Interaction of seismic sources in the Apennine belt

Physics and Chemistry of the Earth 63 (2013) 25–35

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Interaction of seismic sources in the Apennine belt Marcello Viti ⇑, Enzo Mantovani, Nicola Cenni, Andrea Vannucchi Dipartimento di Scienze Fisiche, della Terra e dell’Ambiente – Università di Siena, Via Laterina 8, 53100 Siena, Italy

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Article history: Available online 30 March 2013 Keywords: Post-seismic relaxation Seismotectonics Apennines Central mediterranean

a b s t r a c t By analyzing the historical seismicity of the Apennines since AD 1000, we have identified four major seismic sequences occurred from 1349 to 1353, 1456 to 1461, 1688 to 1706 and 1915 to 1920. Each of these crises is characterized by several strong earthquakes and by the fact that seismic activity progressively migrated from the Southern to northern Apennines. In order to provide a plausible explanation for this phenomenology, we have elaborated a detailed seismotectonic model of the Apennines, compatible with the implications of plate tectonics in the central Mediterranean region. In our model, the seismic motion of a tectonic block, induced by tectonic forces, stresses the surrounding blocks eventually causing further earthquake activity. The temporal delay between the triggering shock and subsequent strong earthquakes depends on the rheological properties of the crust–mantle system, that control post-seismic strain and stress perturbations. We argue that the spatio-temporal distribution of major earthquakes which have occurred in the Apennine belt during the most intense seismic crises since AD 1300 is consistent with the short-term implications of the ongoing tectonic setting in the central Mediterranean area and with the expected effects of post-seismic relaxation, quantified by numerical experiments based on the stress diffusion model. The results obtained point out an agreement between earthquake occurrence and the arrival of the predicted maximum amplitude of post-seismic strain and strain rate perturbation in the relevant seismic zones. Moreover, in most cases the strain regime induced by post-seismic perturbation may be compatible with the geometry and kinematics of fault systems recognized in the zones considered. These results might delineate an important tool for recognizing the zones most prone to next strong earthquakes in the study area. Ó 2013 Elsevier Ltd. All rights reserved.

1. Introduction It is widely recognized that large earthquakes cause timedependent deformation and redistribution of strain and stress in the surrounding regions, a phenomenon usually defined as postseismic relaxation (e.g., Feigl and Thatcher, 2006 and references therein). Post-seismic relaxation following strong shocks has been detected by means of space geodesy, radar interferometry and gravity measurements (e.g., Pollitz et al., 2006; Panet et al., 2007; Ryder et al., 2007; Ergintav et al., 2009; Ozawa et al., 2011). Moreover, several works suggest a causal link between the diffusion of post-seismic strain perturbation and earthquake triggering. Long-range (hundreds of km) and long-term (years) interactions of seismic sources have been recognized by the analysis of the spatiotemporal distribution of earthquakes in various seismic zones of the world (e.g., Anderson, 1975; Rydelek and Sacks, 1990, 2001, 2003; Pollitz and Sacks, 1995; Pollitz et al., 1998, 2004; Mikumo ⇑ Corresponding author. Tel.: +39 0577233826; fax: +39 0577233938. E-mail addresses: [email protected], [email protected] (M. Viti), enzo.mantovani @unisi.it (E. Mantovani), [email protected] (N. Cenni), [email protected] (A. Vannucchi). 1474-7065/$ - see front matter Ó 2013 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.pce.2013.03.005

et al., 2002; Marsan and Bean, 2003; Freed et al., 2007; Lay et al., 2009). The occurrence of a major earthquake at a sector of a plate border triggers a stress and strain perturbation that may significantly increase the probability of earthquakes in the zones where the induced strain change is most favourably oriented with respect to the nature and geometry of seismogenic faults. Strain perturbations propagate at rates of the order of tens of kilometres per year, controlled by the coupling of the upper elastic layer with the underlying viscous (or viscoelastic) layer of the crust–mantle system (e.g., Rydelek and Sacks, 1990; Pollitz et al., 1998; Viti et al., 2003, 2012). This process can be quantified by numerical techniques, using the Elsasser’s (1969) stress diffusion model. Such procedure has been used to investigate the role of post-seismic relaxation in triggering seismic activity in the Central and Eastern Mediterranean (Mantovani et al., 2001, 2008, 2010, 2012; Cenni et al., 2002, 2008; Viti et al., 2003, 2012). In this work we discuss the possible occurrence of this phenomenon in the Apennine belt, by analysing the compatibility of the spatio-temporal distribution of major shocks with the mechanical implications of the present tectonic setting and with the effects of post-seismic relaxation induced by strong earthquakes of the analysed seismic sequences.

