribution of benthic taxa was made following the Loeblich and Tappan classification [11], species were mainly determined following Cimerman and Langer [12] and Sgarrella and Moncharmont-Zei [13].
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Figure 1: Location map and lithology of the studied core (C90-1m)
3
Age Model
The uppermost 40 cm b.s.f. (below sea floor) of core C90-1m were dated by alpha spectrometry measurements of 210 Po (assumed to be in secular equilibrium with its daughter, 210 Pb) at ISMAR – CNR, Bologna, following the procedures proposed by Frignani and Langone [14]. In addition, a number of levels were prepared for 214 Pb counting via gamma spectrometry using a germanium detector [15] in order to provide estimates of supported 210 Pb (i.e. in situ produced 210 Pb from the decay of 226 Ra) and to check the assumption of its constant activity. In order to support 210 Pb dating, the activity of 137 Cs was measured via gamma spectrometry using coaxial intrinsic germanium
detectors [15]. The 210 Pb activity–depth profile in core C90-1m shows an exponential decline with depth (Figure 2), suggesting a constant sediment accumulation over the last century. Consequently, the sediment accumulation rate was calculated for the first 40 cm b.s.f. by applying a Constant Flux–Constant Sedimentation model [16] to the activity–depth profile of excess 210 Pb (i.e. the difference between total and supported 210 Pb). A mean sediment accumulation rate of 0.20 cm·yr−1 was obtained, thus defining an age of 1802 AD at 40.5 cm b.s.f. (Figure 2). Measured 137 Cs activities are low when compared to those obtained in Northern Adriatic sediments [17], but show a clear trend detectable down to 15 cm (Figure 2). Assuming that the value at 11.5 and the 799
Marine Geology
Figure 2: Age model of core C90-1m. From left to right: magnetic susceptibility and colour reflectance plotted vs time. Age-depth profile with indication of mean sedimentation rate (calculated with CF-CS model); 210 Pb and 137 Cs activity-depth profiles for the first 40 cm b.s.f. of the core C90-1m peak at 7.5 cm b.s.f. can be associated to 1954 AD (first appearance of 137 Cs from nuclear explosions) and 1963 AD (maximum 137 Cs fallout), respectively, and that 2006.5 AD represents the topmost layer of the core, the resulting mean sedimentation rate is 0.18 cm·yr−1 (Figure 2). These values are in good agreement with those obtained from 210 Pb activity–depth profile. The preservation of a clear curve trend of 137 Cs activity suggests that the sedimentation rate has been mostly constant for the last fifty years at the core site. Finally, assuming a constant sedimentation rate of 0.20 cm·yr−1 for the entire core section and applying a linear interpolation back to 100 cm b.s.f., an age of 1480 AD for the lowermost studied layer of the core has been obtained (Figure 2). This age model suggest that the studied marine record covers the last 520 years. 800
4
Planktonic and benthic foraminiferal intervals
Evident changes in quantitative distribution patterns of different planktonic foraminifera species allowed several authors [27, 28, 29, 30, 31, 10], to define a number of eco-biozones used to subdivide the Mediterranean Holocene-late Pleistocene stratigraphic record. The ecobiozone boundaries are identified by events of temporary appearance or disappearance and/or evident abundance peaks of selected species. In this paper, we tentatively propose a subdivision in three eco-intervals (Figure 3) of the last 520 years in the continentalshelf marine record of Salerno Gulf (southern Tyrrhenian Sea), based on a highresolution age model.
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In terms of planktonic foraminifera, the first eco-interval A, from the base of the core (1480 AD) to 1720 AD, is mainly characterised by the absence (or very low abundance) of Globigerinoides quadrilobatus and the high abundance in the Globigerinoides ruber distribution pattern associated with the Globorotalia truncatulinoides l.c. wich shows two distinct peaks during at the Sporer and Maunder events (Figure 3). These events represent phases of least solar activity in the solar cycle of the sun (solar minimum), during which , sunspot and solar flare activity diminishes [21, 22, 18, 19, 23, 24, 20, 25, 26]. Upwards, the eco-interval B, from 1720 AD to 1920 AD, is dominated by Turborotalita quinqueloba and Globigerinita glutinata, associated with a strong decrease in abundance of G. ruber and Globigerina bulloides (Figure 3). Finally, the ecointerval C, from 1920 AD to 2006 AD, is characterised by the progressive decrease in abundance of G. ruber and by a further increase of T. quinqueloba, G. bulloides and of G. quadrilobatus (Figure 3). The distribution pattern of the benthic foraminifera assemblage shows an evident three-partition (eco-intervals) of the sedimentary record: eco-interval A, at the base of section (1480 AD), is characterised by the common occurrence of Hyalinea baltica, Uvigerina mediterranea and Valvulineria bradyana associated to low percentages of Bulimia marginata and Cassidulina carinata and to a relative increase of Bolivina alata in the uppermost part of the interval (Figure 3). The second eco-interval B, spanning from 1720 AD to 1920 AD, is characterized by the dominance of B. alata and V. bradyana, associated to the reoccurrence of H. baltica, after its temporary disappearance during the Maunder event (Figure
3). By contrast, U. mediterranea shows an antithetic distribution pattern respect to B. alata and V. bradyana. Melonis padanum shows the third strong peak at 1740 AD in the uppermost part of the Maunder event (lower part of interval B) (Figure 3). Finally, the eco-interval C characterizes the uppermost part of the core, from 1920 AD to the top (2006 AD) and is dominated by Bulimina aculeata and by the reoccurrence of U. mediterranea, which increases after its minimum at the end of interval B during Damon event (solar minimum) (Figure 3). Contrarily, a strong decrease is recorded in the abundance of B. alata and V. bradyana (Figure 3).
5
Discussion
The fact that the boundaries of the three identified eco-intervals for planktonic and benthic foraminifera are almost coincident (Figure 3) suggests a similar response to climate forcing. In addition, abundance changes in planktonic and benthic fauna seem to reflect the main climate changes recorded in the last 520 years by the solar variability proxies â&#x2C6;&#x2020;14C and Total Solar Irradiance (TSI) [18, 19]. In particular, the lower part of the studied record, which falls within the warm interval predating the onset of the well-known cold Maunder event, is characterised by the concurrence of Globorotalia inflata and G. truncatulinoides suggesting a deep high phytoplankton productivity. This condition is supported by the ecological features of these two taxa considered as high dietary dependence on phytoplankton. In fact, G. inflata and G. truncatulinoides are considered indicative of a deep mixed layer during winter [32] for the Mediterranean area, but G. truncatulinoides always lives 801
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at greater depth than G. inflata. Moreover, it lives in phosphate-rich waters and is a good recorder of thermocline nutrient levels [33]. In addition, the increase in planktonic and benthic foraminifera (as number of specimens per gram of dry sediment), during the upper part of the cold Sporer event, confirms a similar response to climate forcing of sea bottom water and column water environments. Upwards, the turnover between carnivorous (G. ruber) and herbivorous (T. quinqueloba and G. glutinata) opportunistic planktonic foraminifera at 1715 AD, represents the main climatic change occurring within the planktonic foraminiferal assemblage. This turnover occurs after the Maunder event, during which historical chronicles report the occurrence of the coldest winter in Europe for the last half millennium (at 1708-09 AD) ([24] and references therein). This planktonic foraminiferal change may be associated with a possible change of food availability in the water column and a possible and progressive shallowing of the phytoplankton productivity successively to the Maunder event as indicated by the increase of G. bulloides in the last century. Finally, after 1920 AD the planktonic foraminiferal assemblage, characterised by the progressive decrease in abundance of G. ruber and by a further increase of T. quinqueloba, G. bulloides and of G. quadrilobatus, suggests a very high surface water phytoplankton productivity. On the other hand a sudden increase in benthic foraminifer B. aculeata is recorded. This configuration, following the industrial development, could be associated to the building of the dam on the Sele river (one of the main rivers flowing in the Salerno Gulf) in 1934 AD, that caused a ceasing of bottom solid particle transport in favour 802
of an increased fine and nutrient-enriched fraction reaching the sea water system. Furthermore, G. quadrilobatus increased only during warm intervals in the last 520 years (Figure 3). In particular, the two strong peaks, centred at about 1760 AD and 1860 AD, correspond well to the warm intervals between the Maunder and Dalton (solar minimum) and between the Dalton and Damon events, respectively (Figure 3), and the progressive increase in abundance after 1920 AD marks the onset of modern warm conditions. This trend suggests, for the southern Tyrrhenian Sea, a strong relationship between oligotrophic conditions and warm periods as reported by Piva et al. [34] in the Adriatic sea and by Poore et al. [35] in the Caribbean area.
6
Conclusion
The detailed age model, based on 210 Pb and 137 Cs radiometric datings, obtained for studied marine core located on the continental shelf of the Salerno Gulf (Tyrrhenian Sea), allowed us to improve the reconstruction of the main environmental and climatic changes occurring during the last 520 years. In term of the benthic foraminiferal ecosystem, the sea bottom continental marine shelf environment for the last 520 years is relatively enriched in species that tend to dominate under more eutrophic conditions. The turnover in the planktonic foraminifera (carnivorous vs herbivorous-opportunistic species) at â&#x2C6;ź1715 AD may be associated to the Maunder event (XVII century) that, as reported by historical chronicles, coincides with a strong decrease in economic and human growth due to the strong flood and pestilence events in the study area. Afterwards, the strong increase of herbivorous-
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opportunistic species (T. quinqueloba and G. glutinata), documenting the progressive shallowing of the phytoplankton productivity in the water column, may be associated to the development of agriculture or increased pasturing and/or extensive landuse of coastal studied area. In the XX century, the uppermost climatic event experienced by the Mediterranean area is documented by the onset of modern warm conditions. The planktonic foraminiferal assemblage, particularly from 1934 AD, shows very low values of the warm water taxon G. ruber abundance pattern associated to an increase in the abundance of G. bulloides and G. quadrilobatus. Moreover, af-
ter 1940 AD the progressive increase in the G. bulloides distribution pattern combined with the sudden increase in benthic foraminifer Bulimina aculeata and coupled to the further prominent increase of benthic and planktonic foraminifera (as number of specimens per gram of dry sediment),was recorded. At the same time, as reported by historical chronicles, a dam on Sele River (Salerno Gulf) was built, which possibly changed the amount of grain-size river transport and the nutrient budget of the sea water. Consequently, the modern warm condition along with the anthropic modifications may be suggest a change in the marine environmental coastal-ecosystem of the last century.
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Figure 3: Distribution of selected species of planktonic and benthic foraminifera of the studied core C90-1m plotted vs time, with the position of the recognised eco- intervals. All taxa are quantified as percentages of the total number of planktonic and benthic foraminifera. The concentration is reported as number of specimens per gram of dry sediment. Correlation with tree-ring D14C [18], Total Solar Irradiance (TSI) [19] and sun spot data [20] is also presented. The blue band correspond to the time of the solar activity minima, for the last 550 years. In particular, the Sporer Minimum (AD 14201540), the Maunder Minimum (AD 1645-1715), the Dalton Minimum (AD 1790-1830) and the Damon Minimum (AD 1880-1900) that occurred in the Little Ice Age (LIA, AD 1450-1900) [21, 22, 18, 19, 23, 24, 20, 25, 26].
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Marine Climate Archives and Geochemical Proxies: a Review and Future Investigations on the Mediterranean Sea P. Montagna1,2 , M. Taviani1 , S. Silenzi3 , M. McCulloch4 , C. Mazzoli5 , S. Goldstein1 , R. Rodolfo-Metalpa6 1, Institute of Marine Sciences, CNR, Bologna, Italy 2, Lamont-Doherty Earth Observatory, Columbia University, Palisades, New York, U.S.A. 3, Institute for Environmental Protection and Research, Roma, Italy 4, The UWA Ocean Institute and School of Earth and Environment, Perth, Western Australia 5, Department of Geosciences, University of Padova, Italy 6, International Atomic Energy Agency, Environment Laboratories, Monaco, Principality Of Monaco montagna@ldeo.columbia.edu Abstract Paleoclimate research based on the investigation of geochemical proxies in marine climate archives has been growing considerably during the last two decades, due to the development of more precise analytical systems. Thermal ionization and inductively coupled plasma mass spectrometers, equipped with multi-collectors, enable to obtain precise and accurate isotopic data, an essential requirement for reliable reconstructions of the physical and chemical marine parameters. In the last 5 years, part of the paleoclimate investigation focused on the Mediterranean has been carried out using these cutting-edge analytical techniques through international collaborations among Italian, Australian and American scientists. Specimens of shallowand deep-water corals collected in the Mediterranean Sea, in the Atlantic and Pacific Oceans, have been analysed with laser ablation and solution ICP-MS and with a thermal ionization mass spectrometer. The correlation between Li, Mg, P, normalized to Ca, and the Nd and B isotopic composition of the coral skeletons with the most important marine parameters has enabled to develop and validate new geochemical proxies. The derived calibration equations can now be applied to well-dated fossil corals with the aim to reconstruct the climate variations in the past. This article reviews some of the principal results achieved in the last 5 years by the authors and present some future directions on the application of geochemistry to coral investigation.
1
Introduction
tools for Late Quaternary paleoclimate reconstructions of important marine parameConsiderable effort has been made in re- ters, such as temperature, salinity, nutrient cent years to refine and develop geochem- content and pH evolution [1, 2, 3, 4, 5, 6]. ical proxies in coral skeleton as reliable
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Figure 1: Shallow and deep-water corals collected in the Mediterranean Sea. A. Distal view of the calyx of a live Desmophyllum dianthus showing the orange tentacles and the septa. B. Fossil D. dianthus colonized by sub-recent and live specimens of the same species. C. Distal-lateral view of a 37kyrs old D. dianthus. D. Transmitted light image of the main septum (S1) of a D. dianthus. White and dark microbands form the density banding characteristic of this species. E. 48kyrs old Lophelia pertusa colony retrieved in the Strait of Sicily. F. Cladocora caespitosa subspherical colony from Miramare in the Northern Adriatic Sea. Colony is characterized by distinct corallites, each having its own wall, growing in a continuous rectilinear way. Due to their carbonate mineralogy, corals can be precisely dated by means of AMS 14 C and mass spectrometric U/Th dating [7, 8], and they systematically incorporate trace elements and stable isotopes, thus being capable to provide multi-century, subannual resolution records. Since the interpretation of coral geochemistry is complicated by the overprinting of physiological processes of skeletal formation, more emphasis has been placed to better understand the mechanisms of elemental uptake and isotopic fractionation, in order to retrieve reliable climatic and oceanographic
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data from coral geochemistry (see [9], for a review). Recent investigations on the trace elements and stable isotopic composition of the aragonitic exoskeleton of Mediterranean shallow and deep-water corals (Figure 1) have proved that different corals species thriving in the Mediterranean Sea at different depths in the water column can serve as climate archives [10, 11, 5, 6, 12]. Spatially-resolved geochemical measurements of lithium, boron, magnesium, phosphorus, calcium, strontium and uranium have been carried out on the skeleton of the
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shallow-water coral Cladocora caespitosa and the cold-water coral species Lophelia pertusa, Madrepora oculata, Desmophyllum dianthus, and Caryophyllia smithii using a high-resolution laser ablation ICPMS. This analytical technique allows the analysis of minor and trace elements at fine-scale resolution, in order to geochemically characterize the different microstructures comprising the coral skeleton. The centres of calcification are considered to be the first component formed in the skeleton and are composed of tiny crystals surrounded by fibrous aragonitic bundles, which represent the bulk of the skeleton. Both coral microstructures show characteristic elemental and isotopic patterns, which might alter the interpretation of the environmental signals if not properly taken into account. On the basis of previous results at fine-scale resolution, we decided to focus on specific geochemical ratios (i.e. Li/Mg) that seem to overcome the bias related to the vital effect, giving a unique opportunity to retrieve precise information on the evolution of seawater temperature in the past. Very recently, there has been a growing interest in investigating non-traditional isotopic systems, such as neodymium and boron, in coral skeletons, with the aim to further our knowledge on the water mass dynamics and the pH evolution. Given the analytical difficulties in obtaining reliable Nd and B isotopic data in coral skeletons only few studies have been conducted so far [13, 14, 13] The necessary precision for paleoclimate investigations required the use of more precise and accurate analytical machines, such as multi-collectors spectrometers, either with inductively coupled plasma or thermal ionization sources. Here we briefly review some of the main results that we have obtained during the last five years in studying the geochemi-
cal composition of Mediterranean shallow and deep-water corals with an emphasis on temperature and nutrient sensitive elements. In addition, we present some of the new frontiers in coral geochemistry and the application of these new promising proxies in retrieving paleoclimate data for the Mediterranean Sea.
2
The Mediterranean sea: a natural laboratory for studying climate change
Due to its peculiar geographic position, situated at the boundary between the subtropical and mid-latitudes zones and its water mass dynamics, the Mediterranean Sea can be considered a perfect natural laboratory to understand the effects of climate changes on marine environment from the photic to the bathyal zone. The Mediterranean Sea is a semi-enclosed basin with an â&#x20AC;&#x153;anti-estuarineâ&#x20AC;? circulation: the Atlantic water enters through the Strait of Gibraltar as surface water and it progressively becomes denser while flowing eastward through the Strait of Sicily towards the Levantine basin. In the northern Levantine basin the Modified Atlantic Water sinks to form the more saline Levantine Intermediate Water (LIW), which represents the major constituent of the Mediterranean Intermediate Water and part of the Mediterranean Outflow into the Atlantic Ocean. The LIW flows westward from the Eastern basin between 150 and 600 meter water depth via the relatively shallow Strait of Sicily (see [15], for a review). During wintertime the effect of atmospheric cooling and evaporation produce cold dense waters that cascade off the continental shelf in specific locations of the Mediterranean Sea 811
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such as in the Gulf of Lions, in the Southern Adriatic Sea and in several Aegean shelves. This process forms the deep water system of the Mediterranean Sea, a relevant but still poorly known component of the entire basin. Even though this is a very simplicistic view of the Mediterranean circulation it gives an idea of the complexity of the Mediterranean dynamics, characterized by a general circulation typical of the open ocean. All these basin scale features were and are impacted by climate change and the investigation of some of the most important physical and chemical parameters, such as sea water temperature, nutrient contents and the state of the ventilation, in the past can help to further elucidate the climate evolution and the processes operating in the basin. The Mediterranean Sea behaves as a “miniature ocean” [16], with contrasting water masses flowing from the Eastern to the Western basins and viceversa, therefore representing an ideal casestudy for building up our knowledge not only on the internal Mediterranean dynamics but also on the climate changes on a global scale. The Mediterranean seems to respond faster than the open oceans to climate variations and this gives the unique opportunity to study the effects of the climate change at a shorter time-scale using the coral geochemistry.
