2.a. Model set-up

To simulate the transition from rifting to oceanic spreading, a time-dependent 2D thermo-mechanical numerical model has been used in which the dynamics of the crust–mantle system have been investigated by numerical integration of the three fundamental equations of conservation of mass, momentum and energy:

(1) $$\begin{equation} \nabla \cdot \vec{u} & =& 0\\ \end{equation}$$

(2) $$\begin{equation}- \nabla P + \nabla \cdot \vec{\tau } + \rho \vec{g} & =& 0\\ \end{equation}$$

(3) $$\begin{equation} \rho {c_p}\left( {\frac{{\partial T}}{{\partial t}} + \vec{u} \cdot \nabla T} \right) & =& \nabla \cdot \left( {K\nabla T} \right) + \rho {H_d} \end{equation}$$

respectively, where $\vec{u}$ is the velocity, P is the pressure, $\vec{\tau }$ is the deviatoric stress, ρ is the density, $\vec{g}$ is the gravity acceleration, c_p is the specific heat at constant pressure, T is the temperature, K is the thermal conductivity and H_d is the radiogenic heat production rate per unit mass.

Equations (1), (2) and (3) are solved using the 2D finite-elements code SubMar, which has been exhaustively described by Marotta, Spelta & Rizzetto (Reference Marotta, Spelta and Rizzetto2006). This numerical code uses the penalty function formulation to integrate the equation for the conservation of momentum and the streamline upwind/Petrov-Galerkin method to integrate the equation for the conservation of energy.

The marker in-cell technique has been used to compositionally differentiate crust and mantle rocks.

A viscous-plastic behaviour has been assumed for both materials. The effective viscosity is calculated as follows:

(4) $$\begin{equation} {\mu ^{{\rm{eff}}}} = {\mu _{{\rm{viscous}}}} = {\mu _{0,i}}\exp\left[ {\frac{{{E_i}}}{R}\left( {\frac{1}{T} - \frac{1}{{{T_0}}}} \right)} \right] \end{equation}$$

where μ _{0, i} and E_i are the reference viscosity at the reference temperature T ₀ and the activation energy for the crust (i = c) and the mantle (i = m), respectively, with a maximum value defined by the plastic viscosity assumed equal to 10²⁵ Pa s (Table 1).

Table 1.

Material properties used in the 2-D numerical modelling

^a Dubois & Diament (Reference Dubois and Diament1997) and Best & Christiansen (Reference Best and Christiansen2001); ^bRybach (Reference Rybach, Haenel, Stegena and Rybach1988); ^cRanalli & Murphy (Reference Ranalli and Murphy1987); ^dChopra & Peterson (Reference Chopra and Peterson1981); ^eHonda & Saito (Reference Honda and Saito2003), Arcay, Tric & Doin (Reference Arcay, Tric and Doin2005) and Roda, Marotta & Spalla (Reference Roda, Marotta and Spalla2010).

We account for a brittle behaviour of the crust to define the rheological conditions for mantle serpentinization only, as better specified in Section 2.b.

The material parameters are listed in Table 1. Initially, the lithosphere is mechanically unperturbed and laterally homogeneous. The initial thickness of the continental crust is assumed to be 30 km (Fig. 2), which is in agreement with models envisaging the beginning of extension-transtension onto a lithosphere characterized by a crust with a thickness of approximately 30 km (Muntener, Hermann & Trommsdorf, Reference Muntener, Hermann and Trommsdorf2000; Manatschal, Reference Manatschal2004).

Figure 2.

(a) 2D geometry and numerical set-up of the model. (b) Thermal and rheological profiles at the beginning of the evolution for the hot (red lines) and cold (blue lines) models. The solid lines indicate the effective viscosity profiles, corresponding to the geotherms indicated by the dashed lines.

Two thermal settings are proposed here to satisfy two contrasting pre-rifting settings of the Alpine lithosphere characterized by different depths of the 1600 K isotherm: 80 and 220 km (Fig. 2b). We refer to these simulations as the hot and cold models, respectively. The initial thermal conditions correspond to an almost conductive thermal profile from 300 K at the surface to 1600 K at the base of the lithosphere; an initial homogeneous temperature of 1600 K is assumed below the lithosphere (Fig. 2b).

Boundary conditions are defined in terms of the temperature and velocity. A temperature of 300 K is fixed at the top of the crust and throughout the air–water layer, a temperature of 1600 K is fixed at the base of the model and zero flux is assumed through the lateral sides of the model. We apply an extension rate of 1.25 cm a⁻¹ at both lateral sides of the model throughout the crustal thickness, resulting in a symmetric passive rifting with a total extension rate of 2.5 cm a⁻¹, compatible with the magma-poor nature of the rift (e.g. Manatschal & Müntener, Reference Manatschal and Müntener2009). The 2D domain is closed vertically with shear-free conditions prescribed along the top and the bottom of the model domain, while both crust and lithospheric mantle are allowed to exit the model boundaries allowing the thinning of either crust and lithospheric mantle (Fig. 2a).

Considering that in the Alpine literature the onset of rifting is proposed to occur during 220–200 Ma (Müntener, Hermann & Trommsdorf, Reference Muntener, Hermann and Trommsdorf2000; Manatschal, Reference Manatschal2004; Montanini, Tribuzio & Anczkiewicz, Reference Montanini, Tribuzio and Anczkiewicz2006; Piccardo, Padovano & Guarnieri, Reference Piccardo, Padovano and Guarnieri2014) and that the oldest rocks of the ophiolitic associations belonging to the Liguria–Piemonte Ocean have been dated at 175–160 Ma (Tribuzio, Thirwall & Vannucci, Reference Tribuzio, Thirwall and Vannucci2004; Rossi et al. Reference Rossi, Lahondère, Cocherie, Caballero and Féraud2012; Kaczmarek, Müntener & Rubatto, Reference Kaczmarek, Müntener and Rubatto2008; Li et al. Reference Li, Faure, Lin and Manatschal2013, Reference Li, Faure, Rossi, Lin and Lahondère2015), we run the simulation over a time span of 60 Ma to match the time interval needed to generate the oldest gabbro intrusions in the natural system.