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2. Tectonic setting of the Apennine belt A very accurate analysis of post-early Pleistocene deformation in the central Mediterranean area suggests that the most recent tectonic activity in the Apennines has mainly been determined by the fact that the outer part of the belt, stressed by the Adriatic plate, has undergone a belt parallel shortening, mainly accommodated by lateral extrusion and uplift of orogenic wedges (e.g., Viti et al., 2006; Mantovani et al., 2009, 2010; Viti et al., 2011). Such mobile belt is constituted by the Molise-Sannio (MS) wedge in the southern Apennines, the eastern sector of the Lazio-Abruzzi carbonate platform (ELA) in the central Apennines, and the Romagna–Marche–Umbria (RMU) and Toscana–Emilia (TE) wedges, in the northern Apennines (Fig. 1). In the southern Apennines, the decoupling of the MS wedge from the inner belt is mostly accommodated by extensional/transtensional deformation at the NW–SE and E–W trending Irpinia, Benevento and Matese fault zones (e.g., Ascione et al., 2003, 2007; Caiazzo et al., 2006; Viti et al., 2006). In the central Apennines, the decoupling between ELA and the western, almost fixed, part of the same platform, is accommodated by two major NW–SE sinistral transtensional fault systems (L’Aquila and Fucino), as suggested by neotectonic deformation and earthquake focal mechanisms (e.g., Cello et al., 1997, 1998; Amoruso et al., 1998; Piccardi et al., 1999, 2006; Pace et al., 2002; Viti et al., 2006; Mantovani et al., 2009).

When one of the above fault systems is activated by strong earthquakes, ELA accelerates and increases belt-parallel compression on the RMU wedge. The consequent belt-parallel shortening of the northern Apennines, accommodated by the outward escape of the RMU and TE wedges, causes extensional to transtensional deformation in the axial part of the chain (e.g., Boncio and Lavecchia, 2000; Piccardi et al., 2006; Elter et al., 2012), where the extruded wedges separate from the inner fixed belt, and thrusting at the outer border of the mobile blocks, where they override the Adriatic foreland (e.g., Boncio and Bracone, 2009; Boccaletti et al., 2010). The escaping Apennine bodies only involve the sedimentary cover, decoupled from its crustal basement at seismogenic depth (of the order of 10 km) by mechanically weak lithological horizons, as evidenced by seismic surveys (e.g., Finetti et al., 2005; Mirabella et al., 2008; Cenni et al., 2012). In particular, a weak Triassic evaporitic layer (Burano formation, e.g., Martinis and Pieri, 1964) forms the base of the Meso-Cenozoic sedimentary cover of the RMU and TE wedges (e.g., Anelli et al., 1994; Ciarapica and Passeri, 2002, 2005). The overall weakness of the Burano formation is related to the presence of evaporite (anydhrite) levels among dolonstones (e.g., De Paola et al., 2008). The foliated fabric of anydhrite could largely reduce the frictional strength along fault planes cutting that layer (Collettini et al., 2009a). Moreover, hydrothermal circulation of silica-rich fluids through dolonstones may produce phyllosilicates (talc), which further lowers fault friction (e.g., Abers, 2009; Collettini et al., 2009b).

Fig. 1. Tectonic sketch of the central Mediterranean area evidencing the Apennine wedges (grey) which move faster than the inner belt (see text for comments). Solid arrows indicate the kinematics of the African and Adria plates with respect to Eurasia (Mantovani et al., 2007, 2009; Viti et al., 2011). Empty arrows indicate the kinematics of the Apennine wedges carried by Adria. (1) African continental domain. (2) Adriatic continental domain. (3) Ionian oceanic domain. (4) Apennine orogenic wedges carried by Adria. (5) Quaternary volcanism; and (6–8) Compressional, extensional and strike-slip structures. ELA, Eastern Lazio-Abruzzi carbonate platform; MS, Molise-Sannio wedge; RMU, Romagna–Marche–Umbria wedge; and TE, Toscana–Emilia wedge. Main fault systems: As, Asolano; Aq, L’Aquila; Be, Benevento; Cn, Carnia; Fu, Fucino; Fo, Forlivese (Romagna Apennines); Ga, Garfagnana; Ir, Irpinia; Lu, Lunigiana; Mt, Matese; Mu, Mugello; Ri, Riminese; UP, Upper Pescara Valley; UTV, Upper Tiber Valley; and Vt, Viterbese.

