The structural evolution of this basin and of the surrounding Apenninic belt (Figures 3a-c) may be subdivided in three main phases, well differentiated in space and time (e.g. Kastens et al., 1988; Sartori, 1990; Sartori and Capozzi, 1998).
From the Tortonian to the middle Messinian (roughly 9-6 My), crustal stretching, with a nearly E-W to ENE-WSW extensional trend, only occurred in the northwestern Tyrrhenian zone, lying north of the Selli fault (Figure 3a). Negligible, or minor, accretionary activity developed in the adjacent belt (Northern Apennines) during this phase (e.g. Vai, 1987; Sgrosso et al., 1988; Borsetti et al., 1990; Castellarin et al., 1992).
From the middle Messinian to the late Pliocene (6-2 My), intense crustal stretching, with a roughly E-W extensional trend, only occurred in the central Tyrrhenian zone (Magnaghi-Vavilov basin), while intense accretionary activity affected the whole Apenninic belt, with particular regard to the southern arc, comprising the Southern Apennines and Calabria (e.g. Patacca et al., 1993). During this phase, the Northern Apennines began to be affected by thrust activity along their external front and tensional deformations in the internal zone (e.g. Elter et al., 1975; Castellarin and Vai, 1986; Vai, 1987; Bigi et al., 1989; Bartole, 1995).
Around the late Pliocene (2 My), crustal stretching ceased in the central Tyrrhenian basin and started, with roughly NW-SE extensional trend, in the southernmost Tyrrhenian zone, leading to the formation of the narrow and elongated Marsili basin. Accretionary activity ceased in the Southern Apennines (e.g. Cinque et al., 1993) and accelerated in the external Calabrian arc and Northern Apennines (e.g. Boccaletti et al., 1985; Finetti and Del Ben, 1986; Castellarin and Vai, 1986; Bartole, 1995).
In the following, we discuss on how the above deformation pattern may be interpreted as an effect of the proposed extrusion mechanism.
The fact that the opening of this basin was not accompanied by accretionary activity at the related trench zone, i.e. the Northern Apennines, implies that no lithosphere subduction has occurred during this extensional event. This inference is corroborated by the fact that the Adriatic domain, which lays in front of the Northern Apennines, was characterized by a continental crust (e.g. Boccaletti et al., 1980; Serri et al., 1991) and, thus, its consumption would have certainly left clear imprints, in terms of accretionary material, at the trench zone. This evidence implies that the formation of the northwestern Tyrrhenian basin was not connected with the development of a T-A-BA system and, thus, it does not represent an example to be discussed in this work.
A possible explanation of this extension and of the lack of coeval subduction in the Northern Apennines has been proposed by Mantovani et al. (1997), who interpreted this event as a consequence of the divergence between the confining blocks, i.e. the fixed Corsica-Sardinia microplate and the NEward drifting Adriatic promontory, triggered by the activation of the Giudicarie decoupling fault system.
On the other hand, it must be pointed out that the opening of the Northern Tyrrhenian basin cannot be explained by the slab pull model, since the lack of coeval lithosphere consumption at the trench zone contradicts one of the basic implications of the above model, i.e. the roll back of the slab.
The boundary conditions which determined the opening of this basin developed around the upper Messinian (5-6 My), when the pre-Apulian low buoyancy zone was completely consumed and a continental collision occurred between the Adriatic and the Aegean-Balkan systems, roughly in correspondence to western Greece (e.g. Mercier et al., 1989). This last event caused the suture of the above consuming boundary and the subsequent efficient transmission of the westward push of the Tethyan orogenic system on the Adriatic promontory. After this event, the Adriatic plate undertook a clockwise rotation around a pole roughly located in the northern Pannonian area (Mantovani et la., 1997, 2000a). The proposed kinematic pattern of this block is illustrated in Figure 3b. This kinematic change was allowed by the decoupling of the Adriatic plate from the surrounding regions, with particular regard to the African block. This decoupling was most probably achieved by the formation of a left lateral shear zone, the Victor Hensen-Medina fault, located in the central Ionian area along the prolongation of the Kefallinia fault (Figure 3b). Post late Miocene tectonic activity, with a transtensional regime, has been recognized in the Victor Hensen (Hieke and Wanninger, 1985; Hieke and Dehghani, 1999) and Medina faults (Rossi and Zarudzki, 1978; Ryan, 1978; Della Vedova and Pellis, 1989).
Another important tectonic event which allowed the change of the Adriatic kinematics was the detachment of the Iblean-Ventura microplate from the African block. This event was a consequence of the roughly E-W convergence between the Adriatic/northern Ionian block and the continental African margin (Tunisia), which caused the detachment and NWward escape of an African fragment, i.e. the Iblean-Ventura microplate (Figures 3b and 3c). The strike-slip lateral guides of this extrusion were the Taormina fault, to the East, and the Sicily Channel fault system, to the West. This last fault, being characterized by a leaky transform geometry in its central sector, became the site of localized extensional deformation in a pull-apart style, with the opening of some troughs (e.g. Finetti and Del Ben, 1986; Reuther, 1987; Argnani, 1993; Mantovani et al.,1997).
On its turn, the lateral escape of the Iblean-Ventura microplate caused other secondary extrusion processes in the Tyrrhenian area. In particular, the NWward indentation of the Iblean-Ventura microplate into the orogenic zone which lay in front of it, and east of Sardinia, caused the eastward expulsion of crustal wedges, at the expense of the western Apulian and Ionian zones. The oceanic-like character of the litosphere subducted during this phase is suggested by petrological evidence (Serri et al., 1993; Francalanci and Manetti, 1994; Beccaluva et al., 1994). The above extrusion process may explain the renewal of orogenic activity that occurred around the late Miocene in the southern Apenninic arc, after some My of relative tectonic quiescence (Bigi et al., 1989; Patacca and Scandone, 1989). In the wake of the westward extruding wedges, crustal stretching occurred in the central Tyrrhenian area, with the formation of the Magnaghi-Vavilov basin (Sartori and Capozzi, 1998).