3
Previous geochemical studies of Mediterranean corals for paleoclimate reconstructions
Previous studies on the geochemical composition of shallow and deep-water coral skeletons from the Mediterranean Sea have 812
been mainly focused on the development of geochemical proxies for the seawater temperature and the nutrient content reconstruction of the water column [10, 5, 6]. Silenzi et al. [10] analysed for the first time the geochemical composition of the zooxanthellae-bearing coral Cladocora caepitosa, a species living between 5 and 40 meters water depth and one of the most important marine biocostructural organisms in the Mediterranean Sea [17]. The Sr/Ca composition of the high-density bands, which represent part of the exoskeleton forming during the wintertime, was found to be significantly correlated to the winter sea surface temperature (SST), providing the first calibration equation. Montagna et al. [6] refined the Sr/Ca equation using a laser ablation ICP-MS on a living coral collected in the Northern Adriatic Sea. In addition, a suite of elements, such as Li, B, Mg, P, Ca and U were analysed at fortnightly resolution. A new set of calibration equations were derived between B/Ca, Mg/Ca and U/Ca against SST. The results proved that this Mediterranean species can serve as a reliable recorder of the SST, although the “vital effect” can influence the elemental uptake (see below). At present, several samples from different locations in the Mediterranean Sea are being geochemically studied using bulk solution and laser ablation ICP-MS with the aim to produce precise calibrations to be applied to the fossil records. An extensive review of these results can be found in Montagna et al. [12].
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Figure 2: Sampling locations of the live Lophelia pertusa specimens used for Li/Mg measurements.
4
New frontiers in coral geochemistry
In the following pages we will discuss some of the new developments in coral geochemistry and the application of these new promising proxies in reconstructing the seawater temperature, the nutrient content, the water dynamics and the pH evolution of the Mediterranean Sea from the photic zone to the bathyal environment.
4.1
P/Ca ratios in the skeleton of deep-water corals: a proxy for seawater nutrient chemistry
The knowledge of the past seawater concentration of phosphorus, one of the ma-
jor bio-limiting nutrient, is important to estimate the contribution of the â&#x20AC;&#x153;biological pumpâ&#x20AC;? to the levels of atmospheric CO2 . Therefore, evaluating past changes in the export of biological production to deep waters is a major issue. Unfortunately, most of the geochemical proxies used to quantify the seawater paleo-phosphorus concentration suffer to varying degrees of additional environmental factors, which potentially complicate the paleoceanographic reconstructions. We developed for the first time a direct method to reconstruct seawater paleo-phosphorus concentration by analysing the P/Ca encoded in the skeletal aragonite of the deep-sea coral species Desmophyllum dianthus [5]. The measurements were performed using a laser ablation ICP-MS and the beam was focused on the outer face of the main septum (Fig813
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Figure 3: Map of the Mediterranean Sea with the sites where fossil deep-water corals are already available for geochemical investigations. Most of these samples have been U/Th dated with ages spanning the last ca. 500kyrs. ure 1). A calibration equation was derived through the correlation between the coral P/Ca ratios of several specimens collected worldwide from a range of geographic locations and the corresponding dissolved inorganic phosphorus (DIP) concentration. The application of this equation to precisely dated fossil corals allowed us to characterize the trophic state of the intermediate water in the Western Mediterranean Sea at the end of the Younger Dryas period. Further fine-scale investigations enabled to improve the previous P/Ca vs. DIP equation, taking into account the presence of micro-size domains likely enriched in micro-apatite [18]. This proxy still represents the only direct method available so far to obtain information on the phosphate content of individual water masses in the past. Given the right fossil samples, the coral P/Ca proxy allows to quantify the fluxes of nutrients to intermediate and deep-water environments. Moreover, it can be used to reconstruct the
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past ocean productivity, thereby furthering our understanding of the biological functions of the Mediterranean Sea in regulating atmospheric CO2 .
4.2
Li/Mg ratios in scleractinian corals: a new precise paleothermometer
Sensitive elements used in corals suffer to varying degrees of limitations due to physiological processes [19, 20, 21, 22, 23]. The coral physiology is actively controlling the chemical composition of different portions of the skeleton, and the temperature reconstructions are often complicated and biased by the overprint of this â&#x20AC;&#x153;vital effectâ&#x20AC;?, likely related to the differential skeletal growth rate. Very recently, the authors of the present contribution carried out a detailed geochemical study of the skeletal aragonite of shallow and deep-water corals, both in the field and in aquaria. By focusing a laser ab-
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lation system on different skeletal portions of living samples of Cladocora caespitosa and Lophelia pertusa (Figure 2), we found that the distribution of most of the elements analysed, including Li, B, Mg, Sr and U is microstructure-related and largely depends on the different calcification mechanisms between the centres of calcification and the fibrous aragonite [24]. If these fine-scale variations translate into temperature, they provide a temperature range between 10 and 18°C, depending on the proxy and the calibration applied. Clearly, the chemical composition of the coral skeleton is not primarily controlled by the small temperature fluctuations in deep ocean sites, which are on the order of ± 2°C at most. In addition, we found that this micron-size scale heterogeneity is ubiquitous among different coral species, both zooxanthellate and azooxanthellate, implying similar physiological processes during the skeletogenesis [24]. For the first time we investigated the possibility of deconvolving environmental from physiological effects by analysing Li, Mg and Ca. Lithium and magnesium are highly correlated in shallow and deep-water corals and both seem to be similarly affected by the coral physiology. In order to correct for this “vital effect” Li/Ca ratios have been normalized for Mg/Ca ratios, providing a positive and highly significant correlation with the in-situ water temperature and suggesting a pure temperature control on Li/Mg ratio. Regardless of the exact mechanism that controls the uptake of Li and Mg, it seems clear that we can significantly improve the seawater temperature reconstructions with a precision of ± 0.8°C and with a particular sensitivity at low temperature such as those found in the deep ocean.
4.3
Nd-isotopes in deep-water corals as a novel ocean water mass tracer
At present, the main paleo-circulation tracers used in paleoceanography are stable carbon isotope ratios (δ 13 C) of dissolved inorganic carbon and Cd/Ca in foraminifera, sortable silt, 231 P a/230 T h ratios and radiocarbon. Unfortunately, all these proxies are subject to limitations and vary as a function of additional biological and environmental factors, including remineralization, biological productivity and air-sea gas exchange. Nd isotopes have been shown to be a promising new proxy to trace provenance and water mass mixing, being not fractionated by biological processes in the water column. 143 N d/144 N d ratios vary in the Earth as a result of β − decay of 147 Sm, and in the ocean the values reflect the age of the continental sources of dissolved Nd, acting as a fingerprint of the dissolved Nd source regions. Since the residence time of Nd in the ocean is in the order of 5001000 years [25], it can be effectively used as a tool to reconstruct the movement of water masses. This proxy has been successfully applied on marine cores, in the dispersed authigenic ferromanganese oxide precipitated in sediments [26, 27, 28], on foraminiferal shells [29] and on fossil fish teeth [30]. Only very recently, deepwater corals have been tested for their Ndisotopic composition, showing a close relationship with the composition of the ambient seawater [31, 32], and opening new possibilities to obtain water mass signals and quantify mixing of water masses via pared neodymium isotopes and radiocarbon analyses of absolutely dated (U/Th) fossil corals. The landmasses bordering the Mediter815
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ranean Sea have a distinct Nd-isotopic composition and this makes it possible to distinguish between different regional inputs [25, 33, 34, 35]. In particular, the Western Mediterranean Sea is characterized by Atlantic values entering the Mediterranean through Gibraltar whereas the Eastern basin displays more radiogenic values reflecting the input of major river such as the Nile [33]. The study of the Nd-isotopic composition of deep-water corals collected in strategic areas of the Mediterranean Sea (e.g. the Strait of Sicily) will provide a unique opportunity to investigate the modification of the Mediterranean circulation in the past in relation to climate change variations. Several living deep-water coral species (Lophelia pertusa, Madrepora oculata, Desmophyllum dianthus and Caryophyllia smithii) have been collected during a recent ISMAR-CNR cruise (MEDCOR cruise) in the Strait of Sicily at different depths in the water column, together with seawater samples. These samples will provide a unique opportunity to precisely calibrate the Ndisotopic composition of the aragonite exoskeleton with the ambient seawater. They will also increase the extensive CNR collection of Mediterranean deep-water corals already available for geochemical studies (Figure 3), and most of which already U/Th and AMS 14 C dated. Modern calibration experiments are essential to test a geochemical proxy that will be subsequently applied to the fossil coral collection. The Nd-isotopic signals of dated corals will be combined with more traditional tracers, such as radiocarbon, obtained from the same specimen in order to trace water provenance, mixing and rates of overturning circulation within the Mediterranean Sea during the Late Quaternary. Finally, those data may be further used to better 816
constrain reservoir effects for of marine benthic organisms.
4.4
14
C-dating
Recontructing the seawater pH evolution in the Mediterranean Sea using boron isotopes in corals
pH is an important marine parameter that enables to evaluate part of the ocean carbonate system. As a result of human activity the ocean uptakes more CO2 , which leads to a drop in pH, a process called “Ocean acidification”. To better understand the future impact of ocean acidification in the Mediterranean Sea, in terms of physical, chemical and biological properties, long-term continuous seawater pH records are needed. Unfortunately, very little is known about the present-day pH variability in the Mediterranean Sea and paleopH data are completely missing (CIESM, 2008). Recently, it has been suggested that the analysis of boron isotopes in calcifying organisms can provide accurate seawater pH reconstructions [36]. The theory behind the B-isotopes as pH proxy is quite well established and is based on relatively simple equations. In aqueous solutions boron exists as two species, boric acid and borate ion, with the proportion of the two species being pH dependent. The two species show large isotopic fractionation (∼ 20 ) due to different B-O vibrational energy and molecular geometry [37]. Since only the borate ion is postulated to be incorporated into marine carbonate [38], and given that the proportions of the aqueous species change according to pH, the boron isotopic composition in marine carbonate should also be a function of the ambient seawater pH during calcification [39]. While B-isotopes have been mainly mea-
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sured on foraminiferal shells to reconstruct paleo-seawater pH at glacial-interglacial scale [40, 41, 42], investigations on corals are still in their infancy, with only a few calibrations undertaken using cultured corals [43, 44] and three reconstructions using modern long-lived [4, 45] and midlate Holocene corals [46]. These first studies are very promising, indicating that the boron isotopic composition in corals actively responds to changes of the ambient seawater pH, with a precision better than Âą 0.02 pH units [45]. However, pH reconstructions seem to be potentially affected by the coral microstructures (i.e. centers of calcification and fibrous aragonite) [47, 15], and further studies are essential to determine the exact mechanisms controlling the B-isotopes variations in corals. Very recently, we used positive thermal ionization mass spectrometry to analyse the B-isotopic composition of two C. caespitosa colonies cultured for one year under two different pCO2 conditions: ambient pCO2 (ca. 400Âľatm) and elevated pCO2 (ca. 700Âľatm). The results clearly support the evidence that the B-isotopic composition of this coral species is pHdependent and open new perspectives in reconstructing the paleo-pH evolution in the Mediterranean Sea using long-lived C. caespitosa colonies [48].
the past is represented by the possibility to combine a series of geochemical proxies obtained from the same precisely dated specimen. As stated before, scleractinian corals can incorporate a suite of minor and trace elements in their aragonite exoskeleton and they can be precisely dated by U/Th and AMS 14 C methods. In addition, compared to sediment cores, which are often affected by bioturbation, corals offer a continuous and usually undisturbed record of the Late Quaternary. If suitably calibrated, the different geochemical proxies can be combined together to reconstruct in detail the paleo-circulation of the Mediterranean Sea and its physical and chemical state through time. Until now, due to the lack of the proper samples and precise and accurate analytical techniques it has been difficult to reliably reconstruct the paleoclimate in the Mediterranean Sea at highresolution. This novel multi-component approach will enable to answer specific paleoclimatic questions: How sensitive was the Mediterranean Sea to climatic events occurring during the last glacial and interglacial cycles in the Northern Hemisphere? How did the input/output water fluxes change through the sills, such as the Strait of Sicily? How did the Mediterranean circulation regime and the long term variability in the surface and deep layers change during the time? In the recent past the circulation of the Eastern Mediterranean 5 A geochemical multi- Sea undergone a drastic change with the Sea replacing for some years the proxy approach for re- Aegean previously dominating Adriatic Sea as the constructing the paleo- main source of the Eastern Mediterranean Waters (EMDW). This unique event climate in the Mediter- Deep is termed the Eastern Mediterranean Transient (EMT) and the real causes are still ranean Sea unknown although several hypotheses have The advantage of coral skeletons in recon- been proposed [49]. This type of circulastructing the environmental parameters in tion changes might have also occurred dur817
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ing previous periods and it is important to document these variations and investigate the link with the climatic events. These aspects remain largely unknown and one way to better understand the possible causes behind those changes is the detailed study of the coral geochemistry.
6
Conclusions
Shallow and deep-water corals from the Mediterranean Sea represent useful tools for the reconstruction of some of the most important marine parameters, such as seawater temperature, nutrient content, pH variations, seawater composition and water mass variability. Previous studies on the geochemical composition of living specimens of Cladocora caespitosa collected in the Mediterranean Sea revealed a strong correlation between Sr/Ca and B/Ca ratios with sea surface temperature [10, 6]. A refined study of the fine-scale geochemical pattern of C. caespitosa and the deep-water corals Lophelia pertusa enabled to identify a new promising temperature proxy (Li/Mg), which is independent of the vital effect, serving as a good candidate to reliably estimate temperature variations in the past. More precise and accurate analytical techniques, such as multi-collector ICP-MS or TIMS, can now be employed for the investigation of B and Nd-isotopes in corals. First results of the boron isotopic composition of scleractinian corals seems to re-
flect the ambient seawater pH during the calcification process. This is very promising for the reconstruction of the paleo-pH variability at different depths in the water column and for the understanding of the carbonate system. The Nd-isotopic composition of deep-water corals has been shown to be directly related to the ambient seawater composition. Based on this finding, Ndisotopes can be used to reconstruct the water mass dynamics, being able to track the different water masses. We now dispose of several powerful geochemical tracers and proxies that, combined together, can be applied to precisely dated corals to reliably characterize the physical and chemical state of the Mediterranean Sea in the past.
7
Acknowledgements
We would like to thank the Captains, Officers, the Crew and shipboard staff of the many CNR coral cruises onboard RV Urania. This research benefited from financial support from ESF Euromargins/Eurocore ”Moundforce”, EU ”Hermes”, EU ”Hermione”, EU ”MedatArchives”, Firb ”Aplabes” and Foundation Prince Albert II ”COMP” projects. We are grateful to Alessandro Remia, Matthias L´opez Correa, Lorenzo Angeletti, Alessandro Ceregato, Agostina Vertino and Helmut Zibrowius for great assistance during the various cruises. This is ISMAR scientific contribution n. 1724.