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3. Interconnection between seismic sources in the Apennine belt The tectonic interpretation described in Section 2 may help to explain the spatio-temporal distribution of major earthquakes that have occurred in the Apennine belt in the past centuries (Viti et al., 2003, 2004; Mantovani et al., 2010). In particular, we focus attention on the four most intense seismic crises since AD 1300 (Fig. 2). Information about historical seismicity comes from the official Italian earthquake catalogue (Rovida et al., 2011) and from relevant literature (e.g., Gasperini et al., 1999; Guidoboni and Comastri, 2005). 3.1. Major seismic sequences The first seismic crisis (Fig. 2a) may have started with a strong earthquake (September 9, 1349, M = 6.6) at the inner extensional border of the MS wedge. That decoupling event may have allowed such wedge to accelerate, increasing belt-parallel push on ELA (Fig. 1) and consequently shear stress at the major NW–SE fault systems of that platform. This could explain why a strong earthquake (M = 6.5) soon occurred at the L’Aquila fault system. In turn, such decoupling event may have allowed the ELA wedge to accelerate, enhancing belt parallel push on the RMU wedge in the northern Apennines, where seismic activity soon increased at the inner extensional border of this wedge (Fig. 2a), particularly in the Upper Tiber Valley (December 25, 1352, M = 6.4; January 1, 1353, with macroseismic intensity I = IX MCS). One could note that the above

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seismic sequence in the Apennine belt has been preceded by a very strong shock in the Eastern Southern Alps (January 25, 1348, Carnia, M = 7, see Fig. 3), which may have favoured the NNW ward motion of Adria and, consequently, of the Apennine wedges that it carries (Fig. 1). The second seismic crisis (Fig. 2b) might have started in the southern Apennines with three strong decoupling earthquakes that on December 5, 1456, struck the internal extensional border of the MS wedge (Benevento, M = 6.3 and Molise, M = 7.2) and ELA (Upper Pescara Valley, southeast of L’Aquila, M = 5.8). As in the previous case, the acceleration of these wedges might have emphasized the belt-parallel push on ELA and consequently shear stress at the major fault systems of this zone, where, in fact, a strong shock occurred at the L’Aquila fault system some years later (November 27, 1461, M = 6.4). The consequent increase of tectonic load on the northern Apennines, due to the acceleration of the ELA, might be responsible for the occurrence of a strong earthquake at the internal border of the RMU wedge (April 26, 1458, Upper Tiber Valley, M = 5.8). The third seismic crisis (Fig. 2c) may have started at the inner border of the MS wedge (June 5, 1688, M = 7). Other seismic decouplings at the same extensional boundary occurred on September 8, 1694 (M = 6.8) and March 14, 1702 (M = 6.5). The consequent stress increase at the inner border of the ELA might be responsible for the occurrence of two strong events (M = 6.7) on January 14 and February 2, 1703 at the L’Aquila fault system, which ruptured that platform up to the Umbria-Marche Apennines (Cello et al., 1998), and of the major event in the southernmost sector of the same fault

Fig. 2. Distribution of major earthquakes (M P 5:0) in the Apennine belt during the four most intense seismic crises (A–D) discussed in the text. Numbers inside or close to largest circles (M P 5:5) indicate the year of occurrence. Data taken from Guidoboni and Comastri (2005) and Rovida et al. (2011). Tectonic scheme and symbols as in Fig. 1. The Apennine wedges that are carried by Adria are indicated in grey.

