High resolution seismic data coupled to Multibeam bathymetry of Stromboli island collected in the frame of the Stromboli geophysical experiment: implications with the marine geophysics and volcanology of the Aeolian Arc volcanic complex (Sicily, Southern Tyrrhenian sea, Italy)

Evolution of the Tyrrhenian-Apennines system

The irregular conformation of the African and Euro-Asiatic continental margins, with the interposed Adriatic microplate (Panza et al. 1980; Pontevivo and Panza 2002) has given very complex the geodynamics of the Mediterranean region (Figure 2). In the Middle-Late Cretaceous (about 80 My B.P.) the opening of the Atlantic Ocean has determined an inversion of the relative movement of the African plate with respect to the Eurasiatic one and has produced as a main result the closure of the Thetys ocean. The process of closure ended in the Eocene with the Alpine orogenesis (Dewey et al. 1973).

Starting from the Oligocene the phases which have determined the present-day configuration of the western Mediterranean reflect a progressive migration of the Adriatic-Ionian lithosphere in the frame of a complex subduction system, migrating westwards (Malinverno and Ryan 1986; Figure 2). From the end of the Oligocene and the beginning of the Miocene (about 26 My B.P.) the extensional phases, to the west of the trench in the back-arc area, have produced a rifting of the Ercinian crust of the European foreland, which has given origin to the opening of the Ligurian-Provenzal Ocean, with a counterclockwise rotation of about 25°-30° of the Sardinia-Corsica block, happened from 30 and 16 My B.P. (Van der Voo 1993). During this period the first compressional phases of the Apenninic orogenesis verified, during which the more internal tectonic units moved eastwards overlying the outermost ones. In the Early-Middle Tortonian (about 12 My B.P.) another compressional tectonic phase verified, during which the uplift of part of the central Apennines happened.

The opening of the Tyrrhenian basin started from the Late Tortonian with a process of rifting which began along a N-S direction next to the eastern margin of the Sardinia-Corsica block (Lavecchia 1988; Sartori 1989). During the Plio-Pleistocene, the extensional process migrated eastwards superimposing, during space and time, to the compressional tectonics involving the Apenninic nappes during their progressive advance towards the Adriatic foreland. During recent times the Apennines thrust and fold belt is in compression along the outer front, while along the axis of the chain prevail extensional faults, whose dislocations have generated the most important earthquakes of the Apenninic ridge (Istituto Nazionale di Geofisica e Vulcanologia 2005).

From the Pliocene to the Pleistocene a set of compressional phases has accompanied the superimposition of the most inner tectonic units on the outer sectors of the Adriatic foreland. The orogenesis has determined a large scale flexure, gradually migrating eastwards, with the consequent formation of the Apenninic foredeep, consequently involved in the processes of thrusting. The present-day foredeep is named the Padan-Bradanic foredeep and is located on continental crust and includes Plio-Quaternary sediments for an overall thickness of 9 km (Royden 1993). The Apenninic orogenesis has produced a crustal shortening along a E-W direction, not uniform along all the chain. This different kinematics has determined a not uniform curvature of the Apenninic ridge, which can be distinguished in two sectors (Figure 3): the northern Apennines showing a concavity towards the Tyrrhenian sea, extends from the Monferrato to the Latium-Abruzzi region; the southern Apennines, from the Molise-Abruzzi region to the Basilicata region, which doesn’t show an evident curvature and the Calabria-Peloritani Arc, with a greater curvature and concavity towards the Tyrrhenian sea (Patacca et al. 1989; 1990; Doglioni 1991).

The Northern Apennines represent an anomalous zone with respect to the Southern Apennines, since it had a different evolutionary zone. In fact, its genesis is related to the Alpine orogenesis, for the most part included among the Sangineto fault and the Soverato-Capo Vaticano fault and to the evolution of the Maghrebids, for the sector between the Soverato-Capo Vaticano fault and the Sicily (Boccaletti et al. 1990). The northern and southern sectors are put in contact along a tectonic lineament, the Ortona-Roccamonfina line, which represents a boundary between two domains with a different kinematics.

The opening of the Tyrrhenian sea

Regarding the evolution of this part of the Mediterranean several geophysical and geodynamic models have been purposed, which evidence the complex evolution of this sector.

The extensional phases, started about 30 My B.P., in the Ligurian-Provenzal basin and then continued in the Tyrrhenian basin, happened in an episodic manner in a general setting of continental collision. The two extensional phases are separated among them by about 5 My B.P. and their opening is strictly linked to the African lithosphere subducting towards W-NW.