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[26] R.L. Rutberg, S.R. Hemming, and S.L. Goldstein. Reduced North Atlantic Deep Water flux to the glacial Southern Ocean inferred from neodymium isotope ratios. Nature, 405:935–938, 2000. [27] A.M. Piotrowski, S.L. Goldstein, S.R. Hemming, and R.G. Fairbanks. Intensification and variability of ocean thermohaline circulation through the last deglaciation. Earth and Planetary Science Letters, 225:205–220, 2004. [28] A.M. Piotrowski, S.L. Goldstein, S.R. Hemming, and R.G. Fairbanks. Temporal relationships of carbon cycling and ocean circulation at glacial boundaries. Science, 307:1933– 1938, 2005. [29] D. Vance, A.E. Scrivner, P. Beney, M. Staubwasser, G.M. Henderson, and N.C. Slowey. The use of foraminifera as a record of the past neodymium isotope composition of seawater. Paleoceanography, 19, 2004. [30] E.E. Martin and H.D. Scher. Preservation of seawater Sr and Nd isotopes in fossil fish teeth: bad news and good news. Earth and Planetary Science Letters, 220:25–39, 2004. [31] T. van de Flierdt, L.F. Robinson, and J.F. Adkins. Deep-sea coral aragonite as a recorder for the neodymium isotopic composition of seawater. Geochimica et Cosmochimica Acta, 74:6014–6032, 2010. [32] K. Copard, C. Colin, E. Douville, A. Freiwald, G. Gudmundsson, B. De Mol, and N. Frank. Nd isotopes in deep-sea corals in the North-eastern Atlantic. Quaternary Science Reviews, 29:2499–2508, 2010. [33] A.E. Scrivner, D. Vance, and E.J. Rohling. New neodymium isotope data quantify Nile involvement in Mediterranean anoxic episodes. Geology, 32:565–568, 2004. [34] C. Jeandel, T. Arsouze, F. Lacan, J.C. Dutay, N. Ayoub, and P. T´echin´e. Nd isotopic compositions and concentrations of the lithogenic inputs into the ocean: a compilation, with an emphasis on the margins. Chemical Geology, 239:156–164, 2007. [35] P. Montagna, S. Goldstein, M. Taviani, and N. Frank. Neodymium isotopes in biogenic carbonates from the Mediterranean Sea: reliable archives of water mass circulation. GEOTRACES Mediterranean workshop. 4-6 October 2010, Nice, France, 2010. [36] B. H¨onisch and N.G. Hemming. Ground-truthing the boron isotope paleo-pH proxy in planktonic foraminifera shells: partial dissolution and shell size effects. Paleoceanography doi:10.1029/2004PA001026, 19:2499–2508, 2004. [37] H. Kakihana, M. Kotaka, S. Satoh, M. Nomura, and M. Okamoto. Fundamental studies on the ion-exchange of boron isotopes. Bull. Chem. Soc. Jpn, 50:158–163, 1977. 821
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[38] A. Vengosh, Y. Kolodny, A. Starinsky, A.R. Chivas, and M.T. McCulloch. Coprecipitation and isotopic fractionation of boron in modern biogenic carbonates. Geochimica et Cosmochimica Acta, 55:2901–2910, 1991. [39] N.G. Hemming and G.N. Hanson. Boron isotopic composition and concentration in modern marine carbonates. Geochimica et Cosmochimica Acta, 56:537–543, 1992. [40] A. Sanyal, N.G. Hemming, G.N. Hanson, and W.S. Broecker. Evidence for a higher pH in the glacial ocean from boron isotopes in foraminifera. Nature, 373:234–236, 1995. [41] G. Foster. Seawater pH, pCO2 and [CO2 −3 ] variations in the Caribbean Sea over the last 130 kyr: A boron isotope and B/Ca study of planktic foraminifera. Earth and Planetary Science Letters, 271:254–266, 2008. [42] B. H¨onisch, N.G. Hemming, D. Archer, M. Siddall, and J.F. McManus. Atmospheric Carbon Dioxide Concentration Across the Mid-Pleistocene Transition. Science, 324:1551–1554, 2009. [43] B. H¨onisch, N.G. Hemming, A.G. Grottoli, A. Amat, G.N. Hanson, and J. Bijma. Assessing scleractinian corals as recorders for paleo-pH: Empirical calibration and vital effects. Geochimica et Cosmochimica Acta, 68:3675–3685, 2004. [44] S. Reynaud, N.G. Hemming, A. Juillet-Leclerc, and J.P. Gattuso. Effect of pCO2 and temperature on the boron isotopic composition of the zooxanthellate coral Acropora sp. Coral Reefs, 23:539–546, 2004. [45] G. Wei, M.T. McCulloch, G. Mortimer, W. Deng, and L. Xie. Evidence for ocean acidification in the Great Barrier Reef of Australia. Geochimica et Cosmochimica Acta, 73:2332–2346, 2009. [46] Y. Liu, W. Liu, Z. Peng, Y. Xiao, et al. Instability of seawater pH in the south China sea during the mid-late Holocene: evidence from boron isotopic composition of corals. Geochimica et Cosmochimica Acta, 73:1264–1272, 2009. [47] D. Blamart, C. Rollion-Bard, A. Meibom, J.P. Cuif, A. Juillet-Leclerc, and Y. Dauphin. Correlation of boron isotopic composition with ultrastructure in the deep-sea coral Lophelia pertusa: implications for biomineralization and paleo-pH. Geochem. Geophy. Geos., 8(doi:10.1029/2007GC001686), 2007. [48] J. Trotter, P. Montagna, M. McCulloch, S. Silenzi, S. Reynaud, G. Mortimer, S. Martin, C. Ferrier-Pag`es, J.-P. Gattuso, and R. Rodolfo-Metalpa. Quantifying the pH “vital effect” in the temperate zooxanthellate coral Cladocora caespitosa: validation of the boron seawater pH proxy. Earth and Planetary Science Letters, 303:163–173, 2011. [49] A. Theocharis, B. Klein, K. Nittis, and W. Roether. Evolution and status of the Eastern Mediterranean Transient. Journal of Marine Systems, 33-34:91–116, 2002. 822
Digital Terrain Model and Morpho-Structural Analysis of the Palinuro Seamount S. Passaro1 , G. Milano2 , S. Ruggieri1 , M. Sprovieri1 , E. Marsella1 , L. Giordano1 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, National Institute of Geophysics and Volcanology, Napoli, Italy salvatore.passaro@iamc.cnr.it Abstract We present a morpho-structural analysis of the Palinuro volcanic complex (Italy, Eastern Tyrrhenian Sea) by a high resolution Digital Terrain Model, obtained by swath bathymetry data collected during Cruise Aeolian 2007. A combined morphologicmorphometric approach helped to highlight that the volcanic complex has a subelliptical shape, extending 55 km E-W and 25 km N-S, with a well constructed summit at 84 m depth, rising from 3000 m depth from the NE bounds of the Marsili basin, and shows several volcanic feature, some of them not yet reported, grouped in the three major morpho-volcanic units displaced along its preferential elongation. The western and central sectors are probably emplaced by the coalescence of collapsed calderas, locally revived by recent volcanic activity, whereas the eastern sector reveals a strongly dissected and tectonized structure. The central sector is marked by distinct volcanic cones. The pronounced rim apparently not obliterated from erosional events of one of these cones suggests a volcanological rejuvenation of this sector. The different volcanic styles along the E-W direction of the Seamount maybe the product of different formation environments related to the ocean/continental transition from the Marsili Plain to the shallower escarpment of the sedimentary shelf of the Gulf of Salerno. The presented data add new information towards the better understanding of the regional volcanism and geodynamic processes in the south-eastern Tyrrhenian Sea.
1
Introduction
The Palinuro Seamount (hereafter PS) is a volcanic complex located in the South Eastern Tyrrhenian Sea at the transition zone between the Aeolian volcanic arc, the Marsili ocean-like basin and the Southern Italy passive continental margin. Its length, about 50 km along the E-W direction, and its position suggest that this volcanic complex should be a structural key for better understanding of the structural and vol-
canic processes affecting the SE Tyrrhenian Sea. Due to the scarce geophysical information and the particular position occupied by the Palinuro Seamount in the Tyrrhenian Sea (Figure 1), some key issues are currently unresolved. Among the open questions, it is not clear if PS is: a) a part of the Aeolian Volcanic Arc (e.g. [1]); b) independent from Aeolian Arc and controlled by pre-Pliocene tectonic structures (e.g. [2]); and/or c) a flank volcano complex of the Marsili ocean-like basin (e.g.
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Figure 1: A) Location map of the Palinuro Seamount (PS). MB=Marsili Basin, MS=Marsili Seamount, VB=Vavilov Basin, VS=Vavilov Seamount. B) Shaded relief of PS derived from the DTM. WZ, CZ and EZ indicate areas characterized by different morpho-volcanic styles. Capital letters indicate the location of the main volcanic features quoted in the text. [3]). In addition, the presence of swallow volcano-like seismicity detected between PS and the Calabrian coast [4] suggest that PS, or its southeastern sector, could be active. Despite the several geophysical surveys carried out in this sector of the Tyrrhenian Sea (e.g., [5, 6]), a detailed morphology of this seamount was not yet reported. On November 2007, a geophysical survey was carried out by IAMC-CNR in the southeastern Tyrrhenian Sea on board of the R/V Urania with the aim to identify potential active volcanic and hydrothermal vents. During the cruise, new high resolution multibeam data were acquired, allowing to build an high resolution Digital Terrain Model (DTM) of the Palinuro volcanic complex. The DTM interpretation, coupled with geomorphological and morphometric analysis, provided new insight on the role of this Seamount in the context of the eastern Tyrrhenian Sea ([7] and [8]). In this paper we summarize the main morphological and structural features of the Palinuro
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volcanic complex resulting by a joint morphometric and morphologic approach.
2
Geological setting of the Palinuro Seamount
The south-eastern sector of the Tyrrhenian Sea is characterized by the presence of two relatively deep (more than 3000m bsl) basins made by oceanic crust [9], as a result of a double stage of evolution (e.g., [10]). The first stage was ruled by E-W rifting that led to the formation of the older Vavilov ocean-like basin (4.3 â&#x20AC;&#x201C; 2.6 Ma, e.g.[5]). A subsequent ESE-directed extension affected this sector, thus leading to the formation of the Marsili ocean-like basin. The older Vavilov basin include the Vavilov Seamount, which is characterized by the presence of geomagnetic anomalies with inverted shapes that testify an emplacement during a inverted geomagnetic epoch [11]. The younger Marsili basin (1.8 â&#x20AC;&#x201C; 0.2 My, e.g. [5]), characterized by a
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roughly circular area of some 8000 km2 and a flat seafloor at a mean depth of 3500 m, and is morphologically dominated by the Marsili Volcano, a large N15°E trending seamount that results to be quite similar to mid-ocean slow-spreading ridges [12]. Large-scale subduction and roll-back of the Ionian lithosphere and extensional processes produced the onset of volcanism throughout the Tyrrhenian Sea and the surrounding coasts. Aligned along a global horseshoe shape around the Marsili Basin is located the Aeolian Arc, constituted by seven sub-aerial volcanoes (Stromboli, Vulcano, Lipari, Salina, Alicudi, Filicudi Panarea) and by several submarine edifices (i.e. Glabro, Alcione, Lametini, Alicudi, Filicudi, Eolo, Enarete, Sisifo, etc.), essentially made by calcalkaline and shoshonitic rocks. The Palinuro Seamount (0.8 - 0.3 Ma,[13]) is a volcanic complex that lies between the Marsili basin to the south, the Southern Appenninic Chain to the east and it is bounded by a sedimentary basin to the north (Figure 1A and B). This zone represents the transition between the sedimentary shelf of the Salerno Gulf, which is the only sector characterized by the absence of volcanism, the Marsili ocean-like basin and the Aeolian Calcalkaline volcanism. The Southern Appenninic Chain represents a part of the arcuate orogenic belt including to the south the Calabrian Arc and the Sicily Maghrebides (Figure 1A) whose evolution is associated to the simultaneous subduction of a west-dipping lithospheric slab and the back-arc extension in the Tyrrhenian Sea (e.g. [14]). Previous studies on PS focused mainly on volcanic products. Bottom sampling ([15] and references therein), shown lavas and sediments associated with volcanism (i.e. crusts and iron- and manganese-bearing nodules). Multibeam swath bathymetry
[12] described the general physiographic characteristics of the volcanic complex, documenting a length of about 50 km along the W-E direction. According to [16], PS was emplaced in a complex way and probably over a long period of time (0.5 Ma, [13]), along E-W striking deep-seated fault system. This fault may represents the northward tear fault of the Calabrian-Arc subduction (e.g. [2]), but its role in the geodynamic context of the south-eastern Tyrrhenian Sea is still debated.
3
Data and methods
Data presented in this paper were acquired during the second leg of the multidisciplinary survey R/V Urania Aeolian 2007”cruise, managed by IAMC-CNR in November, 2007. Multibeam swath bathymetry was performed using a keelmounted RESON Seabat 8160, 50 kHz, 3000 m range, 126 beams, 150° aperture. The data were recorded by the PDS2000 navigator system, positioning system by DGPS and real Time MRU data provided by Kongsberg’s MAHRS and SGBrown gyrocompass. The CDT data collected by Seabird 911 were used for proper speed of sound profiling. Data were processed by the PDS-2000 software, and Digital Terrain Models (DTM) at the resolution of 25 m (UTM, WGS84) were produced, covering an area of more than 1000 Km2 . Finer resolution grids were provided on shallower areas. In order to obtain the most objective interpretation of the volcanic features of PS, a combined morphologic-morphometric approach has been utilized. The DTM was used to produce morphometric. Computations, carried out by using both well-known available GIS tools and “ad hoc” fortran rou825
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tines (e.g., [17]). As is well know in literature (e.g. [18]), gradient, aspect, tangential curvature and profile curvature are useful tools to obtain an overall view of the seamount morphologic items. Further details on the morphologic aspects of PS have been obtained using â&#x20AC;&#x153;Opennessâ&#x20AC;? [19] is defined as a maximum or minimum angle over a specified region, elevation histogram plots and an elevation versus average slope. Elevation histogram plots was considered for a wide range of application ([17], and references therein) such as the identification of marine terraces, since slight dipping surfaces become relative maxima in the elevation histograms (e.g., [20]). Elevation versus average slope (e.g. [21]) is a morphometric tool that helps to classify slopes as a function of depth at which they occur, and help to identify potential flat or dipping terrains when they systematically arises at specific depths.
4
Morphometry of the Palinuro Seamount
The plots of the morphometric analysis (Figure 2A, B, C and D) outline the existence of three groups of volcanic morphologies (e.g. Figure 2A and D) characterized by different dimensions (larger in the western zone) and displaced as a sequence of partially overlapping or coalescent volcanic features on the seamount ridge. Such a pattern suggests an W to E gradual change in the morpho-structural setting. The ridge crest appears to be disarticulated, in correspondences of the transition between different groups (Figure 2B), very likely as the results of the effect of morpho-structural controls. Upward, openness maps of PS (Figure 2E) 826
show ridge crests while downward openness is useful to highlight valleys, providing detailed views of the terrain. Valley reticulate is reported by using the tangential curvature plot (Figure 4C), whereas both upward openness map (Figure 4E) and profile curvature (Figure 4D) allow to exactly locate major crests interpretable as faults, major domes and calderic rims. Erosional base levels and general trends of the whole PS was established through stack of profiles (along and across PS ridge) and through the use of elevation histograms, which consist in an elevation versus percentage of the area plot. Since PS strikes N100°E, we extracted elevation profiles for DTM following this preferential orientation (some examples are reported in Figure 3a and b) and its conjugate pattern (Figure 3d and e). Stacked profiles were obtained by calculating average values on distance versus elevation plots, taking into account 25 m spaced elevation profiles extracted from DTM. Stack computations pointed up the overall northward and eastward shoaling nature of PS. PS elevation histogram show a clearly visible peak in correspondence of the 1582 m bsl depth, that represents the northern erosional base level of the Seamount (Figure 3g), whilst is 2807 m bsl for the southern slope. This asymmetry, also observable on the stacked elevation profiles (Figure 3C and F), is due to the presence of a sedimentary basin bounding the north side of PS [7]. The elevation versus average slope (Figure 3h) recognizes 2 maxima and 2 minima. Minima slope values evidence the erosional base levels, that agrees and confirm what previously reported (Point 2, 1582 m bsl, and Point 3, 2807mbsl in Figure 3g), while maxima are located at 205 m (30%, point 1 in Figure 3h) and at 3000 m bsl (40%, point 2). A number of morphological, ecologi-
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Figure 2: Morphometric computations carried out from the DTM of PS. A) Gradient, B) Aspect, C) Tangential curvature, D) Profile curvature, E=Openness (upward), F=Slope. Location is in Figure 1. cal, and hydrological processes are well described by the use of local slope and aspect. Largescale and regional trends, slope instability and anisotropy can also be enhanced by the use of such parameters [4]. Following existing standards, we plotted the aspect versus slope computation (Figure 3i) that emphasizes the existence of a SSW predominant component.
putations (Figures 2 and 3), we described the main volcanic features of PS (see [7] and [8] a for details). The DTMdepths range from 3200 to 84 m, and show a roughly elliptical shape extending about 55 km along N100째E and 25 km in the NS direction. The morphology reveals a very articulated summit, characterized by different cone-like, flat and amphitheatrelike structures rather than a single volcanic edifice (Figure 1B). In particular, 5 Main morphological amphitheatre-like shapes are present in the westernmost sector of DTM (WZ), conifeatures cal shapes in the central sector (CZ) and Through the DTM analysis and taking into dissected shapes in the eastern zone (EZ). account the previously morphometric com- These differences are also well outlined by 827
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Figure 3: Location (a), plot (b) and average values (c) of the N100°E structural trend elevation profiles extracted from the DTM of PS; d, e and f: same as above but referred to the N10°E structural trend. g is the percentage distribution of depths (elevation histogram) whereas h and i are the elevation versus slope plot and the aspect by slope distribution, respectively. morphometric computations.
5.1
The western sector: “amphitheatre”-like shapes
Two morphostructures, with a global “amphitheatre”-like shapes, are present in this sector (see panels W1 and W2 in Figure 4). Due to their general morphology, the presence of some rim segments and the global concave profile, we interpreted these features as collapsed calderas. Feature A (panel W1; see [7] for details) consists in a 4x4 km flat-top surface located at about 828
1600 m depth, characterized by the presence of an inner hill inside the flat surface. Feature B (panel W2; see [7] for details) is particularly interesting. It is characterized by a semi-elliptical shape, with a N 115° E elongation, and wide extension (7.5x4.5 km, i.e. 3 times the extension of the Monte Somma-Vesuvio volcanic complex caldera) at an average depth of 1200 m bsl, (Figure 2, 3). This area is bounded by a crest, the depth of which ranges between 1200 (SSE) to 800 (NW) m bsl, and is abruptly interrupted in the NE part by a relief of about 500 m (Figure 4B).
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Figure 4: Main volcanic features of PS. W1 and W2 are calderas which characterized the western zone whereas C1/C2, C3 and C4 are the volcanic cones detected in CZ. E1 represents the tectonized volcanic feature inside EZ. Zone and frame locations are in Figure 1 (panel B). Its shape indicates a possible analogy with calderas formed through mixed explosions and lateral collapses of volcanic flanks (e.g., Mount St. Helen), because of the global horseshoe top morphology and the marine slide evidences in correspondences of the southern side-slope (40-55 degrees, Figure 2F). Residual rim are present northwestward and from S to SE with respect the centre of collapsed caldera. The compound NE rim of B (panel W2) has been structurally modelled by erosion and by roughly N60°E faults. The south-western scarp of B shows a composite morphology interpreted as a ridge formed by erosion.
5.2
The central sectors: canic cones
vol-
The central sector of PS present a large number of cone-like morphologies and ridges. Two volcanic cones with flat circular tops of about 7-800 (C1) and 2,500 m (C2) in diameter are characterized by a 30° slope-break value in all directions up to 1000 m. C1 reaches 157 m bsl, whereas C2 84 m bsl. C1 and C2 top surfaces are slightly WNW dipping. Passaro et al., [7] proposed the name “Piotr’s cones” for these two edifices in honour of the researcher Piotr Mikejcik of the IAMCCNR of Mazara del Vallo, who lost his life in the Sicily seas during the sinking of the Thetis oceanographic vessel. On C top, evidences of sea-level still-stand relict morphologies (marine terrace) located at 90, 829
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100 and 123 m bsl were detected as a consequences of global sea-level lowering of the past [17]. Feature C3 is characterized by a roughly crater-like top which extends for about 2.1 km2 and shows flanks strongly marked by the presence of morphostructures with a N10°E and N115°E preferred orientations, whereas the western margin seems bounded by a N-S trending structure (Figure 4). In its central part, this edifice shows a 75 m deep crater with top, located at 570 m bsl (Figures 2, 4). Relief C4 shows an elliptical shape with a N110°E elongated major axis (Figures 2, 4) characterized by a multifaceted top surface. This latter is located at 600 m bsl and seems strongly tilted toward N297°E, with a N30°E structural fabric (Figures 2, 4).
