Journal of the Virtual Explorer

As Merla (1951) early recognized, the eastward moving NA compressional belt has generated progressively younger foreland basins (foredeep and piggy back basins, Ricci Lucchi, 1986; Ori et al. 1986) that were successively incorporated in the fold and thrust belt: the Adriatic Sea represents the younger and easternmost foredeep, active until at least the early Pleistocene. A regional review of the flexural foreland basins of the Apennines has been recently offered by Casero (2004).

The areal distribution of the Miocene-Quaternary foreland basins of the NA is shown in Fig. 1: however, it is necessary to keep in mind that the present-day distribution of these deposits, which presently crop out mostly in the syncline areas, does not necessarily reflect the original shape and size of the original basins, since the width and reciprocal distance between the basins has been substantially reduced by later orogenic contraction.

The diagram in Fig. 4 summarizes the ages of the successions infilling the post-Burdigalian, syntectonic foreland basins, along a transect from Perugia to Ancona, grossly corresponding to the CROP03 profile, thus illustrating the overall eastward migration, through time, of the compressional deformation.

The syn-compressional basins typically evolve through three subsequent stages (fig. 4) reflecting the depositional zones of a foreland basin system (e.g. Decelles and Giles, 1996). The resultant sedimentary succession consists of, bottom to top: i) forebulge/backbulge deposits, typically consisting of hemipelagic marls (e.g. Schlier Fm.); ii) foredeep deposits, mainly turbidite sandstones, with variable amounts of intercalated marls (e.g. Marnoso-Arenacea Fm.); iii) wedge top deposits, showing a larger variability of sedimentary environments and facies. During their evolution these basins were affected by frequent variations in the direction of the paleocurrents (e.g. longitudinal from NW vs. transversal from SW), as well as in the provenance of the sediment supply (e.g. Alps vs. Inner Apennines). Synsedimentary tectonics is reflected by the recurrent occurrence of large-scale mass wasting complexes, emplaced in the inner margin of the Oligocene-Miocene foredeep basins of the NA (Lucente and Pini, 2008): these complexes, including both extrabasinal (i.e. olistostrome) and intrabasinal displaced sediments, also show progressively younger ages moving from west to east. The Pliocene-Quaternary denudation complexes, recognised by Ori et al. (1986) in the seismic profiles through the Adriatic foredeep may be modern analogues of the Miocene mass wasting complexes.

Along the CROP03 transect, the Miocene-Quaternary foreland basins of the NA have been divided into three major groups, corresponding to, from West to East: i) the Marnoso-Arenacea foredeep; ii) the Inner Marche thrust-top basins; iii) the Marche-Adriatic foredeep.

1. The Marnoso-Arenacea Fm. (Late Burdigalian – Late Serravallian) was deposited in a 'complex foredeep’ (sensu Ori et al., 1986), presently cropping out in the region comprised between the eastern front of the Tuscan Units and the main ridge of the Apennines (Fig. 1). The onset of the foredeep was driven by the internal embrication of the innermost Tuscan units, that later overthrust the inner part of the Marnoso-Arenacea basin (Barchi et al., 1998b). Synsedimentary tectonics during the Late Serravallian is suggested by the presence, within the Marnoso-Arenacea succession, of olistostromes and proximal fan deposits, as well as of coeval, slumping phenomena (Ricci Lucchi and Pialli, 1973; Ridolfi et al., 1995). At present this formation is involved in four main tectonic units, consisting of progressively younger successions, moving from West to East (Menichetti and Pialli, 1986). The sedimentation is closed by the Lower-Middle Tortonian M. Vicino sandstones, a thrust-top basin, developed over the easternmost part of the foredeep, immediately to the west of the main Apennines Ridge (fig. 1).

2. The Inner-Marche thrust-top basins (Cantalamessa et al., 1986) developed in the Tortonian-Messinian time interval, when no foredeep basin of regional extent existed, but there were only minor basins, restricted between the growing Umbria-Marche folds. Most of these basins (Urbania, Serraspinosa, M. Turrino, Fabriano, Camerino) are presently located above the main ridge of the Umbria-Marche Apennines, or immediately to the west (M. Vicino) and to the east (Laga) of the mountain belt (fig. 1).