In the Central Mediterranean the first evidence of the subduction of the Adriatic microplate started at the beginning of the Paleocene (about 60 My B.P.) with the occurrence of flysch deposits and high pressure and low temperature metamorphism (Jolivet et al. 1998). During the Oligocene (about 32 My B.P.) the first indications on the geodynamic evolution of the area come when volcanic arcs form in the area Sardinia-Provence: the extension starts at the back of the Apenninic accretion wedge (Beccaluva et al. 1989).

Between the Oligocene and the Early Miocene it begins a rifting phase, starting to the formation of the Ligurian-Provenzal Basin and to the rotation of 25°-30° in a counterclockwise sense of the Sardinia-Corsica block with an average velocity of extension of the back-arc basin of about 3–4 cm/year (Van der Voo 1993).

The extension and the subsidence happened during the Late Tortonian sign the beginning of a new episode of expansion. The lithospheric rifting causes the separation of the Calabria block and begins the opening of the southern Tyrrhenian basin leaving at its back the Sardinia and Corsica blocks. After about 5 My B.P., contemporaneously to the lateral shift of the Calabria block, a new process of rifting starts, testified by the syn-rift deposits in Sardinia and Calabria (Sartori 1989; Sartori et al. 1989). New oceanic crust forms into two separate basins (Bigi et al. 1989), first in the Magnaghi-Vavilov basin (4–5 My B.P.) and then in the Marsili basin (< 2 My B.P.). These ones are considered as small back-arc basins, being linked to the occurrence of the Ionian lithospheric slab, subducting towards NW (Marani and Trua 2002; Figure 4).

After the development of these back-arc basins, the volcanism of the southern Tyrrhenian basin is migrated from W to SE, from the Sardinia to the present-day Aeolian Arc (Serri 1997; Kastens et al. 1988; Bigi et al. 1989; Sartori 1989; Savelli, 2001), developing the present-day configuration of the Marsili back-arc basin. The average rate of opening of the southern Tyrrhenian basin during the last 5 My B.P. is of about 6 cm/y (Malinverno and Ryan 1986; Patacca et al. 1990).

During present times the lithospheric structure shows two thinned regions of oceanic crust (30–40 km) separated by lithosphere relative to the Sardinia continental block thick about 80 km (Faccenna et al. 2001); moreover, the basin is characterized by a crustal thinning, which reduce more and more up to less than 10 km in the sector towards SE (Panza et al. 2003).

The correlation of the data on the lithospheric thickness with those ones regarding the distribution of the heat flow in the Tyrrhenian region (Erickson 1970; Hutchinson et al. 1985; Mongelli 1991; Mongelli and Zito 1994) has evidenced that the Tyrrhenian sea can be divided into three parts: the northern Tyrrhenian sea, the central Tyrrhenian sea and the southern Tyrrhenian sea. The southern Tyrrhenian sea is the widest and thinnest part and shows the deeper parts of the basin (more than 3500 m). To the south-east of this area the highest values of heat flow have been recorded (Morelli 1970). The central part of the Tyrrhenian sea shows the lowest values of the heat flow (Zito et al. 2003). In the basin located northwards, in proximity to the Tuscany, the heat flow values are very high (160 mw/mq).

This trending of the heat flow seems to be linked to the active magmatism occurring in the Tyrrhenian basin and then to the subduction rate of the lithospheric slab: in fact, in the part of the Calabrian Arc where there is a faster retreating, a strong extension of the lithosphere may be observed, coupled with high heat flows. On the contrary, in the southern Apennines, where the retreating of the subduction system is very low, active volcanism is lacking, extension rate is minimum and the heat flow is low. In the central-northern Apennines, where the subduction is faster, there are active seismicity and magmatism in Latium and Tuscany and high heat flows.

Geodynamic models on the extension of the Tyrrhenian sea

In order to explain the geodynamic evolution of the Mediterranean area and in particular, of the Tyrrhenian-Apennines system several evolution models have been purposed, through which the observed structural complexities should be explained.

The main aspects which have to be considered in a general model are the relative movements between the African and European plates, the coexistence of compressional and extensional regimes, respectively on the outer and inner fronts of the Apenninic chain, the eastward migration of the above mentioned tectonic regimes, the distribution of the seismicity along the Apenninic Arc, the geochemical and petrological characteristics of the Tyrrhenian and peri-tyrrhenian magmatism, the thermal flux and the palaeomagnetic evidence.