2.b. Conditions for mantle serpentinization

Magmatic-poor margins formed by the rifting of continental crust are characterized by the occurrence of serpentinized peridotites within a broad continent-ocean transition (e.g. Pérez-Gussinyé et al. Reference Pérez-Gussinyé, Reston, Phipps Morgan, Wilson, Whitmarsh, Taylor and Froitzheim2001; Manatschal, Reference Manatschal2004) and by the lithostratigraphy of the ophiolitic sequences (Mevel, Caby & Kienast, Reference Mevel, Caby and Kienast1978; Lagabrielle & Cannat, Reference Lagabrielle and Cannat1990; Chalot-Prat, Reference Chalot-Prat, Foulger, Natland, Presnall and Anderson2005; Manatschal et al. Reference Manatschal, Sauter, Karpoff, Masini, Mohn and Lagabrielle2011; Li et al. Reference Li, Faure, Lin and Manatschal2013). The role of serpentinization of the lithospheric mantle during continental rifting and the transition to oceanic spreading has been extensively discussed by Pérez-Gussinyé et al. (Reference Pérez-Gussinyé, Reston, Phipps Morgan, Wilson, Whitmarsh, Taylor and Froitzheim2001, Reference Pérez-Gussinyé, Morgan, Reston and Ranero2006). They assume that mantle serpentinization occurs when the overlaying crustal layer is under brittle conditions, so that faults can cut across the crust allowing hydrous fluids to penetrate the mantle, and mantle matches the appropriate pressure and temperature conditions for the stability field of serpentine.

In order to implement mantle serpentinization and the consequent rheological and compositional changes, we check whether the pressure and temperature of each mantle-type marker match the stability field of serpentine and the overlying crustal layer is under brittle conditions. To define whether the overlying layer is under brittle conditions, we use a simplified formulation of Byerlee's law criterion:

(5) $$\begin{equation} {\sigma _{{\rm{brittle}}}} = \beta \cdot y\quad \quad {\rm{with}}\quad \quad \beta = 16\;{\rm{MPa}}\;{\rm{k}}{{\rm{m}}^{ - 1}} \end{equation}$$

where y is the depth, and compare the brittle strength σ_brittle with the temperature- and pressure-based plastic strength:

(6) $$\begin{equation} {\rm{Regime}} \Leftarrow \min \left\{ {{\sigma _{{\rm{brittle}}}},{\sigma _{{\rm{plastic}}}}} \right\}. \end{equation}$$

2.c. Conditions for partial melting

During an active extension of a continental lithosphere, the temperature in the lithospheric mantle increases as a consequence of the upwelling asthenospheric flow. If the pressure and temperature conditions of peridotite solidus are matched, partial melting of the lithospheric mantle occurs and gabbroic-basaltic melts form. We assume here that once the mantle partial melting occurs in the system, the oceanic lithosphere starts to form. This implies an instantaneous transfer of the mantle melt to the surface in agreement with the estimates of the rate of magma ascent across the continental and oceanic crust (Clague, Reference Clague1987; Turner et al. Reference Turner, George, Evans, Hawkesworth and Zellmer2000).

In order to individuate the beginning of oceanic spreading and to identify the extension of the partially molten mantle region, we check the pressure and temperature conditions of each mantle-type marker during the evolution. When the P–T conditions reach the dry solidus field of peridotite (Rogers et al. Reference Rogers, Blake, Widdowson, Parkinson and Harris2008):

(7) $$\begin{equation} {P_{\rm{m}}} \le \alpha \cdot {T_{\rm{m}}} + \beta , \end{equation}$$

where α = 0.00789792857 GPa K⁻¹ and β = – 11.1071202 GPa, its typology is changed from mantle type into potential partially molten mantle type.

The extension of the partially molten mantle region is shown in Figure 3 distinguishing the oceanic lithosphere (dark green), formed by serpentinized mantle hosting gabbros and basalts that can be produced once partial melting conditions are attained, from the mantle that serpentinized before the occurrence of partial melting (light green).

Figure 3.

Successive stages of the tectonic evolution predicted by the hot (panels a_i) and cold (panels b_i) models at different times after the beginning of the forced extension. Black and red dashed lines correspond to 800 K and 1500 K isotherms, respectively. Ages refer to the time span from the beginning of the simulations.

In the present form, the compositional and rheological changes consequent to partial melting are not implemented.

3.a. Structural configurations

Figure 3 shows the structural configurations of the crust and lithosphere at different stages of the evolution of the hot model (panels a_i, left side) and the cold model (panels b_i, right side). The crustal boundaries coincide with the envelope of the crustal type markers. The base of the lithosphere is thermally defined by the 1600 K isotherm (red dashed line).

During the initial phase of forced extension in both models a progressive thinning of the crust is predicted, mostly concentrated around the position of the future ridge. The crustal thinning occurs very early in the cold model, approximately 1 Ma after the beginning of the extension (Fig. 3b₁), while more than 10 Ma passes before thinning becomes significant in the hot model (Fig. 3a₁). After 15.4 Ma (hot model, Fig. 3a₂) and 4.4 Ma (cold model, Fig. 3b₂), the thermal thinning is localized around the position of the future ridge. These times mark the beginning of mantle serpentinization in both models. The long time interval necessary in the hot model for the beginning of deformation localization therefore seems related to the higher thermal state at the base of the continental crust, which results in a different rheological behaviour of the shallow lithospheric mantle as discussed at the end of Section 2.c.