5.3
The Eastern sector: younger tectonized structures
The main feature of the EZ is the presence of a relative high (E1) placed at 950 m bsl, showing an elliptical shape elongated N110°E, with major and minor axes of 6,000 and 4,500 m, respectively (Figure 4). The well defined top of E1 is marked by a N-S/N10°E distinct ridge, whereas southeastward it progrades to a sector of the slope where a deep border of detachment is discernible (Figure 1B and 2).
5.4
Slope instability of the major Piotr's cone top
From a morphological point a view, caldera-stage can be considered as the more ancient volcanic-stage with respect to the conic-shapes. Absolute dating of volcanic products on the area are lacking, however, on the basis of morphological ev830
idences we inferred that PS volcanic features are progressively younger from W to E. The morphological differences between W and E sectors may be due to the emplacement of PS over the escarpment, and to a plausible differentiation of volcanic products both due to its shoaling nature and to the global differentiation of volcanic supply related to the landward (east) progressive crustal thickening. Morphometric computations remark the differences between EZ, CZ and WZ, and also lead to hypothesize a general, relative uplift of the CZ, in particular, with respect to WZ. As a consequence of the: i) presence of marine terraces of abrasion on the shallower cone [8], ii) CZ (relative) block uplift [7], iii) presence of hydrothermal vents on PS top (see [15], and [22]) and iiii) a slight tilting of WZ [8], slope instability is marked in this sector, mainly over shallower C1 and C2 cones. These aspects suggest that some of the CZ valleys could have been produced by lateral spreading phenomena.
6
Concluding remarks
The main morphological and structural features of the Palinuro volcanic complex, derived by the analysis of a new, high resolution DTM resulting by more than 1,000 km2 of multibeam data, have been reported. Morphometric computations allowed to well highlight morphological and volcano-tectonic features. Data clearly show that the Palinuro volcano complex has a roughly elliptical shape, extending for about 55 km along N100°E and 25 km in N-S directions, characterized by a very multifaceted morphology. Its articulated summit consists of a group of overlapped and/or coalescent volcanic edifices, that have been interpreted as a sequence
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of volcanic cones and collapsed calderas. Relic ridge-like shapes are identifiable both in the central and in the western sectors. The central sector of the seamount is characterized by the presence of distinct volcanic cones. One of these clearly shows a volcanic crater with a pronounced rim not obliterated by erosional events. The eastern zone seems to be completely different from the other zones, being structurally controlled and representing a potential el-
ement of transition toward the mainland. Different erosional base levels are found both in the N-S and E-W directions. Varied volcanic styles (e.g. amphitheatre-like and cone-like shapes, tectonized structures) can be identified along the main axis of volcano complex, possibly reflecting changes of the environments of formation of the seamount in terms of water column depth and changes in magma supply, due to the eastward crustal thickening.
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High Resolution Multichannel Seismic Survey in the Gulf of Pozzuoli V. Di Fiore1 , M. Sacchi1 , A. Rapolla2 , F. Molisso1 , A. Cicchella1 , E. D’Aniello2 , V. Spiess3 1, Institute for Coastal Marine Environment, CNR, Napoli, Italy 2, University of Napoli “Federico II”, Napoli, Italy 3, University of Bremen, Bremen, Germany vincenzo.difiore@cnr.it Abstract We present the preliminary results of the oceanographic cruise CAFE-07 Leg 3, conducted in the Napoli and Pozzuoli Bays in January 2008, on board of the R/V URANIA of the CNR. During the cruise it was carried out the acquisition of a grid of ca. 800 km of high-resolution multichannel reflection seismic profiles, working in parallel with two source systems, the seismic source GI-GUN Sodera of 1,7 l and the mini GI-GUN Sodera of 0,4 l, and hydrophones operating simultaneously at different frequency ranges. The MTU Shallow Water streamer have a length of 50 m, with 48 channels and the hydrophone spacing of 1m; the streamer of H.Villinger working group have a length of 100 m with 16 channels over 6 groups of a length of 6.25 m. Here we present only the results related to the acquisition by GI Gun Sodera seismic source using the MTU shallow water streamer. Data processing, aimed at reduction of random noise of the dataset, included removal of unwanted coherent events. Several sequences processing were applied on the dataset to improve the stack section. The aim of the cruise was the understanding of the stratigraphic-structural setting of the Pozzuoli Bay area, with specific reference to the major offshore volcanic features; the results of this research represent a significant contribution to the understanding of the Late Quaternary evolution of the Campi Flegrei area.
1
Introduction
During the cruise CAFE-07 conducted onboard of the R/V Urania, from 10 to 21 January, 2008, in the Naples and Pozzuoli Bays, over, 150 high resolution multichannel reflection seismic profiles of were collected [1]. This area represents a very active segment of the eastern Tyrrhenian margin during the late Quaternary and may be regarded as a key region to analyze the interplay between tectonics and explosive
volcanism associated with rifted back-arc margins. A total of ca. 800 km of seismic profiles were acquired during CAFE-07 cruise. The seismic grid consisted in a dense network of profiles with average distance of about 150 m between navigation routes, in order to obtain a quasi 3-D seismic coverage of the Pozzuoli Bay. Several seismic survey have been carried off this segment of the Campania region in the last 40 years. However all previous multichannel seismic in-
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vestigations had a lower resolution with respect to the survey presented in this study. In this paper, we illustrate the results of the seismic processing results and discuss the preliminary interpretation of a selected multichannel profiles from the Gulf of Pozzuoli. The sequence of data processing was specifically designed in order to improve the imaging of volcaniclastic units and reduce the effects of multiple reflections and reverberation.
2
Geological settings
The Gulf of Pozzuoli extends offshore the volcanic district of Campi Flegrei, a province of active volcanism located NW of the city of Naples. The area lies along the continental margin of the Campania region, Eastern Tyrrhenian Sea that evolved during the Quaternary as extensional basin filled by siliciclastic, epiclastic and volcaniclastic deposits (e.g. [3, 2]). The basin is roughly elongated in a NEâ&#x20AC;&#x201C;SW direction across the continental magin, and is dissected into minor horst and graben structures [4, 5, 6, 7]. The regional tectonic pattern is characterized by dominantly NESW, NW-SE and subordinately N-S trending faults, and associated volcano-tectonic lineaments. The Campi Flegrei itself is a complex volcanic structure, that has been interpreted as a nested caldera [8] resulting from two main collapse phases, namely associated with the eruption of the Campania Ignimbrite (CI, 39 ka) [9] and the Neapolitan Yellow Tuff (NYT, 15 ka) [10]. Structural data show that the Campi Flegrei caldera has been subject to a generalized subsidence since 15 ka B.P: and an inner-caldera late stage resurgence, associated with brittle deformation and fracturing of the NYT 836
caldera floor into separated blocks (e.g. [8, 11]). After NYT eruption, a period of intense activity occurred in the area between 15 and 11 ka, separated by quiescence intervals. According to some authors (e.g. [11]), volcanism and quiescence intervals are strictly related to the phases of caldera deformation. The intense activity of the last 12 ka has been grouped into three periods: Epoch I (from 12 to 9.5 ka), with more than 34 explosive eruptions and distribution of the vents along marginal faults of the NYT caldera; Epoch II (from 8.6 to 8.2 ka) with at least 6 explosive eruptions and vents located along the north eastern structural margin of the NYT caldera; Epoch III (from 4.8 to 3.8 ka) with ca. 16 explosive and 4 effusive eruptions and location of vents in the north eastern sector of the NYT caldera. These unrest periods were separated by two period of relative quiescence, at 3.5 ka and 1 ka, respectively. Other studies also suggest a genetic relationship between the geometry of the regional tectonic pattern and the location of volcanic centers [12, 13, 14, 15, 16, 15, 17]. Particularly, in the Pozuouoli Bay a series of volcanic banks and submarine volcanic edifices, (e.g. Pentapalummo, Nisida and Miseno Banks), may be observed. These volcanic features are ostensibly aligned along the NE-SW trending Magnaghi-Sebeto fault line as well as along the NW-SE structural discontinuities. Magnetic and gravimetric data from of the Bay of Naples [15] seem to confirm a tectonic control on the Campania volcanism, as suggested by the consistency between the location of the main volcanic structures, the pattern of magnetic anomalies, and the major NE-SW and NW-SE trending fault systems [15, 18, 17]. The anal-
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Figure 1: Flow diagram of the sequence processing used in paper. ysis of airborne gravimetric and magnetic data [19], revealed a sub-circular structural depression in the central part of the Pozzuoli Bay, that has been interpreted as the collapsed caldera associated with the NYT eruption [20]. On the basis of the substantial correspondence between seismic features and gravimetric anomalies, Bruno [16] has proposed that the NE-SW (Magnaghi-Sebeto line) and the NW-SE faults systems actually bound the Campi Flegrei (NYT) caldera.
3
Data Processing
The procedure adopted for the processing of multichannel seismic data consisted of several phases [13]. The basic geometry and the parameters of the acquisition system utilized during the cruise are summarized below: - Streamer length 50m; - Offset 6m; - Hydrophone interval 1m; - Sampling interval 0.250 ms; 837
Marine Geology
Figure 2: Localization of the profiles 90 and 108 on the digital terrain model of Naples Bay (from [2] modified). data deconvolution. Spiking pre-stack and - Shot interval 6m; post-stack deconvolution was applied to - Time window recording 3 s; - Seismic source: mini-GI gun SODERA improve temporal resolution, while predictive deconvolution [23] was utilized to pre1,7 l. dict and remove repetition in the recorded Prior to preliminary processing, a data seismograms. quality control as well as the definition of In addition to pre-stack and post-stack the geometry of the acquisition system and predictive deconvolution, another multiple the trace editing has been conducted. Au- suppression technique, based on FK dip filtomatic Gain Control (AGC) was applied tering, was applied. FK multiple attenua[21] for a best visualization of traces. In or- tion is a process combining: der to reduce the spatial aliasing, especially the geometric effect of migration [22] a 1. velocity analysis; trace interpolation was performed. The 2. forward Normal Move Out (NMO) correction; procedure for transforming the data to the 3. FK dip filtering, and; Tau-P domain (Diebold and Stoffa, 1981), 4. inverse NMO correction. consists in the weighting of each transformed sample, based on a coherency measure, and then the inversion of the transformed data to the new spatial sampling. A significant of the processing was spent on 838
This process provides a further gain in multiple energy suppression over stacking. Filtering performs more quickly than other multiple suppression methods, and it
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Figure 3: Profile 90 and its interpretation. TWT: two-way travel time (ms); CDP: Common Depth Point ensemble number is useful for multiple suppression on prestack data. In the FK domain, it is possible to discriminate between primary and multiple events, on the basis of their different velocity. A qualitatively high velocity analysis was performed by the semblance function that is defined as a normalized crosscorrelation. An effective velocity analysis is a key prerequisite for obtaining a good seismic stack by increasing the S/N ratio. As a final step in the pre-stack process, we have applied the Dip Moveout (DMO) correction [24]. In fact, the application of DMO to pre-stack data together with the post-stack time migration data emulates a full pre-stack migration, with the advantage of being computationally faster. After DMO, the stacked data were migrated in time, using Kirchhoff migration [25]. The last process applied to the post-stacked data was Frequency-Space (FX) deconvo-
lution that is a reliable multi-channel noisereduction filter. It preserves the most dominant dipping energy while removes random noise or dips with very low energy [26]. Figure 1 summarizes the sequence processing adopted in this work. In summary, our processing aimed at:
1. reduction of random noise in the data 2. removal of undesired coherent events 3. reduction of spatial aliasing by means of trace interpolation on Common Shot Point (CSP) gathers 4. improvement of resolution of the seismic wavelet with spiking deconvolution algorithms 5. repositioning of reflectors according to their correct location in the space-TWT domain by means of dip moveout and post-stack time migration. 839
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Figure 4: Profile 108 and its interpretation. TWT: two-way travel time (ms); CDP: Common Depth Point ensemble number
4
Interpretation
Seismic profiles 90 and 108 are two examples of a grid of the multichannel reflection seismic profiles acquired during oceanographic cruise CAFE-07 [1]. The location of profiles is displayed in Figure 2. A detail of the profile 90 (Figure 3) illustrates the volcanic structure of Nisida Bank and a series of volcaniclastic units (here referred to as Nisida complex), located between Nisida Bank and Nisida Island. The Nisida Bank represents a remnant of a tuff cone. This unit lies above an erosional unconformity that can be associated with the major sea-level fall that occurred during the last glacial maximum and is covered in turn by uppermost Quaternary -Holocene deposits (18-6 ka BP). The Transgressive Systems Tract (TST) unit is represented by siliciclastic deposits 840
associated with the rapid rise of sea level and with the following sea level highstand. These siliciclastic units, characterized by moderate thickness, are partly fed by the materials deriving from the dismantling of the surrounding volcanic areas. A detail of the profile 108 (Figure 4) shows a recent epi-superficial magmatic intrusion, the M. Dolce-Pampano structure, that overlies an older volcanic unit, and is covered in turn by the Late Holocene sediments.
5
Conclusions
The seismic data presented in this study were collected with an up-to-date, 24-bit acquisition system that allowed us obtaining very high resolution seismic profiles along with a good penetration of the seismic signal at depth, thus providing unprecedented detailed views of the offshore
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stratigraphy and structures. Most of the previous multichannel surveys had in fact much lower resolution and /or were focused to investigate much deeper structures (e.g. [5]). The seismic and sequence stratigraphic interpretation of the acquired seismic grid revealed a complex stratigraphic and structural setting, dominated by the occurrence of volcanic bodies and siliciclastic depositional units, mostly deriving from the dismantling of the adjacent vents and volcaniclastic units [3, 27, 1]. Volcaniclastic units typically are interbedded with siliciclastic units of the Late Qua-
ternary depositional sequence which represent in terms of chronostratigraphy, the marine deposits that formed between the onset of the post-120 ka BP sea-level fall and the present day. The interpreted sesimic profiles intersect some of important volcanic structures such as Nisida Bank (18-6 ka), the Nisida Complex (6-4 ka) and the M.Dolce-Pampano structure (< 6 ka). Our results reveal an extremely complex stratigraphic and structural setting, likely controlled by the current interplay between volcanic and sedimentary processess.
References [1] M. Sacchi, G. Alessio, I. Aquino, E. Esposito, et al. Risultati preliminari della campagna oceanografica CAFE-07 - Leg 3 nei Golfi di Napoli e Pozzuoli, Mar Tirreno Orientale. Quaderni di Geofisica, 61:1–25, 2009. [2] B. D’Argenio, A. Angelino, G. Aiello, G. de Alteriis, et al. Digital elevation model of the Naples bay and adjacent areas, eastern Tyrrhenian Sea. pages 21–28, 2004. [3] A. Milia. Aggrading and prograding infill of a peri-Tyrrhenian basin (Naples Bay, Italy). Geo-Marine Letters, 19:237–244, 1999. [4] E. Carrara, F. Iacobucci, E. Pinna, and A. Rapolla. Gravity and magnetic survey of the Campanian volcanic area. Southern Italy. Boll. Geof. Teor. Appl., XV:39–51, 1973. [5] I. Finetti and C. Morelli. Esplorazione sismica a riflessione dei Golfi di Napoli e Pozzuoli. Boll. Di Geof. Teorica ed Appl., XVI(62-63):175–222, 1974. [6] Agip. Carta aeromagnetica d’Italia (scala 1:500.000). 1981. [7] A. Milia and A.M.M. Torrente. The influence of paleogeographic setting and crustal subsidence on the architecture of ignimbrites in the Bay of Naples (Italy). Earth and Planetary Science Letters, 263:192–206, 2007. [8] G. Orsi, S. De Vita, and M. di Vito. The restless, resurgent Campi Flegrei nested caldera (Italy): constraints on its evolution and configuration. Journal of Volcanology and Geothermal Research, (74):179–214, 1996.
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[9] B. De Vivo, G. Rolandi, P.B. Gans, A. Calvert, et al. New constraints on the pyroclastic eruptive history of the Campanian volcanic Plain (Italy). Mineral. Petrol., 73:47–65, 2001. [10] A.L. Deino, G. Orsi, S. de Vita, and M. Piochi. The age of the Neapolitan Yellow Tuff caldera-forming eruption (Campi Flegrei caldera, Italy) assessed by 40Ar/39Ar dating method. J. Volcanol. Geotherm. Res, 133:157–170, 2004. [11] M.A. Di Vito, R. Isaia, G. Orsi, J. Southon, et al. Volcanism and deformation since 12,000 years at the Campi Flegrei caldera (Italy). Journal of Volcanology and Geothermal Research, 91:221–246, 1999. [12] V. Acocella, F. Salvini, R. Funiciello, and C. Faccenna. The role of transfer structure on volcanic activity at Campi Flegrei (Southern Italy). Journal of Volcanology and Geothermal Research, 91:123–139, 1999. [13] PP. G. Bruno, V. Di Fiore, and A. Rapolla. Seismic reflection data processing in active volcanic areas: an application to Campi Flegrei and Somma Vesuvio offshore (Southern Italy). Ann. Of Geoph., 45(6):753–768, 2002. [14] P.P. Bruno, A. Rapolla, and V. Di Fiore. Structural setting of the Bay of Naples (Italy) seismic reflection data: implications for Campanian volcanism. Tectonophysics, 372:193–213, 2003. [15] M. Secomandi, V. Paoletti, G. Aiello, M. Fedi, et al. Analysis of the magnetic anomaly field of the volcanic district of the Bay of Naples, Italy. Marine Geophysical Researches, (24):207–221, 2003. [16] P.P. Bruno. Structure and evolution of the Bay of Pozzuoli (Italy) using marine seismic reflection data: implication for collapse of the Campi Flegrei caldera. Bull. Volcanol., 66:342–355, 2004. [17] V. Paoletti, M. Secomandi, M. Fedi, and A. Rapolla. The integration of magnetic data in the Neapolitan volcanic district. Geosphere, 1(2):85–96, 2005. [18] S. Judenherc and A. Zollo. The Bay of Naples (southern Italy): Constraints on the volcanic structures inferred form a dense seismic survey. Journ. Of Geoph. Res., 109, 2004. [19] Agip. Italia, carta magnetica, anomalie del campo residuo, scala 1:500.000, foglio H. 1982. [20] G. Florio, M. Fedi, F. Cella, and A. Rapolla. The Campanian Plain and Phlegrean Fields: structural setting from potential field data. J. Volcanol. Geoth. Res., 91:361–379, 1999. [21] R.E. Sheriff and L.P. Geldart. Exploration Seismology. page 419, 1995.