The geodynamic model developed by Malinverno and Ryan (1986) implies that the Tyrrhenian-Apennines system evolved in relationships to the westward subduction of the Adriatic-Ionian lithosphere. The arc-trench system should progressively migrate eastwards (roll-back) due to the sinking of the subducting plate. In this model two main assumptions exist, the first one that the lithospheric plates have a plastic behavior, the second one that the evolution of a subduction zone is linked to the ratio between the velocity of convergence among the plates and the velocity of retreatment of the overthrusting plate. The geodynamic evolution of the Tyrrhenian-Apennines system should then triggered by the southeastwards retreating of the subduction zone ( Figure 5; Malinverno and Ryan 1986). The described model explains some most characteristic aspects of the Tyrrhenian-Apennines area, such as the radial geometry of the Apenninic and Maghrebide chains with respect to the Tyrrhenian basin, the contemporaneous verification of a compressional regime stress on the outer front of the chain and extensional in the back-arc zone (Frepoli and Amato 1997), the subduction under the Calabrian Arc (Amato et al. 1997) and finally, the calcalkaline composition of the Aeolian volcanism (Barberi et al. 1974; Beccaluva et al. 1994). In fact, the geochemical data demonstrate that the magmatism of the Aeolian islands is compatible with a mantle having a composition between MORB (mid-oceanic ridge basalt) and OIB (oceanic island basalt) enriched of fluids and in some cases also of sediments derived from a subducted oceanic crust.

Another geodynamic model of the Tyrrhenian-Apennines system has been purposed by Doglioni et al. (1999), based on an active subduction in the orogenic system. According to this model, the Tyrrhenian sea is interpreted as a back-arc basin, but its asymmetric shape could be the expression of an astenospheric push on a laterally heterogeneous lithosphere. The lithospheric flux proceeding eastwards is responsible of both the retreating in the same direction of the arc-trench system and of the uplift along the Apenninic chain.

Seismicity of the Tyrrhenian-Apennines system

The seismicity of the Italian Peninsula is characterized mainly by a shallow seismicity under the Apennines, a region of deep crustal seismic events eastwards of the Apenninic chain and sub-crustal earthquakes under the northern Apennines and the Calabrian Arc (Chiarabba et al. 2005). As a consequence, three main sectors, well distinguished from a seismo-tectonic point of view, may be defined, i.e. the Northern Apennines, the Southern-Central Apennines and the Calabrian Arc (Amato et al. 1997).

The Northern Apennines has a seismicity with maximum hypocentral depths reaching the 90–100 km (Amato et al. 1997). The verification of sub-crustal earthquakes, with epicenters located northwards of the 43° of latitude and to the south-west of the chain axis, has suggested the occurrence of a subducting plate under the northern Apennines (Selvaggi and Amato 1992). The crustal seismicity is widespread along all the orogenic axis and occurs in the inner part of a belt wide about 50–70 km perpendicular to the chain, corresponding to the maximum topographic heights (Amato et al. 1997).

In the Southern-Central Apennines there is only a crustal seismicity, related to extensional structures (Amato and Montone 1997) and concentrated along a belt, wide about 30–50 km and NW-SE trending (Amato et al. 1997).

In the Calabrian Arc a sub-crustal seismicity has been mainly recorded with hypocenters reaching 450 km of depth (Giardini and Velonà 1991; Selvaggi and Chiarabba 1995; Amato et al. 1997). On the other side, the shallow seismicity shows events with a magnitude greater than 4.5 and a component of compressional deformation localized along the northern coast of the Sicily. The shallow events are clearly confined to the west of the Aeolian Archipelago and appear scattered and rare in the western part, while the intermediate and deep seismicity is confined to the west of the Aeolian islands (Pondrelli and Piromallo 2003). The seismicity of the last twenty years shows the existence of a subduction of oceanic lithosphere under the Calabrian Arc; it defines a Benioff plan thick 40–50 km, wide 200 km and NW dipping with an angle of 70° up to depths of 400 km (Holcomb 1989; Amato et al. 1997; Lucente et al. 1999; Chiarabba et al. 2005).

Tomographic interpretation of the southern Tyrrhenian basin

The Southern Tyrrhenian basin has been the subject of many tomographic studies aimed at reconstructing the lithospheric structure under the sea. These types of seismological images allow to define the velocity anomalies of the P waves under the study area, translating in the spatial definition of the lateral heterogeneities existing in the lithosphere and in the upper mantle in terms of fast and slow zones with respect to a reference velocity model. The first ones are interpreted as colder and denser materials with respect to the surrounding environment and usually associated with subducted lithospheric remnants; the second ones are instead interpreted as hotter and less dense bodies, which are typical expressions of melted volcanic materials or, at sub-crustal depths, of astenospheric fluxes (Hirahara and Hasemi 1993; Piromallo and Morelli, 1997).