The progression of forced extension leads to the occurrence of the crustal break-up which occurs at 31.4 Ma in the hot model, approximately 16 Ma after the beginning of the mantle serpentinization (Fig. 3a₃), and at 7.4 Ma in the cold model, only 3 Ma after the beginning of the mantle serpentinization (Fig. 3b₃). The occurrence of crustal break-up is followed by the exhumation of the serpentinized mantle replacing the thinned continental crust at the floor of the basin. In the hot model, the mantle serpentinization progresses even at the base of the continental crust for more than 250 km away from the ridge. In contrast, in the cold model deep-seated serpentinization does not occur. This is a consequence of the thinner continental crust characterizing the hot model, allowing the occurrence of mantle at shallower depths with respect to the cold model in which continental lithospheric mantle resides at greater depths and P–T conditions are inappropriate for serpentinization.

The pressure and temperature conditions at the base of the lithosphere in the hot model are favourable for partial melting for a relatively short time after the crustal break-up (approximately 5 Ma, Fig. 3a₄). In contrast, in the cold model favourable pressure and temperature conditions are reached after a relatively long time from the break-up (approximately 15 Ma, Fig. 3b₄). Although there are different time intervals between the different stages (thinning, serpentinization and partial melting) in the hot and cold models, once crustal thinning starts comparable time spans pass before the beginning of the mantle partial melting in both models (approximately 21 Ma and 18 Ma for the hot and the cold models, respectively; Fig. 4).

Figure 4.

Thermal and velocity fields of the system at different times after the beginning of the forced extension for the hot (panels a_i) and cold (panels b_i) models. Time scale in the centre of the figure represents the simulation duration for both models, on which the onset of serpentinization, crustal break-up and partial melting are located by arrows (red for the hot and blue for the cold models). Time intervals between serpentinization and partial melting onset are indicated by dashed lines.

At the beginning of the extension, the low strength of the lithosphere in the hot model reduces the efficiency of stress transmission up to shallow depths, making the localization of crustal thinning slower than that for the stronger lithosphere of the cold model. Once the crustal thinning is localized, the successive evolution is dominated by local thermal gradients because this stage is comparable in the two models. After their formation, the mantle, which is depleted by partial melting (dark yellow markers for the partial melting area and light yellow markers for the depleted mantle in Fig. 3a₅, b₅), moves laterally below the expanding oceanic crust, forming the oceanic lithosphere (dark green markers in Fig. 3a₅, b₅).

Figure 5.

Surface horizontal crustal velocity predicted by the hot (a) and cold (b) models at different times after the beginning of the forced extension.

Here we assume that the oceanic lithosphere forms after the beginning of the mantle partial melting, when gabbros, basalts and part of serpentinized mantle contribute to the formation of the oceanic crust. Based on the structural configuration of the system at the time of partial melting (Fig. 3a₄, b₄) and after a given time span of 10 Ma (Fig. 3a₅ and b₅ for the hot model and the cold model, respectively), it is possible to estimate the variation in time of the width of the oceanic lithosphere, characterized by a gabbro-basalt-bearing crust, and of the serpentinized mantle denuded before melt generation. In particular, 10 Ma after the onset of the mantle partial melting our results indicate a total basin width ranging over 360–480 km as a function of the initial thermal state of the lithosphere (hot or cold models, respectively). In both models the oceanic lithosphere extends for approximately 200 km, while the denuded serpentinized mantle covers a width of approximately 160 km in the hot model and 280 km in the cold one.

3.b. Thermo-mechanical evolution

The structural configuration discussed above can be better understood if the thermo-mechanical evolution of the system is analysed. Figure 4 shows the thermal and velocity fields of the system at different stages of the evolution for the hot and the cold models (Fig. 4a_i and b_i, respectively). During the early stages of the evolution of both models, the velocity field is controlled by the far-field traction driving a predominantly horizontal velocity pattern though the lithosphere, being the intensity of the mantle upwelling below the future ridge lower than half the intensity of the far-field for the cold model (panel b₁) and even negligible for the hot model (panel a₁). Mantle upwelling increases during the evolution (Fig. 4a₂) and reaches a magnitude comparable to or higher than that of the far-field traction only after the onset of the mantle partial melting (Fig. 4a₃). For both models, the thermal thinning of the lithosphere localizes at the future ridge position when the mantle serpentinization conditions are matched, at approximately 15 Ma (Fig. 4a₁) and 4.4 Ma (Fig. 4b₁) after the beginning of the forced extension for the hot and cold models, respectively. Thermal thinning achieves its maximum when the pressure and temperature conditions are favourable for the mantle partial melting at 36.4 Ma (Fig. 4a₃) and 22.4 Ma (Fig. 4b₃) for the hot and cold models, respectively. The thermal thinning in the two models therefore differs only in the initial stages and, once the thermal destabilization starts, the time span needed to reach the maximum thinning is approximately independent of the initial thermal state (thick red and blue lines along the time scale in Fig. 4).

The different behaviours of the cold and hot models during the initial deformation phase and their similar behaviours at longer times can be ascribed to the differences in the lithosphere strength (Fig. 2b). During the initial phase, the lower strength (in terms of the lower effective viscosities of the lithospheric mantle levels) that characterizes the hot model (solid red line in Fig. 2b) may attenuate the transmission of the far-field traction up to the future ridge position, making the lithosphere thinning a very slow process compared with the cold model.

3.c. Velocity configuration

A more detailed analysis of the surface crustal horizontal velocity (Fig. 5a, b) indicates that both models are characterized by an initial phase during which the mean surface crustal velocity is stationary and increases linearly from the future ridge outwards until it reaches, at the boundary of the model, the velocity value constrained by the far-field traction. This stationary pattern lasts a short time (approximately 5 Ma) in the cold model (Fig. 5a) and a long time (approximately 30 Ma) in the hot model (Fig. 5b), ending when the crustal break-up occurs. This initial phase is followed by a transition phase lasting approximately 9 Ma and 3 Ma for the cold and hot models, respectively, during which the far-field and upwelling flow concur to the extension rate and the mean surface crustal horizontal velocities progressively increase until they reach the value of the far-field traction velocity. At the advanced stages of evolution the surface crustal horizontal velocity remains almost constant from the ridge to the margins for both models, with the exception of a 150–200 km wide region around the ridge where the mean surface horizontal velocity of the crust is mainly controlled by the very intense mantle upwelling flow and may overcome the far-field value.