3. The Marche-Adriatic syn-orogenic succession was deposited in a foredeep in the Late Messinian-Quaternary time interval. The onset of this basin was related to the major compressional phase of the Umbria-Marche Apennines (Late Messinian-Early Pliocene). Soon after its onset, the forebulge of the Adriatic foreland reached the outermost Dinarides and stopped migrating, so that the topography of the Marche-Adriatic foredeep was affected by a set of syn-sedimentary folds and thrusts (complex foredeep, sensu Ori et al., 1986). The structural setting and the timing of deformation of the Marche-Adriatic foredeep, as well as that of its northern equivalent, the Po Plain, have been depicted in detail in many papers, combining surface geology, biostratigraphy and good quality seismic reflection profiles, calibrated by many boreholes (e.g. Pieri and Groppi, 1980; Castellarin et al., 1985; Argnani et al., 1991; Coward et al., 1999; Scarselli et al., 2006).

Summarising, the age of the foreland basins (fig. 4) effectively illustrates the general eastward migration of the contractional deformation from Late Burdigalian in the Tiber Valley zone to Adriatic coast and a regular superposition of depositional environments, where small mobile basins restricted between the growing ridges are superimposed on regionally extended foredeep basins (Ori et al., 1986). However, the migration is not completely regular: the Marnoso-Arenacea and the Marche-Adriatic basins are regionally extended foredeep, related to the main embrication events of the Tuscan Units (Late Burdigalian-Serravallian) and of the Umbria-Marche Units (Late Messinian-Pliocene) respectively. In the intervening time period (Late Tortonian-Early Messinian) no proper foredeep was present, but only minor basins, restricted between the growing Umbria-Marche folds.

The rate of migration of the compressional wave can be estimated by progressively restoring balanced cross-section through the Umbria-Marche fold and thrust belt. A study made by Basili and Barba (2007) along the CROP03 profile concluded that deformation advanced steadily towards NE, at a time-average rate of about 6 mm/yr. More detailed studies, performed where the timing of deformation is well constrained by syntectonic deposits, suggest a not entirely steady migration of deformation, where a major factor of perturbation is the occurrence of major fall or rise of the base level ( e.g. Scarselli et al., 2007).

It is well known that surface transport processes can influence the evolution of a thrust belt, modifying its regular migration from the hinterland to the foreland (e.g. Simpson, 2006). For example, during the Messinian salinity crisis (about 5.6 Ma) the Mediterranean sea level experienced a drop of about 1500 m, followed by a rapid early Pliocene marine ingression. These perturbations affected the rate of migration of the compressional front (and the magnitude of shortening) in the Marche-Adriatic region (Scarselli et al., 2007) and in the Po Plain area as well (Castellarin et al., 1985). Quaternary glaciations may also have had significant effects on the late evolution of the thrust belt and of the related basins.

Simultaneously with the development of the eastward moving NA compressional belt, at its back the TD was involved in extensional tectonics, continuously migrating through time from west to east, from Eastern Corsica to the Tuscan mainland, and deeply dissecting the previously formed compressive structures. This process, described by Elter et al. (1975) earlier, was successively confirmed by other research (e.g. Lavecchia et al., 1984; 1987; Carmignani et al., 1994 among many others).

The diagram in Fig. 4 summarizes the ages of the successions infilling the extensional hinterland basins along the CROP03 transect and its offshore continuation across the Northern Tyrrhenian Sea. In strict analogy with the compressional history, the age of the successions systematically decreases from the Corsica basin through the Northern Tyrrhenian Sea and Tuscan Mainland, to reach the Umbria-Marche Apennines ridge in the Early Pleistocene, where extension is presently active (Bartole et al., 1991; Jolivet et al., 1998b; Pascucci et al., 1999, 2007; Collettini et al., 2006).