All the tomographic models have evidenced a high velocity zone, interpreted as the subducting Ionian slab NW dipping (Cimini 1999, 2004; Lucente et al. 1999; Montuori 2004). The slab shows an evident vertical continuity with a high immersion angle (70°-75°) in the first 400 km of depth, where it is concentrated the most part of the deep seismicity and becomes sub-horizontal at greater depths. In the model of Montuori (2004), obtained by using for the first time teleseismic events recorded by submarine stations, is well reconstructed the lateral extension, which results to be of 200 km under 150–200 km of depth, while it reduces to 100 km at shallower depths.

The high velocity zone is laterally surrounded by low velocity zones in all its vertical extension, which are interpreted as astenospheric fluxes next to the subducting plate, or as the trace of convective cells determined by the subduction (Figure 6). Moreover, the correspondence between the low velocity zone and the Aeolian Arc indicates that the subduction has generated the volcanism observed in the region (Montuori 2004).

The Aeolian islands

The Aeolian Arc is a volcanic structure, long about 200 km located in the inner margin of the Calabrian-Peloritan Arc. The arc is formed by seven emerged volcanic edifices (Alicudi, Filicudi, Salina, Lipari, Vulcano, Panarea and Stromboli) and by several submarine volcanoes surrounding the Marsili basin. The structural elements and the volcanic activity in the area allow to identify three distinct sectors: the western sector (Alicudi and Filicudi); the central sector (Salina, Lipari and Vulcano) and the eastern sector (Panarea and Stromboli; Marani and Gamberi, 2004a, 2004b; Bortoluzzi et al. 2010).

The oldest volcanic activity of the Aeolian Arc is dated back at 1.3 My B.P. and conceals the Sisifo submarine volcano, located in the western sector of the archipelago. Now the only emerged volcanoes which may be considered still active are Stromboli, Vulcano and Lipari.

The type of volcanism of the Aeolian Arc is of convergence between plates and the islands pertain to an arc-trench system, resulted from the collision between the African and Eurasiatic plates, with the occurrence of a subducting slab NW dipping under the Tyrrhenian sea (Barberi et al. 1974).

In the Aeolian volcanism, active from about one million of years, two eruptive phases separated by a period of quiescence in the Late Pleistocene may be distinguished (Keller 1974). During the first phase the islands of Alicudi, Filicudi, Panarea, Lipari and Salina have been formed. During the second phase the completion of these latter ones happened and the birth of Vulcano and Stromboli (Barberi et al. 1974).

Regarding the chemical composition of the erupted products, a time evolution may be observed which can be resumed in three series with increasing contents of potassium (Barberi et al. 1974). They are the calcalkaline series (basalts rich in Al and dacites), corresponding to the oldest products, the andesitic series, rich in K (Lipari and Stromboli) and the shoshonitic series (shoshonitic basalts and rhyolites), regarding the most recent volcanic products (Vulcano, Lipari and Stromboli).

A magmatic chamber at 2 km of depth is located under the sea bottoms separating Lipari from Vulcano and the constant low values of the Sr isotopic rate indicate that the magmas are sub-crustal in provenance and not involved by phenomena of crustal contamination. The shoshonitic nature of the most recent volcanoes indicates, moreover, that the arc is in its latest stages of evolution (Keller 1974; Sigurdsson 2000; Vidal and Merle 2000).

The nature of the submarine volcanoes is compatible with the hypothesis of an expanding marginal basin; in fact, they are composed of basaltic volcanoes, probably related to extensional fractures having a NW-SW trending, with tholeitic products at their base and alkaline products in the highest parts (Barberi et al. 1974).

Stromboli

Stromboli is the northest island of the Aeolian archipelago and has an area of 12.6 km². The volcanic edifice, localized at a latitude of 38.8°N and a longitude of 15.0°E, emerges of 924 meters above the sea level and extends up to 3000 m under the sea level. Its top is composed of two crests having a half-moon shape: the outer one is named “I Vancori”, while the inner one is known as “Pizzo sopra La Fossa” (Figure 7). Both the crests are the remnants of old volcanic edifices. The erosion has carved deep canyons along the slopes of these old edifices and a wide slope covered by the ashes of the recent activity extends from the upper part of the volcano to the south-eastern one. The main onshore drainage system have also been indicated in the sketch map of Figure 7. The active craters are not located at its top, but at 100–150 m under Pizzo Sopra La Fossa in a depression formed 5000 years B.P. as a consequence of the collapse of a part of the volcanic edifice (Figure 7). The craters, still characterized by volcanic activity are now three and their set is defined as a terrace, a structure continuously changing and in gradual increasing upwards. The lava flows run towards a large valley dipping seawards, up to water depths of 1700 meters (Romagnoli et al. 1993). It is located on the NW side of the volcano and is named Sciara del Fuoco and was formed due to large rock falls and flows (Pasquarè et al. 1993; Figure 7).