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[22] V. Bardan. Trace interpolation in seismic data processing. Geophysical Prospecting, 35(4):343–358, 1987. [23] K.L. Peacock and S. Treitel. Predictive deconvolution: theory and practive. Geophysics, (34):155–169, 1969. [24] O. Yilmaz and J.F. Claerbout. Prestack partial migration. Geophysics, (45):1753– 1777, 1980. [25] O. Yilmaz. Seismic Data Processing. 1987. [26] P. Cary and W. Upham. Noise attenuation with 3D FXY deconvolution. CSEG Convention Abstracts, pages 22–23, 1993. [27] A. Milia, M.M. Torrente, M. Russo, and Zuppetta A. Tectonics and crustal structure of the Campania continental margin: relationships with volcanism. Miner. Petrol., 79:33–47, 2003.
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Marine Geology in the Region of the Messina Straits, and a Puzzling Tale of Faults, Earthquakes and Tsunamis A. Argnani Institute of Marine Sciences, CNR, Bologna, Italy andrea.argnani@ismar.cnr.it Abstract The Straits of Messina is one of the most tectonically active areas of the Mediterranean, and has been the site of the destructive 1908 Messina earthquake. In spite of the hazard potential of this marine area, studies based on modern geophysical data are still lacking. A marine multichannel seismic survey was purposely carried out with the aim to outline the fault pattern in the area of the Messina 1908 earthquake, and to better understand its significance within the tectonic frame of the region. Within the Messina Straits, faults have been imaged on the Calabrian side, with a 30 km long NW-trending fault, located at the SW tip of Calabria, that is affecting the sea floor, whereas we did not image any extensional fault plane attributable to the Taormina Fault, on the Sicilian side of the straits. The geodynamic implication is that extension in south-eastern Sicily, on the Ionian side of the Hyblean Plateau, and extension in southern Calabria and Messina Straits belong to two different tectonic systems and, therefore, cannot be mechanically linked. Finally, the damages produced by the1908 ground shaking were aggravated by the effects of a remarkable tsunami, with up to 11 m of run-up height, that followed the earthquake. The origin of the tsunami associated with the 1908 earthquake is not yet fully understood, but geological and geophysical evidences substantially undermine a recent proposal that the 1908 tsunami originated by a large landslide offshore Giardini Naxos.
1
Introduction and Geological Setting
The Messina Straits and its surroundings are one of the most tectonically active areas of the Mediterranean, as indicated by several lines of geological and geophysical evidence. Relatively large earthquakes struck the area in historical times (Figure 1), although the recurrence time for 1908-type events (Mw=7.1) seems to be about 1500 yr [4]. The Sicilian and Calabrian side of the straits are characterised by uplifted ma-
rine terraces. The flight of emergent marine terraces along the coast of Sicily, from Taormina to Briga, were uplifted since 125 ka with rates of 1.07 mm¡yrâ&#x2C6;&#x2019;1 [5]. The occurrence of a large and active extensional fault named Taormina Fault, running offshore, has been inferred on the basis of coastal geomorphology (Figure 1 [6]). According to some investigators (e.g., [7]), the systems of faults of southern Calabria and south-eastern Sicily, on the Ionian side of the Hyblean Plateau (e.g., [8]), both associated to large and destructive earthquakes, can be linked through the Taormina
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Figure 1: Map of the Messina Straits with location of relevant historical earthquakes (after [1]; [2]). Dashed boxes represent poorly constrained events. Fault. However, along the belt corresponding to the inferred Taormina Fault a lack of seismicity is indicated by historical data and recent instrumental records [9] (Figure 1). The assumption that the hypothesized Taormina Fault is part of a single rift system, connecting Calabria to south-eastern Sicily, implies that it represents one of the most hazardous seismic gaps in Italy, a potential site for large future earthquakes. A flight of uplifted marine terraces characterizes also the Calabrian coast of the Messina Straits. Twelve to fourteen orders of terraces have been identified, with the highest terraces, dated Middle Pleistocene, having an elevation of about 1350 m above sea level [10], with an stimated uplift rate of
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about 1.0-1.3 mm·yr−1 in the last 300-400 ka. GPS velocities [11], indicate a NW-SEdirected extension at a rate between 1.7 to 3 mm·yr−1 between the Sicilian and Calabrian sides of the Messina Straits. In particular, GPS-derived interseismic strain appears to fit aseismic slip along a 30°, SEdipping normal fault, locked above about 8 km (Serpelloni, p.c. 2009). Onshore structural studies (e.g., [3]) have shown the occurrence of a fault system that runs along the Sicilian coast, but the faults that have been most active during the Pleistocene are those located on the Calabrian side of the Messina Straits [3]; these faults belongs to two main sets trending NW-SE and NESW and are not obviously related to the N-
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Figure 2: Map showing the traces of fault planes proposed for the Messina 1908 earthquake, with the authors annotated. S trend of the faults inferred to be responsible for 1908 Messina earthquake (Figure 2). Most authors agree that the extensional faults bounding the Reggio Calabria basin have been active through Late Pliocene and Early Pleistocene. However, subsequent fault activity is debated, and some authors suggest that the faults were not active since Middle Pleistocene (e.g., [10]). The large 1908 Messina earthquake (Mw = 7.1 [2]), for which extensional focal mechanisms were obtained, occurred within this tectonic frame. The December 28, 1908 Messina Earthquake has been ranked as one of the most destructive events of the last centuries, and costed the highest toll in
human life in Italyâ&#x20AC;&#x2122;s history of seismicity, with over 80,000 people dieing in the cities of Messina and Reggio Calabria and the surrounding area. Besides buildings collapse and fires, the damages produced by ground shaking were aggravated by the effects of a remarkable tsunami, with up to 11 m of run-up height, that followed the earthquake [12]. Despite such catastrophic effects the location of the causative fault is not fully assessed (e.g., [13]). Inverse modelling of seismograms and geodetic levelling, and geological studies have produced a variety of results in terms of position, direction, length and dip of the fault, with the most recent solutions proposing
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Figure 3: Map with traces of multichannel seismic profiles in the study area. The trace of the supposed Taormina Fault is in red. Onshore faults after Ghisetti [3]. Seismic profiles shown in Figures 4-6 are in thick green lines. long, E-dipping faults trending about N-S (e.g., [14, 2], and references therein Figure 2). At present, the most accepted seismogenic source for the 1908 earthquake is a 40 km-long, blind fault dipping 30째 to the ESE which is thought to accounts for the topography of the Messina Straits (DISS, http://diss.rm.ingv.it). This fault plane, which trends NNE-SSW, has a minimum depth of 3 km, and would crop out along the coast of Sicily (Figure 2). As for the seismogenic fault, the source of the tsunami related to the earthquake is still 848
a matter of debate [15], and contribution from a so far unidentified submarine slide is called upon. In spite of the hazard potential of this marine region, geophysical surveys purposely devised to investigate the neotectonic features are lacking, with the notable exception of the early work of Selli [16] which, however, did not employ modern geophysical techniques. In order to bridge this gap a multichannel seismic survey, aimed at defining the structural pattern of the Messina Straits, was carried out within the frame of INGV-DPC seismolog-
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Figure 4: Seismic profile TAO 09 across the central part of the supposed Taormina fault. Arrows mark the South Calabrian fault. See Figure 3 for location. ical projects (Figure 3 [13]).
2
Discussion
rine strata along the offshore slope (Figure 4), but such deformation cannot be related to footwall uplift of a normal fault. The geodynamic implication is that extension in south-eastern Sicily, on the Ionian side of the Hyblean Plateau, and extension in southern Calabria and Messina Straits, belong to two different tectonic systems and cannot be mechanically linked. The seismological implication is that the lack of earthquakes is not indicating the occurrence of a seismicity gap.
Three key aspects related to the tectonics of the Messina Straits will be briefly discussed below, based on the results obtained from multichannel seismic data. A more detailed presentation of the data and interpretation procedures is given elsewhere [13], with references therein), together with the correlation of seismic units with onshore outcrops and short offshore wells 2.2 that allows to date the seismic units.
2.1
The elusive Taormina Fault
The hypothesized occurrence of a large extensional fault parallel to the coastline and located offshore, between Taormina and Briga (Taormina Fault) can be rejected by seismic data [13]. That stretch of coastline is actively deforming, as suggested by uplifted marine terraces and tilting of ma-
The faults of the Messina Straits and the 1908 Earthquake
Within the northern part of the Messina Straits, the imaged faults are located on the Calabrian side and dip to the west. These faults appear connected to the fault system reported onshore near Reggio Calabria (Figure 3 [3]). A fault trending NW-SE has been mapped between the towns of Reggio Calabria and Messina. The fault 849
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Figure 5: Seismic profile TAO 08 within the Messina Straits. A listric fault (arrows) is flattening at shallow depth. Location in Figure 5. plane dips to the west and displays a listric shape, with growth strata in the lower part of the half graben sedimentary fill, that can be dated as Late Pliocene-Early Pleistocene (Figure 5). Towards the coast of Calabria the fault is sealed by a sedimentary wedge, which can be dated as Middle Pleistocene-Holocene. The profiles approaching SW Calabria show a 30 kmlong west-dipping fault that is affecting the sea floor with a NW-SE trend, which represents the longest lineament observed within the Messina Straits (Figure 6). The W-dipping southern Calabrian fault has a low-angle plane when depth-migrated therefore, assuming that the epicentre is located above the deeper, seismogenic portion of the fault plane, it does not satisfy the 1908 macroseismic intensity observed along the Calabrian coast. However, this fault might be responsibile for the March 28, 1780 and May 22, 1932 earthquakes (Figure 7), two events for which a seismogenic fault is difficult to be identified [1]. The lack of evident extensional faults
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within the narrower part of the Messina Straits might support the interpretation of a seismogenic fault located to the south of this area or, alternatively, of a blind fault located within the northern Messina Straits. Hints suggesting that the occurrence of a blind fault, possibly east-dipping as indicated by seismological studies, are not obvious on seismic profiles in the northern part of the straits. Moreover, given the large magnitude of the 1908 event, it cannot be ruled out that more than one fault were activated at the same time, as already proposed (Figure 2). In this event, the long fault observed offshore southern Calabria could be activated together with a blind east-dipping fault located further to the north, possibly contributing to the tsunami related to the 1908 earthquake the tsunami waves observed all along the east coast of Sicily, in fact, are hardly compatible with a blind fault located within the northern part of the Messina Straits [15].
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Figure 6: Seismic profile TAO 17. Arrows mark the South Calabrian fault. See Figure 3 for location.
2.3
The 1908 Messina Earth- phobathymetry and seismic profiles clearly quake and its related show that the morphology of the area is the product of long-lasting erosion without any Tsunami trace of a 100 yr-old large-scale landslide.
Because of the uncertainty in the location of the causative fault, the modelling of the associated tsunami is not satisfactory [15], and the possibility to have a contribution from a submarine slide has been explored [19, 20, 21]. A. Billi et al. [22] have recently proposed that the tsunami that stroke the coast of the Messina Strait in December 1908 originated from a very large submarine landslide (20 km3 ) located offshore Giardini-Naxos. However, geophysical data and work on tsunami modelling cast doubt on the proposed hypothesis that a large submarine landslide that was caused by the 1908 earthquake is located offshore Giardini-Naxos [23]. The morphological expression of a 100-yr old landslide, of its headscarp and its deposit, should be little affected by subsequent erosion and reworking. High-resolution mor-
3
Conclusions
The data collected during the Taormina2006 survey open some interesting questions concerning the position, direction and extent of active faults within the Messina Straits, where it is difficult to find a single fault that is long enough to account for the Mw 7.1, 1908 Messina earthquake. Moreover, the trend of the observed faults, though consistent with faults onshore Calabria, is different from the trend of many of the faults proposed on the basis of inversion of seismological and geodetic data (Figure 2). Seismic data show that there is no N-S fault cutting the sedimentary successions south of the epicenter of Schick, 1977 (Figure 7). In fact, the only long 851
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Figure 7: Summary map with main faults in the study area. For the 1908 Messina earthquake the followings are indicated: a) the relocation of [17] (ellipse with green star), b) the epicenter of [18] (yellow star), c) the isoseismal X [12], and the breaks in the telegraph cables. The DISS fault plane is indicated in green. fault that is affecting the sea floor is located at the SW tip of Calabria, trends NWSE and dips to the west with a low angle. Such fault parameters are not compatible with the 1908 earthquake, as the hypocentral depth would be located too far to the west. However, this fault could have contributed to the tsunami wave that came together with the 1908 earthquake better than any faults located within the northern part of the Messina Straits. As a working hypothesis, it seems that a system of interconnected faults (e.g.,[3]), partly exploiting preexisting fault planes, can better describe the observed geological (i.e.,
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long term) deformation. This interpretation leaves the possibility to have more than a fault active at the same time an event that has been inferred for the 1908 earthquake (e.g., Figure 2). At present it is difficult to say whether the observed active faults are just the surface expression of a single, deeper and blind seismogenic fault, as suggested by seismological arguments (e.g., DISS), or if they truly represent the complex response of an area that suffered a great deal of geological deformation and where more than one fault can be active at the same time. The large magnitude of the 1908 earthquake makes the first hypothesis
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perhaps more appealing. As far as the 1908 tsunami is concerned, an alternative to a fault that extends to the south of the Straits consists in introducing a submarine slide that adds to the earthquake in contributing to the tsunami. In this regards, the recent proposal that the 1908 tsunami was due to a very large landslide of about 20 km3 , the deposit of which would be located offshore Giardini-Naxos [22], is not supported by evidence [23]. Moreover, modelling of the 1908 tsunami indicates that a much smaller slide, less than 2 km3 , is enough to produce the effects observed along the shores [20, 21]. The location of this potentially
tsunamigenic slide, however, is difficult to constrain from the inversion of poor quality tsunami runup observations, and no landslide deposit that can be related to the 1908 earthquake has been so far identified on swath bathymetry. The 1908 earthquake has been a complex event. At present, the combined action of a blind fault and a submarine slide is thought to be able to account for the observed ground shaking and tsunami. However, given the structural complexity of the area, other solutions, where more than one slide, and perhaps more than one fault, are active at the same time, cannot a priori be disregarded.
References [1] R. Azzaro, F. Bernardini, R. Camassi, and V. Castelli. The 1780 seismic sequence in NE Sicily (Italy): shifting an underestimated and mislocated earthquake to a seismically low rate zone. Natural Hazards, 42:149–167, 2007. [2] N.A. Pino, A. Piatanesi, G. Valensise, and E. Boschi. The 28 December 1908 Messina Straits Earthquake (Mw 7.1): A Great Earthquake throughout a Century of Seismology. Seismological Research Letters, 80:243–259, 2009. [3] F. Ghisetti. Fault parameters in the Messina Strait (southern Italy) and relations with the seismogenic source. Tectonophysics, 210:117–133, 1992. [4] G. Valensise and D. Pantosti. A 125 Kyr-long geological record of seismic source repeatability: the Messina Straits (southern Italy) and the 1908 earthquake (Ms 7 1/2). Terra Nova, 4:472–483, 1992. [5] F. Antonioli, L. Ferranti, K. Lambeck, S. Kershaw, V. Verrubi, and G. Dai Pra. Late Pleistocene to Holocene record of changing uplift rates in southern Calabria and northeastern Sicily (southern Italy, Central Mediterranean Sea). Tectonophysics, 422:23–40, 2006. [6] S. Catalano and G. De Guidi. Late Quaternary uplift of northeatern Sicily: relation with the active normal faulting deformation. J. Geodynamics, 36:445–467, 2003. [7] C. Monaco and L. Tortorici. Active faulting in the Calabrian arc and eastern Sicily. J. Geodynamics, 29:407–424, 2000.
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[8] A. Argnani and C. Bonazzi. The Malta Escarpment fault zone offshore eastern Sicily: Pliocene-Quaternary tectonic evolution based on new multichannel seismic data. Tectonics, 24(TC4009):1–12, 2005. [9] G. Neri, G. Barberi, G. Oliva, B. Orecchio, and D. Presti. A Possible Seismic Gap within a Highly Seismogenic Belt Crossing Calabria and Eastern Sicily. Bull. Seismological Society of America, 96:1321–1331, 2006. [10] B. Dumas, P. Gueremy, and J. Raffy. Evidence for sea-level oscillations by the “characteristic thickness” of marine deposits from raised terraces of Southern Calabria (Italy). Quaternary Sci. review, 244:2120–2136, 2005. [11] E. Serpelloni, G. Vannucci, S. Pondrelli, A. Argnani, G. Casula, M. Anzidei, P. Baldi, and P. Gasperini. Kinematics of the Western Africa-Eurasia plate boundary from focal mechanisms and GPS data. Geoph. J. International, 169:1180–1200, 2007. [12] M. Baratta. La catastrofe sismica calabro-messinese. 28 Dicembre 1908. pages 1–426, 1910. [13] A. Argnani, G. Brancolini, C. Bonazzi, M. Rovere, F. Accaino, F. Zgur, , and E. Lodolo. The results of the Taormina 2006 seismic survey: Possible implications for active tectonics in the Messina Straits. Tectonophysics, 476:159–169, 2009. [14] A. Amoruso, L. Crescentini, and R. Scarpa. Source parameters of the 1908 Messina Straits, Italy, earthquake from geodetic and seismic data. J. Geoph. Res., 107(B4):1–11, 2002. [15] E. Boschi, D. Pantosti, and G. Valensise. Modello di sorgente per il terremoto di Messina del 1908. Atti VIII Convegno GNGTS, pages 246–258, 1989. [16] R. Selli. Geologia e simotettonica dello Stretto di Messina. Nazionale dei Lincei, 43:119–154, 1978.