The transmission of the spreading rate towards the periphery of the domain is likely limited by the fixed far-field prescribed velocity at the boundaries of the model producing, at first sight, an artificial local shortening.

Figure 6 shows the maps of the velocity modules through the lithospheric thickness at different time steps of the evolution, for the hot (Fig. 6a_i) and cold (Fig. 6b_i) models. Until crustal break-up occurs, the linear increase from the ridge outwards observed in the mean surface crustal velocity (Fig. 5) characterizes the entire lithosphere thickness (Fig. 6a₁, a₂ for the hot model and Fig. 6b₁, b₂ for the cold model). A small increase in the velocity is evident within the mantle where a decrease in the viscosity occurs because of the serpentinization. After the crustal break-up, the strongest velocity gradients localize within 200 km around the ridge and are associated with the increase in the magnitude of the upwelling mantle flow (Fig. 4). In the proximity of the ridge, in both models the extension rate at the base of the crust increases by a factor of 2–4 since the continental break-up to oceanic spreading (compare Fig. 6a₂, b₂ to Fig. 6a₃, b₃). Further away, the velocity gradients decrease significantly until they disappear, and the entire continental lithosphere moves as a rigid plate at a rate equal to the traction velocity after the onset of the mantle partial melting. Significant velocity gradients persist within the part of the basin floored by the serpentinized mantle and at the deep lithosphere levels below the ridge, where the intensity of the asthenosphere upwelling overcomes the far-field traction velocity (compare Fig. 4a₃, b₃ to Fig. 6a₄, b₄). At the advanced stages of the evolution after the onset of the mantle partial melting, as already observed in Figure 5 for the mean surface horizontal velocity, even the lithosphere velocity overcomes the far-field traction velocity up to a maximum value of 1.4 cm a⁻¹ compared with the prescribed 1.25 cm a⁻¹.

Figure 6.

Maps of the velocity modules throughout the lithosphere at different times after the beginning of the forced extension for the hot (panels a_i) and cold (panels b_i) models.

4.a. Geological outline

The Alps and Apennines developed during the closure of the Mesozoic Tethys along two opposite subduction zones during different time intervals, during Cretaceous–Oligocene time for the Alpine system and from Eocene time to the present day for the Apennines system (e.g. Dal Piaz, Reference Dal Piaz, Beltrando, Peccerillo, Mattei, Conticelli and Doglioni2010; Handy et al. Reference Handy, Schmid, Bousquet, Kissling and Bernoulli2010; Carminati & Doglioni, Reference Carminati and Doglioni2012). Different erosion percentages and exhumation efficiencies (Carminati & Doglioni, Reference Carminati and Doglioni2012) allowed the exposure of deep structural levels in the axial zone of the Alpine nappe; however, these levels are still buried in the Apennines. In contrast to the Apennines, denudation of the axial part of the belt makes the Alps an important site where it is possible to explore continuous sections of deep and intermediate Alpine and pre-Alpine continental crust and to investigate the metamorphic and igneous effects of the lithospheric-thinning-induced high thermal regime preceding Alpine convergence.

The Alps spread out from the Gulf of Genova to the Vienna Basin; the chain was truncated during the Neogenic opening of the Ligurian–Provençal–Algerian and Tyrrhenian basins (e.g. Bozzo et al. Reference Bozzo, Campi, Capponi and Giglia1992; Séranne, Reference Seranne, Durand, Jolivet, Horvath and Seranne1999; Federico et al. Reference Federico, Spagnolo, Crispini and Capponi2009; Turco et al. Reference Turco, Macchiavelli, Mazzoli, Schettino and Pierantoni2012). Alpine subduction-collision is responsible for the distribution of the continental pre-Alpine and Mesozoic oceanic rocks in four tectonic domains, individuated and explored at the lithospheric scale after the seismic integrated geophysical prospection projects of the whole belt (e.g. Polino, Dal Piaz & Gosso, Reference Polino, Dal Piaz and Gosso1990; Schmid et al. Reference Schmid, Fügenschuh, Kissling and Schuster2004; Cassinis, Reference Cassinis2006). From the internal to the external part of the belt, they are the Southalpine, Austroalpine, Penninic and Helvetic domains (Fig. 1).