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Fig. 3. Modelling of the post-seismic relaxation induced by the strong earthquakes occurred during the 1348–1353 seismic crisis (Fig. 2a). See Viti et al. (2003, 2012) for details about the numerical procedure. (A) Solid rectangles indicate the horizontal projections of the fault zones adopted for simulating the actual seismic sources (Morewood and Roberts, 2000; Pace et al., 2002; Galadini et al., 2005; Guidoboni and Comastri, 2005; Galli and Naso, 2009). Grey arrows indicate the direction of the adopted horizontal displacements (applied to the long sides of the rectangles), simulating the seismic slip estimated by a displacement-magnitude scaling relationship (Wells and Coppersmith, 1994). Displacement amplitudes are 1.2, 0.6, 0.5 and 0.2 m for the Carnia, Matese, L’Aquila and Viterbese shocks, respectively. The adopted model diffusivity is 200 m2 s1. The solid point identifies the epicentre of the two shocks occurred in the Upper Tiber Valley. Diverging and converging arrows near that point respectively indicate the lengthening and shortening axes of horizontal strain induced by post-seismic relaxation. (B and C) Time patterns of the overall strain and strain rate induced in the Upper Tiber Valley by the seismic sources depicted in (A). The vertical lines mark the time of the two shocks that hit the Upper Tiber Valley. The horizontal band in (B) indicates the amplitude of the strain induced by earth tides, assumed as a lower threshold for earthquake triggering.

system (November 3, 1706, Maiella zone, M = 6.8). As noted for the first seismic crisis, the mobility of the Adria plate and the connected Apennine wedges might have been favoured as well by two strong shocks that occurred at northern collisional boundary of that plate (February 2 1695, Asolano, M = 6.5; July 28, 1700, Friuli, M = 5.6). The fourth and most recent seismic sequence (Fig. 2d) may have started with an earthquake at the extensional border of the MS wedge (June 7, 1910, M = 5.9). Anyway the most important trigger in this zone was determined by a very strong shock (January 13, 1915, M = 7) at the Fucino fault systems of the inner border of ELA. The influence of this last decoupling event on the tectonic instability of the northern Apennines, with the mechanism dis-

cussed earlier, is clearly suggested by the fact that in the period following the Fucino earthquake most seismic zones lying at the inner and outer borders of the RMU and TE wedges were activated by other events with M > 5.5 (May 17 and August 16, 1916, Riminese, M = 5.9 and 6.1; April 26, 1917, Upper Tiber Valley, M = 5.9; November 10, 1918, Forlivese, M = 5.9; June 29, 1919, Mugello, M = 6.3; September 7, 1920, Lunigiana–Garfagnana, M = 6.5). Such an exceptional time concentration of major shocks in the northern Apennines (six major events in just 6 years) is quite anomalous with respect to the known seismic history of these zones. The fact that such a seismic crisis occurred just after the 1915 Fucino decoupling event, which is presumed to increase tectonic load on the northern Apennines (Viti et al., 2004) can hardly be taken as

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casual. In this regard, Viti et al. (2012) point out that the probability of random occurrence of the above seismic sequence is almost negligible (<106), if one assumes that the related sources were independent one another. The fact that the seismic response of the northern Apennines following the 1915 Fucino earthquake was much more intense than in the other cases given in Fig. 2, may be tentatively explained as follows. When a seismic slip occurs at the more external L’Aquila fault system, the decoupled ELA wedge is relatively narrow, and thus mainly stresses the outermost sector of the RMU wedge, as occurred in the first three seismic crises (Fig. 2a–c). When, instead, seismic decoupling develops at the more internal Fucino fault system, as occurred in 1915 with the Avezzano earthquake, the sector of ELA which accelerates is significantly wider and consequently stresses the whole RMU wedge, emphazising the seismic effects of such perturbation in most seismic zones of the northern Apennines. 3.2. Computation of post-seismic relaxation The mechanical implications of the proposed tectonic setting may provide plausible explanations for the relative location of major earthquakes during the four discussed seismic sequences, but cannot easily account for the time distribution of main shocks. Significant insights into this last aspect may be gained by quantifying