Atti Accademia
[17] A. Michelini, A.J. Lomax, A. Bono, A. Nardi, B. Palombo, A. Rossi, and the INGV SISMOS Group. Relocation of instrumentally recorded, historical earthqukes in the Italian region. Geophysical Research Abstracts, 6(07642), 2004. [18] R. Schick. Eine seismotektonische Bearbeitung des Erdbebens von Messina im Jahre 1908. Geol. Jahrb., 11:3–74, 1977. [19] A. Piatanesi, S. Lorito, , and F. Romano. Il Grande maremoto del 1908: analisi e modellazione. 2008. [20] S. Tinti, A. Armigliato, F. Zaniboni, R. Tonini, G. Pagnoni, S. Gallazzi, A. Manucci, and P. Pontrelli. Quale sorgente per il maremoto del 28 Dicembre 1908 nello Stretto di Messina? Terremoto, frana sottomarina o entrambe? GNGTS, 27° Convegno Nazionale, Extended Abstract, pages 191–192, 2008. 854
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[21] M. Favalli, E. Boschi, F. Mazzarini, and M.T. Pareschi. Seismic and landslide source of the 1908 Straits of Messina tsunami (Sicily, Italy). Geoph. Res. Lett, 36(L16304):1–6, 2009. [22] A. Billi, R. Funiciello, L. Minelli, C. Faccenna, G. Neri, B. Orecchio, , and D. Presti. On the cause of the 1908 Messina tsunami, southern Italy. Geophys. Res. Lett., 35(L06301):1–5, 2008. [23] A. Argnani, F.L. Chiocci, S. Tinti, A. Bosman, M.V. Lodi, G. Pagnoni, F., and Zaniboni. Comment on “On the cause of the 1908 Messina tsunami, southern Italy” by Andrea Billi et al. Geoph. Res. Lett., 36(L13307):1–2, 2009.
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The Salerno Valley (Campania Continental Margin, Southern Italy): Highlights into the Stratigraphic-Structural Setting and the MorphoBathymetry of a Pleistocene Half-Graben G. Aiello, E. Marsella, V. Di Fiore Institute for Coastal Marine Environment, CNR, Napoli, Italy gemma.aiello@iamc.cnr.it Abstract New results on the seismic stratigraphy and morpho-bathymetry of the Pleistocene sedimentary basin of the Salerno Valley focused on the regional seismic stratigraphy of the Southern Campania passive continental margin (Southern Tyrrhenian sea, Italy) are presented here. Original data include Multibeam bathymetry as well as multichannel and single-channel seismic profiles, recently collected onboard the R/V Urania of the National Research Council of Italy (oceanographic cruise SISTERII). The Salerno Valley represents a half-graben sedimentary basin, whose identification has been controlled, during the Early Pleistocene, by the master fault Capri-Sorrento Peninsula, showing vertical throws of 1500 metres, which downthrows the Meso-Cenozoic carbonatic acoustic basement under the sedimentary basin. The geologic interpretation of multichannel seismic profiles has enabled the identification of a main unconformity, located at depths ranging from 3000 to 3500 metres under the sea bottom and correlated to the top of the Meso-Cenozoic carbonatic sequence, cropping out onshore in the Sorrento Peninsula structural high. This unconformity bounds upwards the carbonatic acoustic basement, strongly deformed by normal faulting, and represents the base of the Pleistocene basin filling of the Salerno Valley. The basin filling, with an overall thickness exceeding 1000 metres, is characterized by parallel and continuous seismic reflectors alternating with chaotic intervals (gravity mass deposits).
1
Introduction
New multichannel and single-channel seismic profiles and morpho-bathymetric data, consisting of high resolution Multibeam bathymetry of the Salerno Valley have been analysed in order to investigate the tectono-sedimentary evolution of this sector of the Southern Tyrrhenian margin during the Pleistocene and to provide a better understanding of the extensional basins
offshore Southern Italy. In a similar way to other back-arc basins, the Tyrrhenian sea is an area of ongoing extension inside large-scale convergence zones between the continental plates of Europe and Africa. Tyrrhenian extension started about 10 Myr ago, leading to the Pliocene formation of oceanic crust [1, 2, 3]. A deep and narrow Benioff zone, plunging from the Ionian sea towards the Southern Tyrrhenian Basin testifies the occurrence of an east-
Marine Geology
ward migrating subduction plan of the eastern Mediterranean lithosphere [4]. From the Oligocene to the Recent, the subduction has generated the Western Mediterranean and the Tyrrhenian back-arc basins, as well as the accretionary wedge constituting the Southern Apenninic Arc. The extension in the Tyrrhenian sea started in the Late Miocene and produced the Vavilov Plain during the Pliocene and the Marsili Plain during the Quaternary [5, 6, 7, 8, 4, 9]. In the last years, the acquisition of large-scale morpho-bathymetric data on the submerged portions of volcanic edifices and, as a general rule, in the oceanic basins and in the seas, have furnished interesting data, giving a new impulse to the knowledge of the structure and the geological evolution of the oceans. Relatively to the Italian seas, the recent realization of Multibeam bathymetric data in the Tyrrhenian sea [10] has furnished important geological data, collecting new morpho-bathymetric hints on volcanic, tectonic and gravitational processes at several scales. The aim of this paper is to highlight new morpho-bathymetric and seismostratigraphic evidences on the Salerno Valley based on the integrated geological interpretation of Multibeam bathymetry and multichannel seismics. The obtained results are here framed in the regional geological context of the Campania continental margin (see the section 3) in order to improve the geological knowledge on the tectono-stratigraphic setting of the extensional basins offshore Southern Italy. A sketch map is reported (Figure 1) showing the location of the study area and the simplified geological setting of the Campania region. The Salerno Valley is a Pleistocene half-graben basin, whose identification and tectonic setting have been controlled by the master fault Capri-Sorrento 858
Peninsula (showing vertical throws in the order of 1500 meters) and is filled by marine clastic and epicontinental sediments. The acoustic basement is composed of Mesozoic carbonate platform units (penetrated by the lithostratigraphic wells “Mina 1”, “Milena 1” and “Margherita mare 1” [11]) and the overlying Miocene siliciclastic units related to the “Liguride Units” and the corresponding foredeep deposits (“Flysch del Cilento”), cropping out on land in the Cilento Promontory [12, 13, 14]. Strong seismic reflectors inclined towards SSE with apparent angles of 10°-15° appear on seismic profiles, indicating the occurrence of detachment levels in the carbonatic multilayer [15, 16]. The extension in the Salerno basin is controlled by listric normal faults, converging on low-angle detachment levels; one of them is located at the top of the Miocene flysch terrains [17]. Bathymetric and high resolution reflection seismic data previously collected in the Salerno Valley [15, 16] indicate that the depositional processes and overfilling of the basin prevail on the continental shelf to the east, while the western sector of the halfgraben is the site of erosional and sedimentary processes still active in the Salerno canyon, showing depths ranging from 600 to 1000 m. A strong synsedimentary tectonics and Pleistocenic uplift of the adjacent sectors of the Southern Apennines are suggested by wedging geometries, tectonic unconformities and “hummocky” facies at several stratigraphic intervals in the Pleistocenic succession. Roll-over mechanisms, or alternatively, a tectonic inversion of the half-graben, are suggested by anticlines with ENE-WSW axes, deforming the Pleistocene sedimentary sequences. This overall tectonic framework suggests a recent age of the extensional processes in the Salerno basin and is responsible for the
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Figure 1: Sketch map showing the location of the study area and the geological setting (in the lower part of the figure) of the Campania Region. The regional seismic stratigraphy of the Campania continental margin has been investigated based on the geological interpretation of the seismic profiles NAM3 (Gulf of Naples) and SAM5, SAM7a and SAM4 (Gulf of Salerno). Key. 1: Meso-Cenozoic carbonates. 2: Siliciclastic flysch deposits. 3: Quaternary volcanic deposits. 4: Quaternary alluvial, coastal and marine deposits. present-day topography of the sea bottom carried out at water depths ranging beand the foreland uplift along the rift shoul- tween 100 m and 1000 m using the RESON ders (Sorrento Peninsula). Seabat 8160 deep-towed Multibeam. The acquisition of Multibeam data has been achieved through a Multibeam coverage of 2 Data acquisition and 50Multichannel seismic profiles have been recorded using two G/I Airguns seismic processing sources by SSI Sodera. The acquisition system is represented by the Geometrics The presented data have been collected Stratavisor 24 bit seismograph with a 48during the oceanographic cruise SISTERII channel. The sample interval was set at 0(December 2004) onboard the R/V Ura5 sec, the windows time at 5 sec, the renia of the National Research Council of ceive interval at 6.25 m and the shot inItaly (Scientific Responsible: Dott. E. terval at 25 m. An advanced seismic data Marsella). The acquisition of data has been 859
Marine Geology
Figure 2: Total coverage map of the Multibeam acquisition during the oceanographic cruise SISTERII. The visualization of the bathymetric data has been realised through a colorimetric scale and using the software PDS2000 (RESON navigation and Multibeam acquisition). processing was aimed at the reduction of the random noise in the data, including the removal of unwanted coherent events and reduction of spatial aliasing. The combination of pre-stack DMO and post-stack data allowed a better localization of the reflectors on the seismic sections. Cycles of velocity analysis and residual static corrections improved the quality of the velocity function and therefore, the NMO correction. Pre-stack spiking deconvolution widened the frequency spectra of the signal and boosted the data resolution. FK filtering on NMO corrected data, CDP gathers and pre and post-stack predictive deconvolution weakened multiple reflections.The geologic interpretation has been carried out on the stacked seismic sections based on the criteria of seismic stratigraphy [9, 20].
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3
Regional seismic stratigraphy
The Campania continental margin owes its complex stratigraphic architecture to the interaction between volcanic and sedimentary processes occurred during the Late Quaternary. Multichannel profiles, coupled with high resolution reflection seismics, Multibeam bathymetry and magnetic profiles, have contributed to the knowledge of the tectono-stratigraphic setting of the margin, both on a regional and on a local scale [15, 16, 18, 21, 22, 19, 23, 24, 25]. The regional seismic stratigraphy of the continental margin is summarized by several studies carried out at the IAMC-CNR based on the interpretation of the seismic data previously acquired in the area (more than 10.000 km of high-resolution singlechannel sparker and air-gun lines). The examination of the Neogene depositional
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Figure 3: Line drawing of the multichannel profile NAM3 (slightly modified after Aiello et al., [18] and Dâ&#x20AC;&#x2122;Argenio et al.,[19]) showing the stratigraphic architecture of the Campania continental margin in the Gulf of Naples. The profile runs along a NW-SE transect from offshore, towards the Sorrento Peninsula. The acoustic basement, represented by Meso-Cenozoic carbonates, is strongly downthrown by normal faulting towards the centre of the bay. The basin filling is composed of several seismic sequences; two of them (units 2 and 4) are represented by wide relic prograding wedges and appear to be the most relevant in the tectono-sedimentary evolution of the area. Key. 1: Acoustic basement (Meso-Cenozoic platform carbonates). 2: Early Pleistocene relic prograding wedge. 3: Wedge-shaped, transgressive unit, mainly developed in slope and basin settings, probably composed of siliciclastic deposits. 4: Middle-Late Pleistocene prograding wedge. 5b: Buried volcanic complexes related to Ischia and Procida eruptions, ranging in age from 55 to 18 ky. 6: Late Quaternary marine and coastal deposits: prograding wedges deposited on the continental shelf; drapes filling intra-platform basins. sequences and unconformities has proven to be constructive, providing some constraints on the tectono-stratigraphic models, mainly in terms of directions and timing of extensional processes on the Campania continental margin. The seismostratigraphic setting of the area is shown by the seismic profile NAM3 (Figure 3 [18, 19]). Several seismic sequences separated by unconformities have been distinguished. The acoustic basement is represented by Meso-Cenozoic platform carbonates (unit 1 in Figure 3), cropping out onshore in the Sorrento Peninsula and organized as a NWdipping monoclinalic structure. The basin filling consists of two prograding wedges (units 2 and 4 in Figure 4), each 20th dis-
tinctive acoustic patterns and seismic facies. They appear to be the most prominent units in the stratigraphic architecture of the area. The oldest one (unit 2) is interpreted as a wide relic prograding wedge, north-westwards dipping, formed by siliciclastic deposits, probably Pleistocene in age and occurring offshore the Sorrento Peninsula. Both the Meso-Cenozoic carbonates and the unit 2 have been probably involved in a tectonic tilting during Pleistocene extensional phases, slightly increasing the steepness of seismic reflectors. On the continental shelf, the seismic reflectors are truncated by a main subaerial unconformity (unconformity B in Figure 3), whose areal extension varies from 2-3 kilo-
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Figure 4: Line drawing of the multichannel line SAM7, showing the stratigraphic architecture of the Campania continental margin in the Gulf of Salerno. The profile runs from NNW to SSE between the Salerno offshore and the Cilento offshore and shows coupled conjugate listric normal faults, related to a master fault (Capo Dâ&#x20AC;&#x2122;Orso fault), NE-SW oriented, strongly downthrowing the Pleistocene basin filling. Key: 1: Early Pleistocene marine deposits, strongly deformed by normal faulting and representing the first seismic unit of the basin filling. 2: Syntectonic sequence, deposited in the central part of the basin, probably during the tectonic activity of the Capo Dâ&#x20AC;&#x2122;Orso fault (Middle Pleistocene?). 3: Late Pleistocene marine deposits, relatively undeformed and representing the late stages of the basin filling. metres to several hundred of meters. The unconformity B indicates a main relative sea level fall and a strong basinwards shifting of coastal and marine facies, accompanied by sedimentary bypass and strong erosion on the shelf and slope. Above the unconformity B, the clinoforms of unit 3 (Figure 3) progressively onlap the slope and basin areas up to the continental shelf. Unit 3 represents a wedge-shaped, transgressive unit, mainly developed in slope and basin settings and probably composed of siliciclastic deposits. Above the unit 3, a wide prograding wedge (unit 4 in Figure 3) develops. It shows well preserved offlap breaks and thickens from the shelf towards the slope. The wedge grades laterally, or alternatively is overlain by the
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volcanic unit 5 and is deeply incised by the Dohrn canyon axes. Above the unit 4 and/or in facies hetheropy with the latter, a wedge-shaped seismic unit. acousticallytransparent and volcanic in origin has been clearly identified offshore the Sorrento Peninsula, where it unconformably overlies the Meso-Cenozoic carbonates and/or the seismic unit 2. Prograding wedges deposited on the continental shelf and/or on the flanks of volcanic edifices and sediment drapes infilling local basinal depressions are interpreted as Late Quaternary marine and coastal deposits. Holocene highstand drape, covering the whole margin, is represented by wedge-shaped or drapeshaped units represented by highly continuous parallel and subparallel reflectors.
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Figure 5: Shaded-relief map of the Multibeam bathymetry of the Salerno Valley, covering the southern continental slope bounding the structural high of the Sorrento PeninsulaCapri island and the northern sector of the Gulf of Salerno. Note the occurrence of a dense network of erosional channels deeply eroding the continental slope south of Sorrento Peninsula. The south-western rim of the Salerno Valley is bounded by two complex, NE-SW trending, morpho-structural highs; on their western flank a wide intra-slope basin occurs, located at a water depth of 700 m. The geologic interpretation of the seismic profile SAM7a (Salerno-Cilento Figure 4) has shown coupled conjugate listric normal faults, dipping towards SSE and NNW. These faults, showing strong downthrows of the clastic multilayer, converge at several stratigraphic levels and appear to be related to a master fault, NE-SW oriented (Capo Dâ&#x20AC;&#x2122;Orso fault). Three main seismic sequences have been recognized in the basin filling. The Early Pleistocene marine deposits, strongly deformed by normal faulting (unit 1 in Figure 4) are overlain by a syntectonic sequence (unit 2 in Figure 4), probably deposited during the activity of the Capo Dâ&#x20AC;&#x2122;Orso fault. A main unconformity fossilizes the seismic sequences involved by extensional tectonics in correspondence to normal faults. It marks the beginning of the deposition of relatively
undeformed seismic sequences, forming a thick prograding wedge genetically related to the Sele river (unit 3 in Figure 4). Further evidence emerges from the interpretation of the seismic profile SAM4 (Figure 5), running from the Salerno offshore to the structural high of the Salerno continental shelf and giving evidence for the tectonic setting of the clastic multilayer, characterized by several normal faults.