The Southalpine domain consists of a south-verging thrust system that has been active since Cretaceous time, involving Palaeozoic continental basement and Permian–Cenozoic cover units, both only locally affected by very-low-grade Alpine metamorphism. The Austroalpine and Penninic domains have been deeply involved in the Alpine subduction and collision system, as accounted for by the high-pressure metamorphic imprints associated with the dominant Alpine fabrics. The Penninic domain consists of mingled crustal slices deriving from both pre-Alpine continental and Mesozoic oceanic lithosphere, the latter tectonically sampled from the subducted Tethys Ocean (e.g. Platt Reference Platt1986; Polino, Dal Piaz & Gosso, Reference Polino, Dal Piaz and Gosso1990; Stöckhert & Gerya, Reference Stöckhert and Gerya2005; Malatesta et al. Reference Malatesta, Crispini, Federico, Capponi and Scambelluri2012; Roda, Spalla & Marotta, Reference Roda, Spalla and Marotta2012; Malatesta et al. Reference Malatesta, Gerya, Crispini, Federico and Capponi2013). In contrast, the Austroalpine domain does not contain Mesozoic ophiolites but is infolded within them and the related Mesozoic sediments all along its external boundary. Finally, the Helvetic domain consists of a Europe-verging thrust system that includes basement and cover slices structured during the late stages of the Alpine continental collision since Tertiary. Due to its shallow-level Alpine tectonic history, the Helvetic and Southalpine units (Fig. 1) broadly preserve pre-Alpine metamorphic, structural and stratigraphic imprints (Fig. 7), whereas Cretaceous–Paleocene Alpine high-pressure rocks are confined within the axial part of the chain in a rootless crustal prism consisting of Penninic and Austroalpine units. These latter are bounded by the Penninic frontal thrust (PF, Fig. 1) towards the European foreland and by the Periadriatic lineament (PL, Fig. 1) towards the Adriatic hinterland (Handy & Oberhänsli, Reference Handy and Oberhänsli2004; Thöni et al. Reference Thöni, Miller, Blichert-Toft, Whitehouse, Konzett and Zanetti2008; Roda, Spalla & Marotta, Reference Roda, Spalla and Marotta2012). According to many authors the wide range of radiometric ages of high-pressure metamorphic imprints (Bousquet et al. Reference Bousquet, Engi, Gosso, Oberhänsli, Berger, Spalla, Zucali and Goffè2004; Goffé et al. Reference Goffé, Schwartz, Lardeaux and Bousquet2004; Handy et al. Reference Handy, Schmid, Bousquet, Kissling and Bernoulli2010; Lardeaux, Reference Lardeaux2014) suggests that they were buried and widely exhumed during subduction of the oceanic lithosphere (European lower plate), accompanied by tectonic erosion of the upper continental plate (Adria) before the onset of continental collision (e.g. Platt Reference Platt1986; Polino, Dal Piaz & Gosso, Reference Polino, Dal Piaz and Gosso1990; Spalla et al. Reference Spalla, Lardeaux, Dal Piaz, Gosso and Messiga1996; Gerya & Stöckhert Reference Gerya and Stöckhert2005; Roda, Marotta & Spalla, Reference Roda, Marotta and Spalla2010; Roda, Spalla & Marotta, Reference Roda, Spalla and Marotta2012; Rubatto et al. Reference Rubatto, Regis, Hermann, Boston, Engi, Beltrando and McAlpine2011). Here, pre-Alpine structural, metamorphic and igneous relics are also preserved even if pervasive Alpine structural and metamorphic reworking shapes them into small-sized and scattered lenses, inhibiting the correlation of pre-Alpine structures at the regional scale.

Figure 7.

Timing of the metamorphic and magmatic events from Variscan to Jurassic time along the Alpine belt. Radiometric ages are plotted including analytical uncertainty intervals (dashed horizontal bars). Variscan evolutions are synthesized by data in Table 2. Permian–Triassic metamorphic and igneous data (continental gabbros) are synthesized by data in Table 3 and represented with their error margin. Ophiolites and mantle age data are listed in Tables 4 and 5. The grey stripe indicates the age of the earliest radiolarian cherts related to Alpine ophiolites (Cordey & Bailly, Reference Cordey and Bailly2007).

4.b. Lithology, structures, P–T conditions and ages

Asthenospheric upwelling associated with lithospheric thinning causes high thermal regimes in the thinned continental lithosphere (e.g. Thompson, Reference Thompson1981; England & Thompson, Reference England and Thompson1984; Thompson & England, Reference Thompson and England1984; Sandiford & Powell, Reference Sandiford and Powell1986; Spear & Peacock, Reference Spear and Peacock1989; Beardsmore & Cull, Reference Beardsmore and Cull2001), and in the Alps the only igneous and metamorphic effects indicating the occurrence of such a thermal state before the Alpine convergence are of Permian–Triassic age. They are detectable along the whole belt, from the Ligurian Sea to the Pannonian Basin, even in domains strongly reworked by the Alpine tectonics and metamorphism. These records consist of a widespread emplacement of Permian–Triassic basic to acidic igneous activity and huge gabbro bodies (Tables 2 and 3; Fig. 1) associated with regional high-temperature – low-pressure (HT-LP) metamorphism, which postdate structures and metamorphic imprints widely developed during the Variscan subduction and collision (Fig. 7). These features are frequently associated with subcontinental peridotites (Table 4) and are mainly confined to the Austroalpine and Southalpine domains (e.g. Lardeaux & Spalla, Reference Lardeaux and Spalla1991; Bonin et al. Reference Bonin, Brändlein, Bussy, Desmons, Eggenberger, Finger, Graf, Marro, Mercolli, Oberhänsli, Ploquin, Quadt, Raumer, Schaltegger, Steyrer, Visonà, von Raumer and Neubauer1993; Bussy et al. Reference Bussy, Venturini, Hunziker and Martinotti1998; Rottura et al. Reference Rottura, Bargossi, Caggianelli, Del Moro, Visonà and Tranne1998; Schuster et al. Reference Schuster, Scharbert, Abart and Frank2001; Stahle et al. Reference Stahle, Frenzel, Hess, Saupe, Schmidt and Schneider2001; Rebay & Spalla, Reference Rebay and Spalla2001; Rampone, Reference Rampone2002; Peressini et al. Reference Peressini, Quick, Sinigoi, Hofmann and Fanning2007; Marotta, Spalla & Gosso, Reference Marotta, Spalla, Gosso, Ring and Wernicke2009; Spalla et al. Reference Spalla, Zanoni, Marotta, Rebay, Roda, Zucali, Gosso, Schulmann, Martínez Catalán, Lardeaux, Janoušek and Oggiano2014).

Table 3.

Permian–Triassic gabbros emplaced in the pre-Alpine continental crust of the Alps and Apennines. Labels listed in the code column are reported in Figures 1 and 9–11 to indicate rock positions and duration of thermal fitting

Table 4.