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the effects of post-seismic relaxation induced by strong earthquakes. As discussed in the Introduction, we assume that post-seismic relaxation may be simulated by stress diffusion through an elastic-viscous model (Viti et al., 2003, 2012). In our two-dimensional approach, the crust–mantle system in the study area (Apennine belt and surroundings) is simulated by a thin sheet, subdivided in a grid of six-nodes, isoparametric quadratic triangular finite elements, which has been generated by an automatic procedure (TWODEPEPÒ commercial software) in order to achieve a suitable triangle shape and distribution (Viti et al., 2003, 2012). Stress diffusion through the model is triggered by imposing an instantaneous displacement at the long side of a thin rectangular box, which simulates the earthquake rupture zone. This condition may allow for the finite size of the fault zone and for magnitude and direction of the coseismic slip. Thus, a typical source model is formed by the relevant source box embedded in a thin sheet simulating the Apennine belt and surroundings. The location, shape and kinematics of the seismic sources considered for the computation of post-seismic relaxation are shown in the Figs. 3–6, each referred to a specific Apennine seismic sequence. Such sources are modelled by rectangles representing the surface projections of the respective fault zones (e.g., Viti et al., 2003, 2012). Seismic slip is simulated by a displacement vector applied to the long sides of the rectangle. As in a previous

Fig. 4. Modelling of the post-seismic relaxation induced by the strong earthquakes occurred during the 1456–1458 seismic crisis (Fig. 2b). (A) Solid rectangles indicate the horizontal projections of the fault zones adopted for simulating the actual seismic sources (Gasperini et al., 1999; Guidoboni and Comastri, 2005). Displacement amplitude is 1.5, 0.35 and 0.16 m for the Molise, Benevento and Upper Pescara Valley shocks, respectively. The solid point identifies the epicentre of the shock occurred in the Upper Tiber Valley. Diverging and converging arrows near that point respectively indicate the lengthening and shortening axes of horizontal strain induced by post-seismic relaxation. (B and C) Time patterns of the overall strain and strain rate induced in the Upper Tiber Valley by the seismic sources depicted in (A). Other information as in Fig. 3.

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Fig. 5. Modelling of the post-seismic relaxation induced by the strong earthquakes occurred during the 1688–1703 seismic crisis (see Fig. 2c). (A) Solid rectangles indicate the horizontal projections of the fault zones adopted for simulating the actual seismic sources (Gasperini et al., 1999; Di Bucci et al., 2005; Galadini et al., 2005). Displacement amplitude is 1.1, 0.8, 0.7, 0.5, 0.5 and 0.1 m for the Sannio, Irpinia, Norcia, Asolano, Benevento and Friuli shocks, respectively. The solid point identifies the epicentre of the shock occurred near L’Aquila. Diverging and converging arrows respectively indicate the lengthening and shortening axes of horizontal strain induced at Norcia and L’Aquila by post-seismic relaxation. (B and C) Time patterns of the overall strain and strain rate induced at Norcia by the Sannio, Irpinia, Asolano, Friuli and Benevento seismic sources depicted in (A). (D and E) Time patterns of the overall strain and strain rate induced at L’Aquila by the Norcia seismic source depicted in (A). Other information as in Fig. 3.

analysis (Viti et al., 2012), we have applied half of the total estimated slip on both sides of the source, in opposite directions (Figs. 3–6). Since seismological data are not available for most of the historical earthquakes here considered, the geometry of the seismic sources has tentatively been reconstructed by taking into account the spatial distribution of macroseismic intensity and evidence about fault pattern near the presumed epicentral zones (e.g., Gasperini et al., 1999; Rovida et al., 2011). In particular, the average length and slip of the various sources have been inferred from the related magnitudes, taken from the official Italian seismic catalogue (CPTI11 by Rovida et al., 2011), using suitable scaling relationship (Wells and Coppersmith, 1994; Viti et al., 2012). The thin sheet, embedding the seismic sources, is subdivided in a grid of triangular finite elements, implemented in a numerical procedure in order to solve the system of partial differential equations governing stress diffusion (Viti et al., 2003, 2012). The parameter controlling the propagation of the post-seismic perturbation through the sheet is diffusivity (D, in m2 s1), which depends on the properties of the elastic-viscous model by the relation D = EHh/g. The terms E and H are the Young elastic modulus and thickness of the elastic-brittle layer, while g and h are the viscosity and thickness of the underlying viscous layer (e.g., Wang, 1995). In previous investigations of post-seismic relaxation in the central Mediterranean region (Viti et al., 2003, 2012; Cenni et al., 2008; Mantovani et al., 2008, 2012), we have explored the sensitivity of numerical experiments to diffusivity, looking for the value of D that allows the best fit of the time delay between triggering and induced earthquakes. The results obtained suggest that such condition is fulfilled by the values D = 600, 400 and 100–200 m2 s1 for the Adriatic, Ionian and Apennine domains respectively. In particular, the diffusivity range considered for the Apennines implies that the viscosity of the viscous layer (corresponding to the lower crust and/or upper mantle) may vary from g = 1018 to 1019 Pa s, in