The reactivation of vertical movements along the Siracusa and Apulian escarpments, representing the borders between the Ionian oceanic lithosphere and the adjacent continental zones, i.e. the Iblean block to the West and the Adriatic plate to the East (Carbone et al., 1982; Auroux et al., 1984; Finetti and Del Ben, 1986) could be interpreted as an effect of the downward flexure of the Ionian lithosphere beneath the extruding southern Apenninic arc.
Another major tectonic event which could be associated with the late Miocenic change of the Adriatic kinematics was the activation of a major SE-NW shear zone, the Schio-Vicenza fault (Cantelli and Castellarin, 1994). This discontinuity allowed a new decoupling of the main Adriatic block from its Padanian sector, this time compatible with the post Late Miocene Adriatic kinematic pattern, induced by the push of the Anatolian-Aegean system (Figure 3b). This change of motion trend in the northern Adriatic zone is also testified by the evidence that the trend of the compressional axis in the sector of the Alps lying East of the Schio-Vicenza line changed from SW-NE to SE-NW (Cantelli and Castellarin, 1994).
The opening of the central Tyrrhenian basin has been alternatively interpreted as an effect of slab pull forces (e.g. Malinverno and Ryan, 1986; Royden, 1993a,b). However, this interpretation can not easily provide convincing explanations for the timing of this extensional event (late Miocene – late Pliocene) and for the peculiar geometry of the stretched zones. At the end of the Western Mediterranean arc-trench migration, around the upper Miocene (12-13 My), a well developed slab was presumably present under the N-S segment of the Iberian-Apenninic belt and the Corsica-Sardinia microplate (Figure 2c), as indicated by volcanological evidence (Bellon, 1981; Beccaluva et al., 1994). It is not clear why this large slab did not undergo roll back in the time interval spanning from 12-13 to 5-6 My and why slab roll back just started around the late Miocene and only occurred under the Southern Apennines and the Calabrian arc, notwithstanding that well developed subducted lithosphere also existed to the north and to the south of this sector.
Another evidence which can hardly be explained as an effect of a slab pull mechanism is the formation of several oroclinal arcs in the Apenninic belt and the counterclockwise rotation of the Southern Apennines with respect to the northern part of the belt (e.g. Bigi et al., 1989; Sartori, 1990).
One must also consider that, to avoid the need of several driving mechanisms, any geodynamic interpretation of the Tyrrhenian-Apennines system should also account for the major tectonic events occurred in the surrounding regions, as, for instance, the complex time-space distribution of tectonic activity in the Pelagian zone and in the Alps (Figures 3b and 3c). No attempts in this sense have been so far made by assuming the slab pull model as driving mechanism.
Around the late Pliocene, the arrival of the continental part of the Adriatic foreland at the Southern Apennines consuming boundary determined the stop of subduction (e.g. Patacca et al., 1993). After this event, the convergence between the Iblean-African domain and the Adriatic plate, no longer accommodated by lithosphere consumption beneath the Southern Apennines, caused the lateral expulsion of the Calabrian wedge, at the expense of the Ionian domain (Figure 3c). The effects of this extrusion process might be recognized in the contemporaneous acceleration of tensional (opening of the Marsili basin) and of compressional (accretionary activity in the external Calabrian Arc) deformation respectively along the internal and external sides of the Calabrian wedge. The roughly SW-NE compressional regime that the Calabrian Arc underwent during this phase might explain the acceleration of tectonic activity with the formation of several troughs and sphenocasms, and uplift in this zone (e.g. Barone et al., 1982; Sartori, 1990; Sagnotti, 1992; Del Ben, 1993; Bordoni and Valensise, 1998).
If the formation of the narrow Marsili basin (Figure 3c) had been an effect of a slab-pull mechanism, the sinking slab should have been confined to a very narrow slice of the Ionian domain. In this case, however, one should find the decoupling zones between the sinking Ionian slice and the lateral non subducting parts of this zone. Two discontinuities which could have played such a role are the Apulian and Siracusa escarpments (Carbone et al., 1982; Finetti and Del Ben, 1986; Sartori et al., 1989; Reuther et al., 1993). However, these features are separated by a distance of several hundreds of km, i.e. much larger than the width of the presumed sinking Ionian slice.
As said before, during this phase the Calabrian Arc was affected by a fast uplift and a significant increase of its curvature. If this deformation pattern was an effect of slab-pull forces, one might wonder why a comparable uplift and disruption did not also affect the Southern Apennines in the previous phase (Messinian-late Pliocene), i.e. during the formation of the Magnaghi-Vavilov back arc basin. Some authors (e.g. Westaway, 1993; Cinque et al., 1993) suggested that the Quaternary uplift in Southern Italy might represent the isostatic response of the shallow structure to the detachment of the underlying slab. However, this hypothesis cannot be advanced for the Calabrian Arc, where the distribution of deep earthquakes (e.g. Anderson and Jackson, 1987), the results of tomographic analysis (e.g. Spakman, 1990; Selvaggi and Chiarabba, 1995; Piromallo and Morelli, 1997; Giardini and Velonà, 1991) and the study of high-frequency seismic wave propagation (Mele, 1998) indicate the presence of a continuous subducted body from shallow depths to about 500 km of depth.