The emerged part of the Stromboli strato-volcano is mainly formed during two distinct cycles of activity (Barberi et al. 1974). The old cycle is composed of rhythmic eruptions of pyroclastic materials and lava flows, giving rise to the formation of all the eastern sector of the island. The recent cycle, during which lava flows have been mainly erupted, is responsible for the formation of all the western half of the Stromboli island. Now the recent cycle continues with the activity of the craters surrounding the Sciara del Fuoco lava flow.

Stromboli is one of the few volcanoes on the earth which has a continuous eruptive activity with periods of few days or decades lasting for about 2000 years. The persisting activity implies a magmatic chamber of great dimensions supplied with continuity and mainly consisting in explosions of moderate energy (Barberi et al. 1993). The explosions last up to 10 seconds and repeat at intervals of about 10–20 minutes and are connected to fluvio-dynamical movements of the magma in the inner of the conduit (Ereditato and Luongo 1997; Chouet et al. 1999; Ripepe et al. 2001; De Martino et al. 2004). They involve a set of fragments in the order of 10²-10⁴ kg (Chouet et al. 1974; Ripepe et al. 1993) and a gas volume in the order of 10³ m³ (Chouet et al. 1974). The sources of these explosions are concentrated at a depth of 200 m about under the top of the craters (Chouet et al. 1997; Saccorotti and Del Pezzo 2000). This moderate activity has been interrupted by episodes of greater entity, accompanied by lava fluxes happened in the 1975 (Capaldi 1978), in the December of 1985 (De Fino et al. 1988), in the 1993 (Bonaccorso et al. 1996) and the last in the December 2002 (Bonaccorso et al. 2003).

The last eruption of the Stromboli, started on the 28 December 2002, has been anticipated by an increase of the explosive activity starting from the May 2002 and by the increase of lavas in the craters from November, when from the northern flank of the second crater a small lava flow erupted, propagating in the upper part of the Sciara del Fuoco. Starting from December the explosions and the heights of the eruptions was particularly intense in the first crater, located to the NW, up to reach the 200 m with respect to the crater. The same day, the activity ended with the opening of an eruptive fracture NE trending, long about 300 m. The opening of the fracture caused a shifting towards NW of the upper part, located to the SE of the Sciara del Fuoco and the formation of small slopes parallel to the direction of the slope originated by the landslide which reached the sea generating a tsunami (Bonaccorso et al. 2003).

Navigation and positioning

The vessel was set up for Multibeam data acquisition and navigation by using the PDS2000 software by RESON. The UTC absolute time was measured and recorded at any shot produced by the PDS 2000 by the Java Daphne software (Stanghellini and Bortoluzzi 2004) interfaced to a Trimble Acutime and to a Differential Global Positioning System (DGPS). The hull-mounted 16 transducer Benthos Chirp system was used. The data flow and performance were controlled through the Communication’s Technology SWANPRO software. The Subbottom Chirp workstation received positions through a sentence by the PDS 2000; positions were therefore recorded on the XTF trace headers as latitudes and longitudes of the DGPS antenna.

Multibeam bathymetry

The acquisition of Multibeam bathymetric data was carried out through one workstation interfacing the RESON8160 Multibeam system. The adopted Multibeam system was the 50 kHz, 150° aperture RESON 8160 Multibeam, having a range of 5000 m. The sonar head was positioned on the skip’s keel using a V-shaped steel frame. A sound velocity probe at the sonar head was directly interfaced to the processor of the Multibeam system, thus providing the necessary real time data for the beam forming. In addition, two datasets were generated and stored on separate computer for back-up of data on HD and CD/DVD. The PDS2000 software was able to build a 20 m Digital Terrain Model (DTM) during the acquisition of the bathymetry in the survey area. The existing Multibeam dataset will therefore be used for an up-to-date regional bathymetric compilation.

The calibration of the Multibeam data has been carried out through the acquisition of some lines. Heading and pitch values have been easily found, whereas the roll values have been difficult to obtain due to the roughness of the sea bottom morphology.

CTD casts

The CTD casts were collected in the study area during the acquisition of Multibeam bathymetry. The CTD data were recorded by using a Mod. 911Plus SBE profiling system. The position of the CTD stations is reported in the Table 2. The raw data were recorded and processed through the processing software respectively named SEASAVE and SBEdata. The sound velocity data from the acquired profiles were immediately imported into the PDS2000 software for the Multibeam data corrections. Moreover an example of CTD data collected in this cruise is shown in Figure 8. The location of the CTD stations is reported in the upper right inset of the above mentioned figure (Figure 8).