Subcontinental and oceanic mantle peridotites from the Alps and Apennines. Labels listed in the code column are reported in Figures 1 and 9–11 to indicate rock positions and duration of thermal fitting

Metamorphic Permian–Triassic imprints are widely recorded in the lower, intermediate and upper continental crust of the Austroalpine and Southalpine domains; only a few records have been recognized in the upper and intermediate Penninic crust of Western Alps, and they are never detected in the Helvetic domain (Fig. 1). In the Penninic domain, HT assemblages mainly developed in sillimanite-bearing metapelites and metaintrusives. They are generally derived from re-equilibration under low-pressure conditions and interpreted as Permian in age (Bouffette, Lardeaux & Caron, Reference Bouffette, Lardeaux and Caron1993; Table 2). Sapphirine-bearing HT- to UHT-IP granulites of the Gruf Complex have been recently petrologically and chronologically investigated, revealing an age of 260–290 Ma for T _max conditions (Galli et al. Reference Galli, Le Bayon, Schmidt, Burg, Caddick and Reusser2011). The exhumation of Penninic HT Permian–Triassic rocks was generally associated both with cooling and heating (Desmons, Reference Desmons1992; Bouffette, Lardeaux & Caron, Reference Bouffette, Lardeaux and Caron1993) and occurred before the Alpine convergence, whereas the exhumation of UHT Gruf granulites occurred at the end of the Alpine convergence from the base of the internal European passive margin where they lay since the Permian lithospheric thinning (Galli et al. Reference Galli, Le Bayon, Schmidt, Burg, Caddick and Reusser2011). In the Austroalpine domain, Permian–Triassic HT metamorphism mainly developed in sillimanite- and biotite-bearing gneisses; it is associated with minor mafic granulites, amphibolites and high-grade marbles (Table 2). HT minerals locally mark mylonitic fabrics within discrete shear zones (Lardeaux & Spalla Reference Lardeaux and Spalla1991; Spalla et al. Reference Spalla, Lardeaux, Dal Piaz and Gosso1991), and exhumation paths can be characterized by cooling, heating or isothermal decompression (e.g. Dal Piaz, Lombardo & Gosso, Reference Dal Piaz, Lombardo and Gosso1983; Stöckhert, Reference Stöckhert1987; Vuichard, Reference Vuichard1987; Lardeaux & Spalla, Reference Lardeaux and Spalla1991; Spalla, Messiga & Gosso, Reference Spalla, Messiga and Gosso1995; Schuster et al. Reference Schuster, Scharbert, Abart and Frank2001; Manzotti & Zucali, Reference Manzotti and Zucali2013). Locally, the exhumation was accomplished up to very shallow structural levels (e.g. Rebay & Spalla, Reference Rebay and Spalla2001), suggesting that some Austroalpine units belonged to a thinned continental crust before being subducted during Cretaceous convergence. In the Southalpine basement, Permian–Triassic HT metamorphism developed in metapelites, mafic granulites, amphibolites and high-grade marbles and mainly re-equilibrated under granulite-amphibolite-facies conditions (Table 2). HT paragenesis marks a pervasive foliation that is locally mylonitic within up to kilometre-thick discrete shear zones that are often steepened by Alpine thrusting (e.g. Bertotti et al. Reference Bertotti, Picotti, Bernoulli and Castellarin1993; Gosso, Siletto & Spalla, Reference Gosso, Siletto and Spalla1997). Mylonitic belts developed under upper amphibolite-granulite- to greenschist-facies conditions are widespread in the central Southalpine domain, accounting for regional-scale pervasive extensional tectonics (Brodie, Rex & Rutter, Reference Brodie, Rex, Rutter, Coward, Dietrich and Park1989; Diella, Spalla & Tunesi, Reference Diella, Spalla and Tunesi1992; Bertotti et al. Reference Bertotti, Picotti, Bernoulli and Castellarin1993) and widely interpreted as related to regional-scale normal faults responsible for the exhumation of HT-LP complexes (e.g. Brodie, Rex & Rutter, Reference Brodie, Rex, Rutter, Coward, Dietrich and Park1989; Handy et al. Reference Handy, Franz, Heller, Janott and Zurbriggen1999) during a continuous evolution from Permian to Triassic or Jurassic time (e.g. Bertotti et al. Reference Bertotti, Picotti, Bernoulli and Castellarin1993). Intrusive stocks emplaced at shallow levels during Permian – Early Triassic time are associated with metamorphic aureoles (Povoden, Horacek & Abart, Reference Povoden, Horacek and Abart2002; Benciolini et al. Reference Benciolini, Poli, Visona and Zanferrari2006; Gallien, Abart & Wyhlidal, Reference Gallien, Abart and Wyhlidal2007). Exhumation paths were characterized by cooling or by increasing temperature during decompression (e.g. Brodie, Rex & Rutter, Reference Brodie, Rex, Rutter, Coward, Dietrich and Park1989; di Paola & Spalla Reference di Paola and Spalla2000) and were generally accomplished under a high thermal regime. HT Permian–Triassic mineral associations have never been detected in the pebbles of Permian conglomerates from the Orobic Alps, suggesting that HT rocks were not yet exposed in their Variscan source areas before the late Permian – Triassic period (Spalla et al. Reference Spalla, Zanoni, Gosso and Zucali2009; Zanoni, Spalla & Gosso, Reference Zanoni, Spalla and Gosso2010).