4
Multibeam bathymetry
morpho-
The morpho-bathymetry of sedimentary basins based on Multibeam surveys represents a research line of increasing interest, mainly for its implications in the 863
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Figure 6: Morphological sketch map of the Salerno Valley superimposed on contour map of the isobaths. The main morphological lineaments have been recognized through the geological interpretation, i.e. the toe of the continental slope, the submarine slides, the erosional gullies, the canyons and/or channels, the breaks in slope, the morpho-structural highs, the intra-slope basins, the depositional areas and the normal faults. In the inset the shaded relief map of the area is reported. coastal and deep sea environmental monitoring in terms of the definition of geologic and environmental hazards. The continental slopes and submarine shelves surrounding the Campania Region are 20th different morphologies, varying depending on the geologic setting of the adjacent mainland, which controls the occurrence of submarine instabilities. The knowledge of these geological processes is subordinated to the use of up-to-date geophysical survey techniques, as the high resolution multichannel and single-channel seismics, the Subbottom Chirp profilers, the magnetic survey and the single-beam and Multibeam bathymetry. Geophysical and geological data have been previously collected to improve the knowledge of submarine instabilities offshore the Campania Region. This allowed for a discrimination between “slow” submarine insta864
bilities (creeping, slumping, deep gravitational deformations) and “fast” submarine instabilities (debris avalanches, rock falls, roto-translational slidings). Their chronostratigraphic framework is finalized to the correlation with the triggering geological processes at a regional scale (seismicity, volcanism, tectonic activity in correspondence to significant faults). The Multibeam data collected during the SISTERII oceanographic cruise have evidenced the high gradients of the continental slope and the occurrence of erosional processes, partly still active, mainly at the toe of the slope, where a dense network of channels occurs. The analysis of seismic profiles has already evidenced the tectonic activity of the fault escarpment as a triggering factor for gravity instabilities active during Late Pleistocene and Holocene; this is evidenced also by slumping deposits, char-
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Figure 7: (a) Localization on detailed DTM of the bathymetric profiles used for the volumetric computation of submarine landslides. (b) Bathymetric profiles on the slide scars. (c) Detailed Digital Terrain Model showing channels and canyons eroding the lower continental slope. acterized by a chaotic seismic facies, intercalated at several stratigraphic levels in the basin filling. On the contrary, in the distal zone of the valley, where deposition prevails, the shallower sector appears in a regime of intense erosion (Salerno canyon) and shows recent tectonic deformations, as antiforms and normal faults, which can be interpreted in the regional framework as hints of intense extensional tectonics, lasting up to recent times [15, 16]. The Multi-
beam bathymetric data have enabled the construction of a 3D DTM reconstruction, covering an area of 1600 square kilometres (Figure 5). The mapped offshore is bounded from the southern Capri offshore to the north-west and from the Sele river mouth to the south-east. The continental shelf of the Salerno Gulf is strongly asymmetrical, depending on the structural domains of this sector of the Apenninic chain. The variability of extension, of the depth
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Marine Geology
Figure 8: Digital Terrain Model of the Salerno Valley and superimposed location of multichannel seismic profiles recorded in the study area. of the shelf break and of the average slope steepness are controlled by the structural and geological setting of the margin areas [26], more than by the glacio-eustatic variation during the Pleistocene. The continental slope surrounding the Salerno Gulf is characterized by structural depocentres with a NW-SE (Apenninic) trending, originated by the extensional phases related to the Tyrrhenian basin, alternating with morpho-structural highs, adjacent to intra-slope basins. Two NESW (counter-Apenninic) trending morphostructural highs have been identified based on the DTM interpretation (Figure 5). A sketch morphological map of the Salerno Valley superimposed on contour map of the isobaths is reported in Figure 6. The main morphological lineaments have been recognized through the geological interpretation, i.e. the toe of the continen866
tal slope, the submarine slides, the erosional gullies, the canyons and/or channels, the breaks in slope, the morpho-structural highs, the intra-slope basins, the depositional areas and the normal faults. The slope is characterized, throughout its extension, by a dense network of gullies, partly reflecting the hydrographic pattern of the corresponding emerged sector, separated by saddles, with variable dimensions. Important submarine slides have been identified on the continental slope to the south of the Salerno Valley. Main breaks in slope are located in correspondence to the - 240 m, - 350 m and 480 m isobaths. Two complex morphostructural highs (slope ridges), NNW-SSE trending, are located on the south-western continental slope. They are separated by an intra-slope basin, having several depocenters (Figures 5 and 6). Minor depositional
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Figure 9: Multichannel seismic profile L6, recorded in the Salerno Valley during the cruise SISTERII. zones occupy several areas of the continental slope. Several, previously unknown channels have been identified on the continental slope offshore the Sele Plain. They are drafted by the three-dimensional reconstruction of Figure 7, representing a detailed DTM of the area. The channels feed the surrounding bathyal plain, starting from - 200 m of water depth. Several drainage axes join into three main channels, reaching the slope rim, located at 440 m of water depth. On the slope, the channels become two canyons. The eastern one (Figure 7) shows pronounced rims and starts, with high gradients, from a water depth of 520 m reaching the 840 m of water depth. The second canyon, located 2.5 km to the south, is located in the same bathymetric interval and shows at its base a related deposit, having a diameter of more than 2500 m and a maximum height of more than 30 m.
5
Seismic stratigraphy
The multichannel seismic grid, superimposed on shaded-relief map of Multibeam bathymetry is reported in Figure 8. The seismic profile L6 shows a E-W trending (in its first part), passing to a NNW-SSE trending (in its second part). It crosses the Salerno Valley at water depths ranging from â&#x20AC;&#x201C; 900 m to â&#x20AC;&#x201C; 500 m (Figures 9 and 10). The seismic profile shows a first part depicting the basin (i.e. the Salerno Valley) at water depths of about - 900 m. Then it rises in correspondence to a break in slope located at about 825 m up to 525 m along the flank of the slope surrounding the western sector of the Salerno Gulf (Figure 8). The geological interpretation of the multichannel profile L6, carried out based on seismo-stratigraphic criteria (Figure 10) has enabled the identification of several seismic horizons, correspond867
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ing to significant unconformities. In particular, an important unconformity, located between 1.7 and 1.8 sec of depth (twoway travel times; corresponding to about 3000-3500 meters) has been correlated to the top of the carbonatic acoustic basement, widely cropping out onshore in the Sorrento Peninsula-Capri island structural high. This unconformity is strongly downthrown by normal faulting and marks the top of the Meso-Cenozoic acoustic basement and the base of the Pleistocene basin filling of the Salerno Valley. The corresponding seismic reflector lacks in lateral continuity and grades towards parallel seismic reflectors, which form a thick sedimentary package along the slope (Figures 9 and 10). The Pleistocene basin filling of the Salerno Valley is organized in eight seismic sequences (3 to 10 in Figure 10), separated by unconformities (B to I in Figure 10), for an overall thickness of about 600 msec (corresponding to about 510 m by using a seismic velocity of about 1700 m/sec for qualitative time to depth conversion). The seismic sequence 2 (Figure 10) represents the oldest sedimentary drape, overlying the continental slope bounding the eastern sector of the Salerno Gulf. It is 20th parallel reflectors, with high amplitude and continuity. Its top corresponds to an unconformity (B reflector), identified by the onlap of the overlying seismic sequence (sequence 3 in Figure 10). The sequence 3 corresponds to the earliest phases of the basin filling of the Salerno Valley. It is distinguished by onlap terminations on the B unconformity and wedge-shaped external geometry overlying the A unconformity. Its upper boundary corresponds to the C unconformity (Figure 10). The seismic sequences 4, 5 and 6 show seismic horizons with high amplitude and lateral continuity. Their overall geometry corre868
sponds to a vertically-aggrading infill, both in the basin and in the continental slope. The F unconformity (Figure 10), located at the top of the seismic sequence 6, indicates a main variation in the aggradational geometries of the basin filling and a variation in the depositional conditions of the sedimentary basin, accompanied by submarine erosion. This is also suggested by the onlap of the sequence 7 on the unconformity F (Figure 10). The sequence 7 is 20th parallel and continuous seismic reflectors, alternating with transparent intervals. The vertical aggradation of the sequences in the basin continues in correspondence to the seismic sequences 8, 9 and 10 (Figure 10). The blanking of the acoustic signal in correspondence to localized seismic intervals indicates the occurrence of chaotic intervals, probably corresponding to resedimented deposits. They are triggered by submarine instabilities occurring in the sedimentary filling of the Salerno Valley. The seismic profile L4 runs on the continental shelf of the Salerno Gulf at water depths ranging from - 150 m to 250 m with a NNW-SSE trending, parallel to the shoreline of the Salerno Gulf (Figure 11). The seismic interpretation has enabled the distinction of several sequences, probably Pleistocene in age, strongly deformed by normal faulting. These sequences highlight a structural framework 20th structural highs and intervening sedimentary basins (Figures 11 and 12). These sequences overlie an acoustic basement correlated to the siliciclastic deposits of the â&#x20AC;&#x153;Flysch del Cilentoâ&#x20AC;? Auct. [14] Figures 11 and 12). A structural high, controlled by a pattern of normal faults bounding the flanks of the structure towards NNW and SSE, has been individuated (Figures 11 and 12). Both its seismic facies and its stratigraphic position suggest that the high is composed of Neo-
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gene siliciclastic terrains correlated to the “Flysch del Cilento ” Auct., widely cropping out in the adjacent emerged sectors of the Southern Apennines [14]. These terrains are 20th alternating turbidite sands and shales and have been intensively deformed by compressional tectonics during the deformational phases of the Apenninic chain and later by extensional tectonics, during the phases of differential neotectonic uplift enabling the formation of horst and graben structures in the whole Apenninic margin [27, 28]. The structural high represents, perhaps, a high of the acoustic basement (FC in Figure 12). The structural high crops out at the sea bottom (between the CDP points 250 and 420), where it is overlain by a drape of recent sediments, characterized by sub-horizontal seismic reflectors for an overall thickness of 70 msec (Figures 11 and 12). It has induced the deformation of the sea bottom itself. The morphology of the sea bottom in the area affected by the occurrence of the high shows a structural control. The pattern of normal faults controlling the deformation of the structural high is 20th a couple of main faults and by some secondary faults having a trending parallel to the main fault. These secondary faults downthrow both the acoustic basement and the first two underlying Pleistocene sequences, delineating the sedimentary basin of the Salerno Gulf (Figures 11 and 12). This downthrowing is evident both towards NNW and SSE, delineating two zones of sedimentary basin. Their filling is 20th almost five seismic sequences, separated by tectonic unconformities, significant to a regional scale and corresponding to main seismic reflectors, identified at the scale of the whole seismic section. The top of the structural high seems to be eroded. It has been affected by probable phases of differential
uplift and erosion, controlling the individuation of a polyciclic erosional surface, at the top of the structure and at the base of the recent sediments (Figures 11 and 12).
6
Conclusions
The Salerno Valley is a Pleistocenic halfgraben, whose identification and tectonic setting have been controlled by the CapriSorrento normal fault (with a probable strike-slip component), showing an average thickness of about 1500 metres [26, 29, 15, 16, 30]. This fault has controlled the origin of the Salerno canyon, where erosional processes and synsedimentary tectonics appear to be still active. The acoustic basement consists of Mesozoic carbonate platform units (drilled by the lithostratigraphic wells “Mina 1”, “Milena 1”, “Margherita Mare 1” [11]) and the overlying Miocene siliciclastic units related to the “Liguride Units” and the related Miocene foredeep deposits (“Flysch del Cilento”), cropping out on land in the Cilento Promontory [12, 13, 14]. The basin filling is composed of marls and marly clays with intercalations of sands and conglomerates and then of marly clays with intercalations of thin sands (Pleistocene). In the Cilento offshore a regional unconformity, probably related to a non-depositional and/or erosional hiatus (Pliocene is completely missing), occurs at the base of the Pleistocene sequences (“Margherita Mare 1” and “Milena 1” exploration wells [11]). The main morphological lineaments have been individuated through the DTM interpretation (Figures 5 and 6). They are represented by the toe of the continental slope, the submarine slides, the erosional gullies, the canyons and/or channels, the breaks in slope, the morpho869
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structural highs, the intra-slope basins, the depositional areas and the normal faults (Figure 6). The slope is characterized, throughout its extension, by a dense network of previously unknown gullies, reflecting the hydrographic pattern of the adjacent emerged sector of the Sorrento Peninsula. Important submarine slides have been identified on the continental slope to the south of the Valley. Two complex morpho-structural highs (“slope ridges” Trincardi and Zitellini [31]), NNWSSE trending, occur on the SW continental slope, separated by an intra-slope basin, having several depocenters. The Salerno Valley shows an example of complex Quaternary filling of a tectonically-controlled sedimentary basin, recording the interactions between the glacio-eustatic sea level fluctuations, the tectonic activity in the source region and the tectonic deformation in the depositional areas, lasting up to recent times. High sedimentary supply, combined with a restrict sediment dispersal, have enabled the deposition of a thick sedimentary succession. The seismic stratigraphic analysis of multichannel profiles has shown that the Pleistocene basin filling is organized in eight seismic sequences (3 to 10 in Figure 10), separated by regional unconformities (B to I in Figure 10), for an overall thickness of about 500 m. The structural setting is controlled by two main listric normal faults, the Capri-Sorrento master fault and the Capo D’Orso master fault. Several high-angle normal faults, parallel to the Capri-Sorrento master fault
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have been developed in correspondence to the shoreline of the Sorrento Peninsula [32, 33]. Synthetic and antithetic normal faults, parallel to the Capo D’Orso fault have been also identified. This tectonic framework is consistent with the overall framework of the Campania continental margin, where NE-SW (Anti-Apenninic), NW-SE (Apenninic) and E-W normal faults are revealed by outcrop and subsurface geological data [34, 35, 33]. NE-SW trending normal faults were the main structural features and gave rise to asymmetrical extensional structures, i.e. half-graben filled with Quaternary deposits and tilted blocks dipping towards the NW [36]. Multibeam data have evidenced that the continental slope surrounding the Salerno Valley is characterized by structural depocentres with a NWSE (Apenninic) trending, originated by the extensional phases related to the Tyrrhenian basins [31, 37]. These depocenters alternate with morpho-structural highs, bounded to their western flank, by a wide intra-slope basin, NE-SW trending. The occurrence of chaotic intervals, interpreted as gravity transport deposits and intercalated at several stratigraphic levels in the sedimentary succession, suggests that the tectonic activity lasted up to the Late Pleistocene in the offshore adjacent to the sedimentary basin, probably in correspondence to NNW-SSE trending normal faults. This evidence, perhaps, is confirmed by the tectonic setting of the morpho-structural highs and related intra-slope basins in the southeastern offshore of the area.
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Figure 10: Geologic interpretation of multichannel seismic profile L6 (see the Figure 9 for the corresponding seismic line). Key: 1) Meso-Cenozoic platform carbonates. 2) Early Pleistocene marine deposits, which can be probably attributed to relic prograding wedges. 3) First seismic sequence of the basin filling, onlapping the B unconformity (located at the top of the seismic sequence 2) and unconformably overlying the unit 1, with parallel and subparallel reflectors. 4) Second seismic sequence of the basin filling, 20th inclined reflectors on the slope and by parallel reflectors in the basin. 5) Third seismic sequence of the basin filling, 20th inclined reflectors on the slope and by parallel reflectors in the basin. 6) Fourth seismic sequence of the basin filling, 20th inclined reflectors on the slope (where it is relatively thinner) and by parallel reflectors in the basin. 7) Fifth seismic sequence of the basin filling, 20th a wedge-shaped external geometry, onlapping the F unconformity (located at the top of the seismic sequence 6). 8) Sixth seismic sequence of the basin filling, 20th inclined reflectors on the slope and parallel reflectors in the basin. 9) Seventh seismic sequence of the basin filling, 20th inclined reflectors on the slope and parallel reflectors in the basin. 10) Eighth seismic sequence of the basin filling, 20th inclined reflectors on the slope and parallel reflectors in the basin.
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Marine Geology
Figure 11: Multichannel seismic profile L4 recorded in the Salerno Gulf parallel to the Tyrrhenian shoreline
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Figure 12: Geologic interpretation of the multichannel profile L4 (see the Figure 11 for the uniterpreted seismic profile). Key. FC: Acoustic basement. Cenozoic siliciclastic deposits related to the â&#x20AC;&#x153;Flysch del Cilentoâ&#x20AC;? Auct. 1 and 2: Seismic units of the Pleistocene basin filling, in lateral contact with the acoustic basement along normal faults, 20th parallel to subparallel seismic reflectors. 3: Seismic unit of the Pleistocene basin filling , in lateral contact with the acoustic basement along normal faults, 20th parallel to subparallel seismic reflectors. 4: Seismic unit of the Pleistocene basin filling, onlapping the flanks of the structural high of the Cenozoic acoustic basement and 20th parallel seismic reflectors and slight thickness variations. 5: Seismic unit of the Pleistocene basin filling, onlapping the flanks of the structural high of the Cenozoic acoustic basement and 20th parallel to subparallel seismic reflectors, slightly deformed by normal faulting. 6: Seismic unit of the Pleistocene basin filling, onlapping the flanks of the structural high of the Cenozoic acoustic basement, infilling intra-basinal depressions with bi-directional onlaps. 7: Late Pleistocene-Holocene seismic unit. 20th low angle progradational reflectors, cropping out at the sea bottom.
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References [1] E. Patacca, R. Sartori, and P. Scandone. Tyrrhenian basin and Apenninic arcs: kinematics relations since late Tortonian times. Memorie della Societ`a Geologica Italiana, 45:425–451, 1990. [2] E. Patacca and P. Scandone. Constraints on the interpretation of the CROP-04 seismic line derived from Plio-Pleistocene foredeep and thrust-sheet-top deposits (Southern Apennines, Italy). Bollettino Societ`a Geologica Italiana (Special Issue), 7:241–256, 2007. [3] E. Patacca and P. Scandone. Geology of the Southern Apennines. Bollettino Societ`a Geologica Italiana (Special Issue), 7:75–119, 2007. [4] R. Sartori. The Tyrrhenian back-arc basin and subduction of the Ionian lithosphere. Episodes, 26(3), 2004. [5] K. Kastens, J. Mascle, C. Auroux, E. Bonatti, C. Broglia, l J.E.T. Channel, P. Curzi, K.C. Emeis, G. Glason, S. Hasegawa, W. Hieke, G. Mascle, F. Mc Coy, J. Mc Kenzie, J. Mendelson, r C. Muelle, J.P. Rehault, A. Robertson, R. Sartori, R. Sprovieri, and M. Torii. Young Tyrrhenian sea evolved very quickly. Geotimes, 31(8):11–14, 1986. [6] K. Kastens, J. Mascle, C. Auroux, and the ODP Leg 107 Scientific Party. ODP Leg 107 in the Tyrrhenian sea: insights into passive margin and back-arc basin evolution. Geol. Soc. Amer. Bulletin, 100:1140–1156, 1988. [7] R. Sartori and R. Capozzi. Patterns of Neogene to Recent rift-related subsidence in the Tyrrhenian domain. Proceedings of the International School Earth and Planetary Sciences, CNR, Siena, pages 147–158, 1998. [8] R. Sartori, G. Carrara, L. Torelli, and N. Zitellini. Neogene evolution of the southwestern Tyrrhenian sea (Sardinia Basin and western bathyal plain). Marine Geology, 175:47–66, 2001. [9] R. Sartori, L. Torelli, N. Zitellini, G. Carrara, M. Magaldi, and P. Mussoni. Crustal features along a W-E Tyrrhenian transect from Sardinia to Campania margins (Central Mediterranean). Tectonophysics, 383:171–192, 2004. [10] M. Marani and F. Gamberi. Distribution and nature of submarine volcanic landforms in the Tyrrhenian sea: the arc vs. the back-arc. Memorie Descrittive della Carta Geologica d’Italia, 64:97–108, 2004. [11] Agip. Temperature sotterranee. Inventory of the data collected by the Agip during the research and the production of hydrocarbons in Italy, pages 1–1390, 1977. [12] B. D’Argenio, T. Pescatore, and P. Scandone. Schema geologico dell’Appennino meridionale (Campania e Lucania). Accademia Nazionale dei Lincei, Atti del Convegno “Moderne Vedute sulla Geologia dell’Appennino”, Roma, 183, 1973. 874
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[13] B. D’Argenio, F. Ortolani, and T. Pescatore. Geology of the Southern Apennines. Geologia Applicata e Idrogeologia, 6:135–161, 1986. [14] G. Bonardi, F.O. Amore, G. Ciampo, P. De Capoa, P. Miconnet, and V. Perrone. Il Complesso Liguride Auct.: stato delle conoscenze e problemi aperti sull’evoluzione pre-appenninica ed i suoi rapporti con l’Arco Calabro. Memorie della Societ`a Geologica Italiana, 41:17–35, 1988. [15] G. Aiello, F. Budillon, G. De Alteriis, L. Ferranti, E. Marsella, and G. Pappone. Late Neogene tectonics and basin evolution of Southern Italy Tyrrhenian margin. Abstract ILP Task Force “Origin of Sedimentary Basins”, Torre Normanna, Palermo (Italy), June 1997, 1997. [16] G. Aiello, F. Budillon, G. De Alteriis, O. Di Razza, M. De Lauro, E. Marsella, N. Pelosi, F. Pepe, M. Sacchi, and R. Tonielli. Seismic exploration of the perityrrhenian basins in the Latium-Campania offshore. Abstract Int. Cong. ILP Task Force “Origin of Sedimentary Basins”, Torre Normanna, Palermo (Italy), June 1997., 1997. [17] M. Sacchi and E. Infuso, S. Marsella. Late Pliocene-Early Pleistocene compressional tectonics in the offshore of Campania. Bollettino di Geofisica Teorica Applicata, 36:141–144, 1994. [18] G. Aiello, F. Budillon, G. Cristofalo, G. De Alteriis, M. De Lauro, L. Ferraro, E. Marsella, N. Pelosi, M. Sacchi, and R. Tonielli. Marine geology and morphobathymetry in the Bay of Naples. In: Faranda F.M., Guglielmo L. and Spezie G. (Eds.) Structures and Processes of the Mediterranean Ecosystems, Springer-Verlag Italy. pages 1–8, 2001. [19] B. D’Argenio, G. Aiello, G. De Alteriis, A. Milia, M. Sacchi, R. Tonielli, F. Budillon, F.L. Chiocci, A. Conforti, M. De Lauro, C. D’Isanto, E. Esposito, L. Ferraro, D. Insinga, M. Iorio, E. Marsella, F. Molisso, V. Morra, S. Passaro, N. Pelosi, and S. Porfido. Digital Elevation Model of the Naples Bay and adjacent areas, Eastern Tyrrhenian sea. Atlante di Cartografia Geologica scala 1:50.000 (progetto CARG), Servizio Geologico d’Italia (ISPRA), 2004. [20] P.R. Vail, R. Audemard, S.A. Bowman, P.N. Eisner, and G. Perez Cruz. The stratigraphic signatures of tectonics, eustasy and sedimentation. An overview. In: Einsele G., Richen W., Seilacher A. (Eds.) Cycles and Events in Stratigraphy, Berlin, Springer-Verlag. pages 617–659, 1991. [21] G. Aiello, A. Angelino, E. Marsella, S. Ruggieri, and A. Siniscalchi. Carta magnetica di alta risoluzione del Golfo di Napoli (Tirreno meridionale). Bollettino della Societ`a Geologica Italiana, 123:333–342, 2004.