The diagram of fig. 4 also shows that magmatic activity also migrates eastward, following the formation of the extensional basin. The magmatic bodies are emplaced late in the extensional history, after the major rift phase, possibly because the emplacement of the magmatic bodies requires that a significant crustal extension has already occurred. See Peccerillo (2005) for a comprehensive review of the Pliocene-Quaternary volcanic processes of the Apennines.

Pascucci et al. (1999) describe the evolution of a typical extensional hinterland basin, consisting of three major subsequent stages: the first pre-rift stage is characterised by narrow bowl-shaped basins or flat-like deposits, whose location is possibly affected by the pre-existing topography and tectonics; during the second syn-rift stage the activity of the major extensional faults promotes the subsidence of triangular shaped, asymmetrical half-grabens; finally, during the third post-rift stage wide bowl-shaped or blanket-type deposits are draped above the previously formed depressions. The three subsequent stages have followed each other through time and space, moving from the westernmost offshore (i.e. Tyrrhenian) areas towards easternmost inshore (fig. 4).

The hinterland basins and the extensional faults driving their evolution have been effectively imaged by many seismic reflection profiles, located on both the Northern Tyrrhenian Sea and the Tuscan Mainland (e.g. Bartole et al., 1991; Pascucci et al., 1999; 2007). Shallow extensional faults, producing the direct superposition of younger over older rocks, exhumed by a subsequent uplift, have been also mapped at the surface (e.g. Zuccale fault, Keller et al., 1994; Collettini and Barchi, 2004) and/or drilled by deep wells in the geothermal areas of Larderello (Batini et al., 1985; Brogi et al., 2003) and M. Amiata (Calamai et al., 1970; Brogi, 2004).

These studies produced significant advances in the knowledge of the extensional process.

The extensional deformation is strongly asymmetric and dominated by a set of east-dipping, low-angle normal faults (Barchi et al., 1998a; Decandia et al., 1998), whose location is strictly connected with the position of the shallow marine and/or continental syn-tectonic basins.

The low-angle normal faults are detached within the upper crust. Beneath Tuscany, the basal detachment corresponds to a prominent reflector (k-horizon, Cameli et al., 1993), located at an average depth of 10 km, and it slightly deepens towards the east, reaching a maximum depth of about 5s (up to 14 km). This is the same depth of both the Brittle/Ductile transition inferred from the thermal field (Pauselli and Federico, 2002) and the cut-off of seismicity (Chiarabba et al., 2005; De Luca et al., 2009). The underlying lower crust shows an anomalously reduced thickness (about 9 km), possibly produced by ductile flow. In this view, the Tyrrhenian Domain has been extended by brittle/semi-brittle fault zones in the upper crust, coupled with prevailing ductile flow in the lower crust.

The easternmost fault of this extensional system is an active, east-dipping low-angle normal fault (Alto-Tiberina Fault, ATF), exposed in the Umbria region, at the western border of the AD (Boncio et al., 1998; 2000; Collettini and Barchi, 2002). The ATF cuts the thickened crust down to a depth of about 14 km, greater than that reached by the older, no longer active faults of the TD, reflecting incipient extension. Strong evidence for the present-day activity of the ATF is furnished by microseismicity surveys, where most of the recorded seismicity fits the trace of the ATF (Boncio et al., 1998; Chiaraluce et al., 2007). The focal mechanisms show extensional kinematics with NNW-SSE trending planes, parallel to the ATF fault, and the resulting stress tensor (with a vertical σ₁ and a ENE trending σ₃) is consistent with both the presently active stress field (Montone et al., 2004) and the Quaternary long-term stress field (Lavecchia et al., 1994; Boncio et al., 2000) of the region. The present day activity of ATF (and of some of its high-angle splays) is also suggested by geomorphological evidence (Cattuto et al., 1995), and confirmed by GPS data, showing about 2.5 mm/yr of NE extension across the High Tiber valley (D'Agostino et al., 2009; Hreinsdóttir and Bennett, 2009), as well as by high resolution seismic reflection survey across the SW margin of the Sansepolcro basin (Delle Donne et al., 2007). ATF is suitable to represent the presently active expression of the extensional wave of the NA.