Permian–Triassic continental gabbros have been detected in the Austroalpine and Southalpine domains of the Alps (Bonin et al. Reference Bonin, Brändlein, Bussy, Desmons, Eggenberger, Finger, Graf, Marro, Mercolli, Oberhänsli, Ploquin, Quadt, Raumer, Schaltegger, Steyrer, Visonà, von Raumer and Neubauer1993; Rottura et al. Reference Rottura, Bargossi, Caggianelli, Del Moro, Visonà and Tranne1998; Stahle et al. Reference Stahle, Frenzel, Hess, Saupe, Schmidt and Schneider2001; Spiess et al. Reference Spiess, Cesare, Mazzoli, Sassi and Sassi2010; Spalla et al. Reference Spalla, Zanoni, Marotta, Rebay, Roda, Zucali, Gosso, Schulmann, Martínez Catalán, Lardeaux, Janoušek and Oggiano2014). The mafic products of the widespread Permian–Triassic igneous activity mainly consist of gabbroic bodies with subcontinental peridotites (Brodie, Rex & Rutter, Reference Brodie, Rex, Rutter, Coward, Dietrich and Park1989; Bonin et al. Reference Bonin, Brändlein, Bussy, Desmons, Eggenberger, Finger, Graf, Marro, Mercolli, Oberhänsli, Ploquin, Quadt, Raumer, Schaltegger, Steyrer, Visonà, von Raumer and Neubauer1993; Schuster et al. Reference Schuster, Scharbert, Abart and Frank2001; Stahle et al. Reference Stahle, Frenzel, Hess, Saupe, Schmidt and Schneider2001; Rampone Reference Rampone2002; Spalla et al. Reference Spalla, Zanoni, Marotta, Rebay, Roda, Zucali, Gosso, Schulmann, Martínez Catalán, Lardeaux, Janoušek and Oggiano2014). They occurred at different structural levels, and the country rocks vary from high-temperature – intermediate-pressure metamorphics to Triassic carbonatic sediments (Sills Reference Sills1984; Handy & Zingg Reference Handy and Zingg1991; Lardeaux & Spalla Reference Lardeaux and Spalla1991; Gallien, Abart & Wyhlidal, Reference Gallien, Abart and Wyhlidal2007; Miller et al. Reference Miller, Thöni, Goessler and Tessadri2011; Spalla et al. Reference Spalla, Zanoni, Marotta, Rebay, Roda, Zucali, Gosso, Schulmann, Martínez Catalán, Lardeaux, Janoušek and Oggiano2014). From a geochemical point of view most of the gabbros have a tholeiitic to alkaline signature, although they are generally considered to be generated from variably contaminated mantle sources in an extensional tectonic regime under a high thermal state associated with lithospheric thinning and rifting (Spalla et al. Reference Spalla, Zanoni, Marotta, Rebay, Roda, Zucali, Gosso, Schulmann, Martínez Catalán, Lardeaux, Janoušek and Oggiano2014 and references in Table 3). The ages of these gabbroic intrusions in the Austroalpine domain cluster around Permian (Table 3). The main Triassic magmatic signal is recorded in the Southalpine domain by Predazzo and Monzoni intrusives and in the Ivrea Zone by alkaline rocks (Table 3). Stahle et al. (Reference Stahle, Frenzel, Hess, Saupe, Schmidt and Schneider2001) interpret this igneous activity as the result of a long-lasting process of active rifting.

Few Permian–Triassic continental mantle slices have been found in the Alps (Table 4). In particular, the North Lanzo body in the Penninic domain is characterized by subcontinental lithospheric mantle protoliths and underwent progressive exhumation during pre-oceanic lithosphere extension and rifting. The South Lanzo body shows impregnation by mid-ocean-ridge basalt (MORB) melts rising from the underlying molten asthenosphere during the rifting stage of the Liguria–Piemonte Ocean (Piccardo & Guarnieri Reference Piccardo and Guarnieri2010). The Lanzo peridotites would therefore represent a lithosphere changing from a thinned continental plate to an ocean–continent transition zone (OCTZ) (Piccardo & Guarnieri Reference Piccardo and Guarnieri2010). In the Austroalpine domain, a small peridotite body has been detected in the Dent–Blanche Nappe (Nicot, Reference Nicot1977). Although no radiometric age is available for this mantle rock, its structural relation with continental rocks of the Valpelline Series and the analogy with similar rocks of the Ivrea Zone suggest a Permian age and therefore a continental affinity for this peridotite.

In the Alps, the transition from rifting to oceanic spreading is indicated by the deposition of post-rift sediments (172–165 Ma, Baumgartner et al. Reference Baumgartner, Bartolini, Carter, Conti, Cortese, Danelian, De Wever, Dumitrica, Dumitrica-Jud, Gorican, Guex, Hull, Kito, Marcucci, Matsuoka, Murchey, O'Dogherty, Savary, Vishnevskaya, Widz and Yao1995; Stampfli et al. Reference Stampfli, Mosar, Marquer, Marchant, Baudin and Borel1998; Bill et al. Reference Bill, O'Dogherty, Guex, Baumgartner and Masson2001; Handy et al. Reference Handy, Schmid, Bousquet, Kissling and Bernoulli2010) postdating syn-rift Triassic deposits (e.g. Gillcrist, Coward & Mugnier, Reference Gillcrist, Coward and Mugnier1987). This transition involved the exhumation and serpentinization of the subcontinental mantle at the Liguria–Piemonte Ocean margins (e.g. Desmurs, Manatschal & Bernoulli, Reference Desmurs, Manatschal, Bernoulli, Wilson, Whitmarsh, Taylor and Froitzheim2001; Manatschal, Reference Manatschal2004; Manatschal & Müntener, Reference Manatschal and Müntener2009). The radiometric ages of the ophiolitic gabbros (Fig. 7, Table 5) clustering around approximately 160 Ma (Mevel, Caby & Kienast, Reference Mevel, Caby and Kienast1978; Li et al. Reference Li, Faure, Lin and Manatschal2013), with older values of 166–183 Ma from the Apennines, Corsica and Erro–Tobbio ophiolitic units (e.g. Tribuzio, Thirwall & Vannucci, Reference Tribuzio, Thirwall and Vannucci2004; Rampone et al. Reference Rampone, Borghini, Romairone, Abouchami, Class and Goldstein2014; Li et al. Reference Li, Faure, Rossi, Lin and Lahondère2015), confine the beginning of the spreading of the Liguria–Piemonte Ocean. Ages of 198 ± 22 (Sm–Nd on gabbro WR) have been obtained by Costa & Caby (Reference Costa and Caby2001) in ophiolites from the Western Alps (Chenaillet) and interpreted by the authors as the signature of lithospheric extension announcing the oceanic spreading. Oceanic gabbros exclusively occur in the ophiolitic sequences of the Penninic domain (Fig. 1). They generally have a MORB affinity and are usually associated with serpentinized mantle but sometimes to volcanic sequences, such as lava flows, pillow basalts and pillow breccias, and oceanic sediments (e.g. Mevel, Caby & Kienast, Reference Mevel, Caby and Kienast1978; Ohnenstetter et al. Reference Ohnenstetter, Ohnenstetter, Vidal, Cornichet, Hermitte and Mace1981; Riccardi, Tribuzio & Caucia, Reference Riccardi, Tribuzio and Caucia1994; Martin, Tartarotti & Dal Piaz Giorgio, Reference Martin, Tartarotti and Dal Piaz Giorgio1994).