Station	Data time UTC	Longitude	Latitude
01	2006-11-28 22.25:42	15:13.42	38:52.25
02	2006-12-02 22:38:01	15:14.00	38:42.21

Subbottom chirp

A largely spaced grid of Subbottom Chirp profiles has been recorded. It has allowed the delimitation of the main units cropping out at the sea bottom, the recognition of the morpho-structures and the calibration of the multibeam geologic interpretation. The seismic grid, radial with respect to the shoreline has been recorded on the continental slope. The location of the seismic profiles has been reported in a sketch navigation map (Figure 1). The map has been constructed using a geographic information system (GIS) importing the files related to the navigation lines recorded onboard by the PDS2000 navigation program. The data processing has been realized using the SEISPRHO software, developed for the elaboration of seismic profiles (Gasperini and Stanghellini 2009). This program processes files recorded in a SEGY format and produces, as a final result, seismic sections as bitmap images. A time variant gain (TVG) has been applied improving the quality of the seismic signal and the visualization of the scattered volcanic sequences.

Geologic interpretation particularly referred to volcanic geomorphology

The first phase of data processing and interpretation consisted of the cartographic restitution of the Multibeam data as bathymetric maps with contour isobaths and shaded relief maps for the geological interpretation of the main morpho-structural lineaments. The geologic interpretation was carried out based on the recognition of the main morpho-structures cropping out at the sea bottom. The interpretation of high resolution seismic profiles allowed us to reconstruct the structural and stratigraphic setting of the continental slope successions.

The stratigraphic units belong to the Late Quaternary depositional sequence (Catalano et al. 1996; Fabbri et al. 2002). From the Late Pleistocene to the Holocene the space and time evolution and the lateral and vertical migration of the coastal marine, continental shelf and slope depositional environments have been recognized in this sequence. The stratigraphic succession has recorded the variations of the accommodation space of the Late Quaternary deposits during the last 4^th order glacio-eustatic cycle, ranging in age between 128 ky (Tyrrhenian stage) and recent times (isotopic stage 5e; Shackleton and Opdyke 1973; Martinson et al. 1987).

The occurrence of outcrops of volcanic acoustic substratum at the sea bottom has not permitted a classical stratigraphic approach applied taking into account the stratigraphic relationships between the acoustic basement and the filling units. The systems tracts of the Late Quaternary depositional sequence, being limited by time-transgressive surfaces, may be considered equivalent to the units bounded by unconformities (Unconformity-Bounded Stratigraphic Units; UBSU; Chang 1975; Nummendal and Swift 1987; Galloway 1989; Sacchi et al. 1999).

The techniques and methods of modern volcanic geomorphology have been adopted and taken into account in the geologic interpretation of the data collected in the Stromboli volcanic island (Grosse et al. 2012; Lahitte et al. 2012; Platz et al. 2012; Procter et al. 2012; Rodriguez-Gonzalez et al. 2012; Thouret 1999; Thouret and Nemeth 2012; Torrecillas et al. 2012).

The significance of volcanic geomorphology has been improved through the quantitative classification of volcanic landforms, which blends morphometry and studies based on ground observations, remote sensing data and laboratory experiments, as well as the diversified use of airborne images and digital data acquired through radar and satellites and combined with the Digital Elevation Models data (Thouret 1999). The monitoring of morphological changes in volcanic areas provides a fundamental contribution to the comprehension of the dynamics of the volcanic systems both during the eruptions and in post-eruptive stages. The rates of geomorphological processes acting at all scales on volcanoes has to be measured to improve the theoretical aspects of volcanic geomorphology (Thouret 1999). The production and comparison of Digital Elevation Models is necessary to document and quantitatively describe the morphological evolution induced by volcanic constructional and destructive processes, such as the emplacement of lava flows, calderic collapses and gravitational instabilities. In particular, for submarine areas, the acoustic techniques as the Multibeam bathymetric surveys allows to investigate the sea floor with an increasing detail and full coverage.

Large scale instability processes, erosional landforms and debris avalanches present a particular relevance in the case study of the Stromboli island. As a consequence of rapid construction, many volcanoes are liable to massive flank or slope failures resulting from structural instability. The slope failures have produced mobile debris avalanches, travelling on long distances beyond the flank of volcanoes at high velocities. A description of geological processes involving the flank failures and the related deposits has been provided in detail by several authors (Voight et al. 1983; Siebert et al. 1987; Crandell 1988; Glicken 1991; Moore et al. 1994).