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[22] G. Aiello, A. Angelino, B. D’Argenio, E. Marsella, N. Pelosi, S. Ruggieri, and A. Siniscalchi. Buried volcanic structures in the Gulf of Naples (Southern Tyrrhenian sea, Italy) resulting from high resolution magnetic survey and seismic profiling. Annals of Geophysics, 48(6):1–15, 2005. [23] E. Marsella, G. Aiello, A. Angelino, P.P.G. Bruno, V. Di Fiore, F. Giordano, N. Pelosi, A. Siniscalchi, C. D’Isanto, and S. Ruggieri. Shallow geological structures and magnetic anomalies in the Gulf of Naples: an integrated analysis of seismic and magnetometric profiles. Bollettino di Geofisica Teorica Applicata, 42(1/2):292–297, 2002. [24] M. Secomandi, V. Paoletti, G. Aiello, M. Fedi, E. Marsella, S. Ruggieri, B. D’Argenio, and A. Rapolla. Analysis of the magnetic anomaly field of the volcanic district of the Bay of Naples, Italy. Marine Geophysical Researches, 24:207– 221, 2003. [25] A. Siniscalchi, A. Angelino, S. Ruggieri, G. Aiello, E. Marsella, and M. Sacchi. High resolution magnetic anomaly map of the Bay of Naples (Southern Tyrrhenian sea, Italy). Bollettino di Geofisica Teorica Applicata, 1/2:99–104, 2002. [26] R. Bartole, D. Savelli, M. Tramontana, and F.C. Wezel. Structural and sedimentary features in the Tyrrhenian margin off Campania, Southern Italy. Marine Geology, 55:163–180, 1983. [27] A. Ascione and P. Romano. Vertical movements on the eastern margin of the Tyrrhenian extensional basin. New data from Mt. Bulgheria (Southern Apennines, Italy). Tectonophysics, 315:337–356, 1999. [28] C. Caiazzo, A. Ascione, and A. Cinque. Late Tertiary-Quaternary tectonics of the Southern Apennines (Italy): New evidences from the Tyrrhenian slope. Tectonophysics, 421:23–51, 2006. [29] R. Bartole. Tectonic structure of the Latian-Campanian shelf (Tyrrhenian sea). Bollettino di Oceanologia Teorica Applicata, 2:197–230, 1984. [30] G. Aiello, V. Di Fiore, E. Marsella, and C. D’Isanto. Stratigrafia sismica e morfobatimetria della Valle di Salerno. Extended Abstract 26th National Congress GNGTS, 495-498, 2007. [31] F. Trincardi and N. Zitellini. The rifting of the Tyrrhenian basin. Geomarine Letters, 7:1–6, 1987. [32] M. Tozzi and F. Capotorti. Tettonica trascorrente della Penisola Sorrentina. Memorie della Societ`a Geologica Italiana, 41:235–249, 1988. [33] A. Milia and M.M. Torrente. Evoluzione tettonica della Penisola Sorrentina (margine peritirrenico campano). Bollettino della Societ`a Geologica Italiana, 116:487–502, 1997. 876
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[34] G. Gars and M. Lippman. Nouvelles donne`es n`eotectonique dans l’Apennin campanien (Italie du sud). C.R. Acad. Sci. Paris, 298:495–500, 1984. [35] D. De Rita and G. Giordano. Volcanological and structural evolution of Roccamonfina volcano (Italy): origin of the summit caldera. Geol. Soc. Spec. Publ., 110:209–224, 1996. [36] M. Mariani and R. Prato. I bacini neogenici costieri del margine tirrenico: approccio sismico-stratigrafico. Memorie della Societ`a Geologica Italiana, 41:519–531, 1988. [37] A. Malinverno and W.B.F. Ryan. Extension in the Tyrrhenian sea and shortening in the Apennines as a result of arc migration driven by sinking of the lithosphere. Tectonics, pages 227–245, 1986.
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Sources and Fate of Organic Carbon on Continental Margins L. Langone, S. Miserocchi, T. Tesi Institute of Marine Sciences, CNR, Bologna, Italy leonardo.langone@ismar.cnr.it Abstract Understanding origin and chemical processes that affect the deposition and preservation of OC on continental margins is crucial for a comprehensive examination of the global OC cycle. Multi-proxy studies (%OC, C/N, δ 13 C, ∆14 C, CuO reaction products, grain-size) were conducted to investigate origin, transport and deposition of OC in the Adriatic continental shelf and Gulf of Lions. The evolution of a 100yr flood layer off the Po River delta revealed deposits up to 36 cm thick proximal to the river mouth. The flood signature of the OC is lost with time by biological and physical reworking in surface sediments, but preserved in high sediment accumulation areas. The main OC sources were identified: 2 allochthonous (degraded soil-derived and fresh vascular plant detritus) and 2 autochthonous phytoplankton (estuarine and marine). Both estuarine and plant detritus OC are confined to the Po river prodelta. On the rest of the Adriatic shelf, a significant decoupling is observed between suspended particles and surface sediments (mainly marine vs. soil-derived, respectively). In the Gulf of Lions, the plant detritus is confined to the Rhˆone prodelta while the estuarine OC is absent. Two sources (soil-derived and marine) describe the OC composition along the dispersal system, with the marine fraction increasing offshore. In the Bari and Cap de Creus canyons, OC is mainly marine, but during dense water cascades the contribution of reworked aged OC increases.
1
Introduction
Altough deltas and continental shelves receive large inputs of organic matter from both autochthonous and allochthonous sources, in literature a general consensus on the effective role of these environments in the biogeochemical C cycle is still lacking [1]. On the other hand, shelves account for ∼90 % of the OC burial in the ocean. Further studies have highlighted that in these shallow environments, OM is efficiently recycled also after deposition by early diagenesis, reducing the ultimate OC burial [2].
The allochthonous material supplied by rivers is a heterogeneous and complex mixture of organic compounds with different chemical characteristics, originating from different sources (reactive freshwater phytoplankton, soil, leaf debris, woody fragments, kerogen etc.) [3, 4, 5]. Much of the OM deposited in prodelta areas is aged due to both the preferential degradation of the marine-riverine labile fraction compared to refractory terrestrial material as well as continental denudation, which can introduce ancient OC [6, 7, 2]. The composition of the materials exported by rivers is also quite dynamic, changing in
Marine Geology
both seasonal and inter-annual time scales [7, 8]. In this context, the role of stochastic events, such as floods and storms, is not yet fully understood and quantitatively defined. To understand the main mechanisms controlling deposition and preservation of the OC in the continental margins is necessary a multiproxy approach able to take into account the dynamic nature and the conditions of high spatial and temporal variability of OM input in these environments. Particular biomarkers (CuO reaction products), such as lignin, coupled to OC contents (%OC) and stable isotopes (δ 13 C), total nitrogen (TN), radiocarbon compostition (∆14 C) and grain-size analyses, allow to characterize source, age, and spatial variability of sedimentary OM. By means of an approach “from source to sink”, it is possible to distinguish the marine and terrestrial fraction of OM; if this latter derived from soils or from vascular plant tissues, if prevail woods, leaves, or needles. In the last 8 years, our group carry out a significant effort to attain a more complete understanding of organic matter transport, deposition and burial on the continental margins, as part of EU, ONR- and Government-funded projects (EuroSTRATAFORM, HERMES, VECTOR, SOMFlood Programs). The aim of this contribution is to summarize the main results so far achieved in the two continental margins: Gulf of Lions (from the Rhˆone prodelta to the Cap de Creus canyon) and Adriatic Sea (from the Po prodelta to the Bari canyon).
ranean, after the building in 1970 of the imposing dike of Assuan on the Nile and the nearly complete exploitation of its waters for irrigation. Both these rivers originates from the Alps and drain large catchment areas (Rhˆone, 97,800 km2 , and Po, 70,000 km2 ). TOC discharge for the Rhˆone has been estimated at 15 × 104 tons y−1 , one third of which is POC. For the Po, POC discharge (13.4 × 104 tons y−1 ) is slightly higher than DOC discharge (12.1 × 104 tons y−1 ). Thus, DOC discharge rates for the two rivers are similar, while, in terms of TOC the Po export rate (25.5 × 104 tons y−1 ) is about 1.7 times higher than that of the Rhˆone, making the former river the largest contributor of organic matter to the Mediterranean. Suspended matter ( 450 samples) in the water column and surface sediment ( 900 samples) were collected along transects perpendicular to the shore in both studied areas. For elemental and stable isotope analyses, samples were pretreated with 1.5 N HCl solution to remove inorganic carbon. OC and total nitrogen (TN) contents were measured using a FISONS NA2000 Element Analyzer. Stable isotopic analyses of OC were carried out on the same samples by using a FINNIGAN Delta Plus mass spectrometer directly coupled to the FISONS NA2000 EA by means of a CONFLO interface for continuous flow measurements. On the basis of elemental and isotopic surficial distribution, 120 samples were selected for alkaline CuO oxidation analyses which were performed in collaboration with Miguel Go˜ni at the College of Oceanic and Atmospheric Sciences (Ore2 Material and Methods gon State University). The Rhˆone and the Po rivers are the two Accelerator Mass Spectrometer (AMS) ralargest freshwater inputs to the Mediter- diocarbon measurements were carried out 880
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Figure 1: Graph of the stable isotopic compositions of organic carbon (δ 13 C) vs carbonnormalized lignin phenol yields and total nitrogen:organic carbon ratio from Po and Rhˆone surficial sediments. The compositions of four possible OC sources (C3 vascular plant detritus, C3 soil OM riverine/estuarine phyto detritus and marine phyto detritus) are also plotted in all graphs to illustrate the relative influence. The main difference between a flood (blue characters) and non-flood (red charaters) event is visible in the lignin content. on selected samples at the National Ocean Sciences Accelerator Mass Spectrometry Facility (NOSAMS, Woods Hole Oceanographic Institution).
3
The autochthonous fraction of OM comes from estuarine or marine origin. Figure 1 show the distribution of parameters in flood or non-flood conditions in Po and Rhˆone prodelta sediments.
Results and Discussion
End-member mixing models based on N/C, δ 13 C and lignin data were applied to quantitatively assess the OM contributors in suspended material and surficial sediments from both study areas. Four end-members define the main OM sources: 2 allocthonous and 2 autochthonous. The terrestrial OM is allochthonous and can be fresh, constituted by vascular plant detritus, or ancient and degraded if soil-derived.
3.1
Po river prodelta and Adriatic continental shelf
A suite of scientific papers describe results on the Adriatic continental shelf [9, 8, 10, 11, 12, 13]. Here, we summarize main conclusions. The Po River (Italy) experienced a 100year flood in the October 2000. The Po flood was characterized by an exception881
Marine Geology
ally long duration and occurred during relatively quiescent physical oceanographic conditions, which allowed for the formation of a fine-grained deposit, thick up 36 cm. A quick-response survey was carried out to characterize suspended matter and sediment delivered to the prodelta and adjacent shelf. Stations were reoccupied 9 times at seasonal frequency in order to follow the time evolution of the flood layer. The survey highlighted there were two major depocenters: one located off the main distributary, Po della Pila, the other in the southern portion of the prodelta off the Tolle, Gnocca and Goro distributaries. In surface sediments peculiar characteristics of the flood OM vanished with time due to biological and physical mixing. Due to sediment sorting, the lignin content increased with time as well as the grainsize whereas the 14 C age of bulk OC age decreased. In areas characterized by the highest flood layer thicknesses, the flood signature was preserved in the sedimentary sequence. The western Adriatic sediments have been grouped into two main biogeochemical partitions: Po and Adriatic mud-wedge. Figure 2 shows lignin contents vs. δ 13 C compositions and lignin. High variability in elemental, isotopic and biomarker data was measured in the Po prodelta area, whereas the southern sediments exhibited a narrower range of values. Suspended OM is compositionally heterogeneous, with contributions from marine phyto-plankton, riverine-estuarine phytoplankton and soil-derived OM. In the Po prodelta area, elemental and isotopic compositions suggest that the phytoplankton (marine and riverine-estuarine) are the major OC constituents in the surface water column. Near the seafloor, soil-derived 882
OC is the most important source of POC. Along the Adriatic shelf, primary production is influenced by land-derived nutrients that promote phytoplankton growth which in turn influences the isotopic composition in the water column. At the same time, a fraction of fine, isotopically depleted, terrestrial material travels southward suspended within the Western Adriatic Coastal Current. The dilution of marine-derived OC with terrigenous material, located in the distal stations, generates a marked longitudinal δ 13 C gradient seaward. Soil-derived OC is the main component of the material supplied by the rivers, specially by the Po river. Riverine phytoplankton and plant fragments are additional important sources. On the contrary of suspended matter, the surficial sediments are dominated by soilderived OM (from 50 to 94%). Frequent physical reworking and re-oxidation of surficial sediments result in an efficient reactor where refractory soil-derived OC, delivered by the Po and Appennine rivers, is rapidly accumulated and preserved. Significant contributions of plant fragments are geographically limited in the Po prodelta, which exhibits the highest plant-fragment fraction in the proximal prodelta area. The riverine-estuarine phytoplankton is confined within the Po prodelta area, where the highest contribution on average is exhibited in the proximal Po prodelta area. In the Bari canyon, the main source of the organic carbon is marine. During cascading event of dense waters spilling out the Adriatic continental shelf, the contribution of OC from kerogene or soil-derived increased due to reworking of old deposits at the shelf edge and canyon head.
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3.2
Rhˆone prodelta and Gulf of sition and the absence of lignin. The advected material that reaches the Lions continental shelf
Tesi et al. [14], Fabres et al. [15] provided a significant improvement of the understanding of OM composition in the Gulf of Lions (GoL) using elemental, isotopic and lignin analyses to evaluate OM origins and to quantify the relative importance of terrigenous and marine OC on surface sediments. The composition of OM from surface sediments in different regions of the GoL are shown in Figures 1 and 3. Figure 3 displays the relationships between δ 13 C and lignin of surficial sediments from diverse depositional settings in the GoL. These settings include the Rhˆone prodelta, where high sediment accumulation rates reach 20 cm y−1 and the mid-shelf mud belt consisting of a fine-grained deposit of Holocene. The linear relationship highlighted in the graph identifies a possible mixing of fine soilderived terrestrial material, isotopically depleted, rich in OC and relatively rich in lignin with marine phytodetritus characterized by low OC, “heavier” isotopic compo-
slope and the head of the canyons has lost its original terrestrial signature and is mainly marine. Furthermore, along the dispersal system, the terrestrial OM adsorbed onto the fine material is degraded by microbial activity. In the Cap de Crues canyon, we point out the existence of an additional pool of aged OC that is supplied to the interior ocean independently of river discharge. This aged material derives from OC included in consolidated sediments deposited during lowstands of sea level. Erosion processes at the shelf edge re-expose this pool of aged OC, temporarily buried in marine sediments, to physical erosion, as well as to chemical and biological degradation.
4
Acknowledgments
Researches carried out in the framework of Projects EU-EUROSTRATAFORM, ONR-EUROSTRATAFORM, VECTOR, EU-HERMES, EU-SOMFlood.
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Figure 2: Plots of lignin contents vs. δ 13 C compositions and lignin vs. 3,5-Bd:V in the Adriatic samples. In both plots, it is apparent the variation of the considered parameters along the sediment dispersion system.
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Figure 3: Surficial OC, δ 13 C and lignin contents for the Gulf of Lions. The linear regression identifies a mixing between soil-derived OC with phytoplankton detritus. Out the prodelta, the marine contribution increases showing depleted isotopic values and low lignin contents in the offshore regions. In the RhË&#x2020;one proximal prodelta, samples from low fluvial discharge fall out the main alignment.
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