Table 5.

Ophiolitic gabbros from the Alps, Apennines and Corsica. The lack of P–T estimates for Corsica gabbros inhibits their comparison with the model predictions. Labels listed in the code column are reported in Figures 1 and 9–11 to indicate rock positions and duration of thermal fitting

Oceanic mantle rocks coming from the Penninic domain (Table 4) are variably serpentinized peridotites. Based on the relict texture, mineralogy and structural relations with gabbros and rodingites, the serpentinized peridotite of Mt Avic (Table 4) is considered to have an oceanic affinity (Fontana, Panseri & Tartarotti, Reference Fontana, Panseri and Tartarotti2008). Although the gabbroic bodies of the Lanzo Massif are considered to have originated during the opening of the Jurassic–Piedmont Ligurian ocean (Lagabrielle, Fudral & Kienast, Reference Lagabrielle, Fudral and Kienast1989; Pognante, Rösli & Toscani, Reference Pognante, Rösli and Toscani1985) or during the earliest stages of the formation of the embryonic oceanic crust (Kaczmarek, Müntener & Rubatto, Reference Kaczmarek, Müntener and Rubatto2008), the associated peridotites show a more complex affinity as already discussed. The Mg-rich gabbroic rocks of Erro-Tobbio complex in Voltri Massif (Ligurian Alps) is considered as representative of syn-rift melt intrusions in thinned lithospheric mantle exhumed at ocean–continent transition domains (Rampone et al. Reference Rampone, Borghini, Romairone, Abouchami, Class and Goldstein2014). They represent the oldest gabbroic bodies of the Alpine Tethys (201–163 Ma, Rampone et al. Reference Rampone, Borghini, Romairone, Abouchami, Class and Goldstein2014). Because rocks from the Alps are widely deformed and metamorphosed during Alpine subduction and collision, we add ophiolites from the Northern Apennines that escaped pervasive subduction-related metamorphic re-equilibration and the unique case of mafic granulite from the continental crust to the collection of natural data (Tables 3–5). The Northern Apennines are characterized by oceanic and continental units (Fig. 1). Oceanic units are divided into two different groups of thrust nappes – the Internal and External Ligurian units (e.g. Marroni & Pandolfi, Reference Marroni and Pandolfi2007) – strongly deformed under low-grade metamorphic conditions (Marroni & Pandolfi, Reference Marroni and Pandolfi2007; Donatio, Marroni & Rocchi, Reference Donatio, Marroni and Rocchi2013). An ophiolitic sequence of Jurassic age and a sedimentary cover ranging in age from Late Jurassic to Paleocene characterize the Internal Ligurian units (Marroni & Pandolfi, Reference Marroni and Pandolfi2007). In the External Ligurian units, Late Cretaceous sedimentary melanges containing slide-blocks of ophiolites occur at the base of the Upper Cretaceous carbonate turbidites (Helminthoid Flysch; Marroni & Pandolfi, Reference Marroni and Pandolfi2007). The Ligurian units were thrust onto the Tuscan nappes during the Oligo-Miocene post-collisional convergence. The successions of the Tuscan units belong to the Adria passive continental margin, recording in sequence rifting-, cooling- and subsidence-related imprints during the opening of the Liguria–Piemonte Ocean (Marroni & Pandolfi, Reference Marroni and Pandolfi2007). Underlying Tertiary metamorphic units are exposed in rare tectonic windows (Fig. 1).

Granulitic gabbros of the External Liguride Unit associated with felsic granulites locally intrude mantle peridotites (Marroni & Tribuzio, Reference Marroni and Tribuzio1996; Marroni et al. Reference Marroni, Molli, Montanini and Tribuzio1998). The emplacement of the gabbroic protoliths occurs at deep crustal levels of late Carboniferous – early Permian age (approximately 290 Ma), and their country rocks are felsic granulites of the lower continental crust. Most likely in association with the subcontinental mantle, mafic and felsic granulites underwent a multistage exhumation beginning during Permian–Triassic time and ending during Late Triassic – Middle Jurassic time, when they were finally exhumed to shallow levels by extensive brittle faulting (Marroni et al. Reference Marroni, Molli, Montanini and Tribuzio1998). The External Liguride Unit is interpreted as an ocean–continent transition zone (Marroni et al. Reference Marroni, Molli, Montanini and Tribuzio1998).

Some Jurassic oceanic gabbros have been detected in the Apennines (Table 5; Riccardi, Tribuzio & Caucia, Reference Riccardi, Tribuzio and Caucia1994; Tribuzio, Riccardi & Ottolini, Reference Tribuzio, Riccardi and Ottolini1995; Tribuzio, Riccardi & Messiga, Reference Tribuzio, Riccardi and Messiga1997; Rebay, Riccardi & Spalla, Reference Rebay, Riccardi and Spalla2015). The oldest gabbro bodies (169–179 Ma) belong to the External Liguride Unit in the Northern Apennines. Despite their N-MORB affinity, Tribuzio, Thirwall & Vannucci (Reference Tribuzio, Thirwall and Vannucci2004) interpreted these gabbros as having developed during an intermediate stage of the rifting process that led to the opening of the Ligurian Tethys and not strictly related to the oceanic spreading.

The peridotite of the External Liguride Unit has been interpreted as the mantle of an ocean–continent transition zone (Marroni et al. Reference Marroni, Molli, Montanini and Tribuzio1998), while the serpentinized peridotite of the Internal Liguride Unit is suggested to be representative of the Jurassic oceanic lithosphere of the Liguria–Piemonte Ocean (Marroni & Pandolfi Reference Marroni and Pandolfi2007; Donatio, Marroni & Rocchi, Reference Donatio, Marroni and Rocchi2013).