Fig. 6. Modelling of the post-seismic relaxation induced by the strong earthquakes occurred during the 1915–1920 seismic crisis (see Fig. 2d). Solid rectangles indicate the horizontal projections of the fault zones adopted for simulating the actual seismic sources (Amoruso et al., 1998; Piccardi et al., 1999; Costa, 2003; Vannoli et al., 2004; Delle Donne et al., 2007; Brozzetti et al., 2009; Di Naccio et al., 2009; Sani et al., 2009). Displacement amplitude is 1.1 m for the Fucino shock, 0.35 m for the Mugello event, 0.26 m for the second Riminese shocks, 0.19 m for the first Riminese, Upper Tiber Valley and Forlivese earthquakes. The solid point identifies the epicentre of the Lunigiana–Garfagnana event. Other information as in Fig. 3. See Viti et al. (2012) for details.

line with most published estimates (see Viti et al., 2012 for a discussion). In the numerical experiments here presented (Figs. 3– 6), we have thus adopted the value D = 200 m2 s1, also assuming

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Fig. 7. Post-seismic relaxation computed in the sites where major earthquakes occurred in the period 1916–1920, after the strong 1915 Fucino earthquake (Fig. 6). (A) Time patterns of the strain (solid line) and strain rate (dashed line) induced in the Riminese zone by the Fucino earthquake. (B) Strain and strain rate in the Upper Tiber Valley, computed as combined effect of the Fucino and the two Riminese earthquakes. (C) Strain and strain rate in the Forlivese zone, computed as combined effect of the Fucino, Riminese and Upper Tiber Valley earthquakes. (D) Strain and strain rate in the Mugello zone, computed as combined effect of the Fucino, Riminese, Upper Tiber Valley and Forlivese earthquakes; and (E) Strain and strain rate in the Garfagnana zone, computed as combined effect of the Fucino, Riminese, Upper Tiber Valley, Forlivese and Mugello earthquakes. Other informations as in Fig. 3. See Viti et al. (2012) for details.

that the diffusivity is uniform through the thin sheet (Viti et al., 2012). In order to investigate the seismic crises described above, we compute the overall post-seismic perturbation resulting from the combined effect of a number of earthquakes occurred at different places and times. As discussed by Viti et al. (2012), each point of the model is stressed by a sum of perturbations that depends on the geometry of the adopted seismic sources, their distance from the considered point and the time intervals separating the events. Previous investigations (e.g., Pollitz et al., 1998; Viti et al., 2003, 2012) suggest that the probability of induced earthquakes is highest in correspondence with the arrival of the maximum strain rate value. This probability is also influenced by the strain amplitude, in that no significant triggering effects are expected for strain values lower than that induced by earth tides (Viti et al., 2003). We would point out that other computational approaches exist that could be adopted in order to investigate how post-seismic relaxation affects the interaction among seismic sources. For instance, the time-dependent strain field could be converted in the related stress field. Locally, for a given fault surface of known orientation, the stress tensor perturbation may be resolved in the normal (DrN) and shear (Ds) stress components. Then, a Coulomb Failure Function can be defined as DCFF = Ds + l0 DrN, where l0 is the effective friction coefficient along the fault (e.g., Steacy et al., 2005). In general, a DCFF increase would indicate that the considered fault has approached failure, although only in relative terms. The above approach has been used to investigate the co-seismic (elastic) and post-seismic (viscoelastic) effect of 10 strong earthquakes (M > 6.5) occurred along the North Anatolian Fault from 1939 to 1999 (Lorenzo-Martin et al., 2006). A main conclusion of that work is that post-seismic perturbations have played an important role in transferring stress from the eastern to western