Massive landslides create specific morphology and deposits, such as horseshoe-shaped re-entraints of the volcanic edifice, such in the case of Ischia island (Chiocci et al. 1998; Aiello et al. 2009, 2012) and high steep escarpments having an amphitheater shape. The debris avalanches typically form a hummocky terrain with water filled depressions and steep flow margins and thick hummocky deposits with block and matrix facies, consisting of unsorted angular-to-subangular debris. A relationship exists between the distance run-out travelled by an avalanche and the failure volume.

As an analytic technique, the volcanic geomorphology can identify the sedimentary facies associations and the facies models for dynamic volcano-sedimentary systems and establish the criteria for recognizing volcaniclastic deposits in old volcanic successions, inferring the role of climatic and tectonic effects on transport and deposition and analyzing the characteristics of sediment gravity flows to determine relevant parameters for modeling their behavior (Thouret 1999).

Morpho-bathymetry and geology

The materials erupted by the Stromboli volcano, which shows a volcanic activity constant during geological time, slide on the surrounding slopes and deposit on the submarine flanks of the edifice, which are disrupted by submarine instabilities (gravitative mass fluxes, debris flows, debris avalanches, slides, rock falls, slumpings and erosion along channels). The flanks of the Stromboli island are characterized by the occurrence of lineaments of gravity instability, frequent on the flanks of the volcanic islands (Chiocci et al. 1998). Proceeding eastwards a chute of anastomized detritus occurs, generated by channellised fluxes of detritus organized in transversal bottom lineaments. To the north of the island, channellised fluxes act on a flat sea bottom, characterized by a low reflectivity; to the west it occurs the submarine continuation of the “Sciara del Fuoco”, representing a main lava flow always active, which can be followed for several kilometers before that it laterally joins the Stromboli canyon.

The Stromboli canyon represents one of the most important canyons of the Tyrrhenian sea and links the shoreline of the Sicily with the Tyrrhenian bathyal plain surrounding the Aeolian islands. The canyon receives channellised fluxes on its right side (among them the Gioia canyon) and mass fluxes on its left side (also on the flanks of the Stromboli volcanic island), where its thalweg touches the Aeolian volcanic archipelago.

A shaded relief image map of the Stromboli island has been constructed and interpreted (Figure 9). The overall extension of the bathymetric survey is of about 910 km² in a bathymetric range from 11 m to 2555 m. The Multibeam data give new interesting evidence to understand the morphological, volcanological and structural setting of the Stromboli island, particularly referring to gravity instability processes, focusing on lateral collapses on the flanks of the volcanoes.

The geological interpretation of the Digital Terrain Model of the Stromboli island shows several main morphological lineaments (Figure 9). On the right of the island it is worth noting the occurrence of the Stromboli canyon, fed by a lateral tributary channel, draining the volcaniclastic input coming from the submerged volcanic edifice of the Panarea island. Two main morpho-structural lineaments have been identified, one having a N-S trending and another having NE-SW trending. The submerged flanks of the volcanic edifice are dissected by morpho-structural lineaments controlled by the activity of submarine instability processes along steep slopes. On the upper right side in the Multibeam map it is possible to see the Strombolicchio plateau, located at water depths ranging between 50 m and 150 m isobaths. On the upper left side in the Multibeam map it is worth note the occurrence of the end of the Angitola canyon, draining the volcaniclastic input coming from the continental slope off Calabria.A slope map of the Stromboli volcanic edifice offshore has been constructed based on Multibeam bathymetric data (Figure 10). The slopes of the upper submerged part of the stratovolcano are minor than 10°. A steep break in slope signs the passage to the lower part of the volcanic edifice, showing slopes ranging between 10° and 25°. The last slope gradient occurs in correspondence to the passage from the basal part of the edifice, characterized by slopes ranging between 0° and 5°. Note that the margin of the submerged “sciaras” (lava flows), of the more pronounced channels and of the Stromboli canyon show slopes in the order of 35°-40°.

Detailed shaded relief maps of the south-eastern and south-western flanks of the Stromboli volcano have been constructed (Figures 11 and 12). The circles indicate the location of the three OBS stations used for a location of a refraction seismic survey carried out during the cruise STRO-06 in a preliminary project phase. It is worth noting the break in slope which signs the passage from the base of the submerged volcanic edifice of the Stromboli island and the bathyal plain, located at around – 2100 m of water depths, at the passage with the volcanic edifice of the Panarea island. In the lower part of the figure a detailed shaded relief map of the south-western flank of the island is reported. The circles indicate the location of six OBS stations positioned before the execution of the wide angle refraction seismics (Marsella et al. 2007a; 2007b). Parasitic vents related to the main volcanic edifice have also to be noted.A shaded relief and contour map of the isobaths of the submerged volcanic edifice has also been constructed (Figure 13). On the right in the map of Figure 13 the Stromboli canyon has been identified, draining the volcaniclastic supply coming from the submerged volcanic edifice of the Panarea island. Two main morpho-structural lineaments, respectively N-S and NE-SW trending have been recognized. The submerged flanks of the volcanic edifice are involved by morpho-structural lineaments controlled by gravitational instabilities along steep slopes (Figure 13). In the upper right of the map the Strombolicchio plateau has been identified at water depths ranging between 50 and 150 m. In the upper left of the figure the lower part of the Angitola canyon has been identified, draining from the volcaniclastic input coming from the Calabria region (Figure 13). The northern flank of the Panarea volcanic edifice has been identified in the lower part of the Figure 13.

The constructed bathymetric profile ABCD (Figure 14) runs parallel to the break in slope joining the steep upper part of the submerged volcanic edifice with its lower part, having less steep slopes. The AB segment crosses the submarine prosecution of the “Sciara del Fuoco” lava flow and shows hints of an intense channelization, while the BC segment is comprised between two minima of water depths in the bathyal plain, crossing a deep channel located in the north-eastern flank of the volcanic edifice, having a structural control (Figure 14). The CD segment crosses the south-eastern flank of the submerged volcanic edifice, characterized by submarine gravity instabilities (Figure 14).

The constructed bathymetric profile ABC runs along the south-eastern flank of the Stromboli volcanic edifice and crosses the Stromboli canyon (Figure 15). The AB segment shows an overall decrease of water depth proceeding from north to south and hints of several channels with several order of dimensions (Figure 15). The BC segment shows a section of the Stromboli canyon, having a flat thalweg and asymmetrical levees (Figure 15).

Interpretation of Chirp seismic data

A sketch map showing the OBS and shot locations around Stromboli has been constructed for a better visualization of the shots of the seismic data collected offshore the island (Figure 16). Some seismic profiles have been processed and interpreted referring to the above mentioned location map. Seismic processing consisted of a ri-lecture of seismic data using the SEISPRHO (software) allowing the generation of bitmap files starting from SGY files of seismic lines (Gasperini and Stanghellini 2009). The processed and interpreted seismic lines are listed in the Table 3.

Several seismic units have been identified based on the geologic interpretation of Subbottom Chirp profiles recorded around the volcanic edifices and interpreted as volcanic acoustic basement pertaining to the volcano and overlying slide chaotic bodies emplaced during its complex volcano-tectonic evolution. They are related to the eruptive activity of Stromboli, mainly poliphasic and to regional geological processes involving the geology of the Aeolian Arc.

The interpretation of the Chirp line Stromboli canyon P (Figure 17) has suggested the occurrence of a volcanic acoustic basement, genetically related to the Stromboli lavas, located at depths among 30 and 60 msec. This basement is overlain by a relatively thick sequence, characterized by parallel seismic reflectors and interpreted as Late Pleistocene fine-grained sediments (Figure 17). This sequence is overlain by another sedimentary sequence, recognized up to the sea bottom and characterized by a scattered seismic signal of high amplitude. It has been interpreted as Late Pleistocene coarse-grained marine sediments.

The interpretation of the Chirp line Stromboli canyon (Figure 18) has confirmed the stratigraphic setting seen in the above mentioned profile. A volcanic acoustic basement has been identified at an average depth of 40 msec. This basement is covered by a seismic sequence, characterized by parallel reflectors and interpreted as Late Pleistocene fine-grained sediments (Figure 18). This sequence is overlain by another seismic sequence, occurring up to the sea bottom and characterized by a scattered seismic signal having a high amplitude. The sequence has been already interpreted as Late Pleistocene coarse-grained marine sediments. Similar results have been obtained through the interpretation of the Chirp line Stromboli canyos (Figure 19).

The interpretation of the Chirp line Stromboli esu has evidenced the occurrence of a volcanic acoustic basement overlain by a seismic sequence interpreted as Late Pleistocene marine coarse-grained sediments (Figure 20). In the first subbottom a thick seismic sequence characterized by hummocky facies has been interpreted as debris avalanche deposits genetically related to the Stromboli volcano (Figure 20). Similar results have been obtained through the geologic interpretation of the Chirp line Stromboli 1 est (Figure 21) and of the Chirp line Stromboli 1 esv (Figure 22). The same seismic sequences have been recognized also on the seismic lines Stromboli 1 esw and Stromboli 1 esx (Figures 23 and 24).