segments of the North Anatolian Fault, influencing the progressive failure of that structure. Notwithstanding these achievements, and the potentiality of the CFF approach, we have chosen to not carry out that kind of analysis in this work. First, the results of strain analysis are directly comparable with those obtained in our previous works for various sectors of the Apennines and Calabrian Arc (Mantovani et al., 2008, 2012; Viti et al., 2003, 2012). Second, in the analysis of post-seismic strain the time delay between triggering and induced shocks is constrained by the peak of the strain rate vs. time curve, provided that the strain exceeds the earth-tide threshold. On the other hand, no similar criteria is available for the CFF approach. Third, the CFF approach needs resolving the stress tensor on welldefined fault surfaces, which would represent major seismic sources. However, this information is often not available for the Apennine chain, where the pattern of seismogenic faulting is very complex and the relationship between mapped fault traces and real seismic sources is still debated. Obviously, this problem is enhanced for historical earthquakes such as those here considered, whose characterization is almost entirely based on macroseismic data (e.g., Gasperini et al., 1999). For the first seismic crisis (Fig. 2a), we have computed the postseismic perturbation induced in the Upper Tiber Valley by the 1348 Carnia event and 1349 Matese, L’Aquila and Viterbese shocks (Fig. 3a). Modelling results show that the 1352 and 1353 Upper Tiber Valley earthquakes took place when strain amplitude was well above the earth tide threshold (Fig. 3b) and strain rate was close to the peak of the related time pattern (Fig. 3c). Moreover, the computed strain (Fig. 3a) shows that at the time of the induced events, the Upper Tiber valley was undergoing a NE–SW horizontal lengthening, fairly compatible with normal faulting recognized in that zone (e.g., Delle Donne et al., 2007; Brozzetti et al., 2009).

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Fig. 8. Strain regimes induced in the northern Apennines by the largest earthquakes occurred in the period 1915–1920, as predicted by numerical experiments (see text and Figs. 6 and 7). Divergent and convergent arrows respectively indicate the lengthening and shortening principal axes of the horizontal strain tensor (further details in Viti et al., 2012). (A) Strain regime induced by the Fucino earthquake, computed at the time of the first Riminese event. (B) Strain regime induced by the Fucino and Riminese earthquakes, computed at the time of the Upper Tiber Valley event. (C) Strain regime induced by the Fucino, Riminese and Upper Tiber Valley earthquakes, computed at the time of the Forlivese event. (D) Strain regime induced by the Fucino, Riminese, Upper Tiber Valley and Forlivese earthquakes, computed at the time of the Mugello event, and (E) Strain regime induced by the Fucino, Riminese, Upper Tiber Valley, Forlivese and Mugello earthquakes, computed at the time of the Lunigiana–Garfagnana event. In each map, the grey circle encloses the computed strain regime nearest to the zone where the respective shock took place.

For the second seismic crisis (Fig. 2b), post-seismic perturbation has been estimated in the Upper Tiber Valley by modelling the 1456 Molise, Benevento and Upper Pescara Valley events (Fig. 4a). Results indicate that the 1458 Upper Tiber Valley earthquakes took place when the amplitude of the overall post-seismic strain exceeded the earth tide threshold (Fig. 4b) and the strain rate was close to its maximum value (Fig. 4c). As in the previous case, the modelled horizontal lengthening at the time of that shock (Fig. 4a) is compatible with active normal faulting recognized in the Upper Tiber Valley (e.g., Delle Donne et al., 2007; Brozzetti et al., 2009). The third seismic crisis (Fig. 2c) has been modelled by computing, at the Norcia zone, the overall post-seismic effects of the 1688 Sannio, 1694 Irpinia, 1695 Asolano, 1700 Friuli and 1702 Benevento shocks (Fig. 5a). Results indicate that the 1703 Norcia event occurred when post-seismic strain was well above the earth tide threshold (Fig. 5b) and the strain rate was fairly close to one of its peaks (Fig. 5c). One could wonder why that event did not occur in correspondence with the previous two strain rate peaks. For the first peak, the absence of any earthquake triggering can be easily imputed to the fact that at that time (around 1689) the induced strain was very low (comparable to the threshold value). For the second peak the absence of any reaction cannot be

explained, since the related strain value was well above the threshold. One can only consider that the third strain peak is related with a much higher strain value. We have also estimated the perturbation induced by the strong 1703 Norcia event at the L’Aquila zone (Fig. 5d and e), that was hit by a strong earthquake about 20 days later. In this case, the post-seismic perturbation is much stronger than for the previous event, due to the proximity of the triggering Norcia event. Anyway, it must be pointed out that the 1703 L’Aquila shock fairly well corresponds to the peak of the induced strain rate (Fig. 5e). Finally, one could remark that the strain regimes computed for the Norcia and L’Aquila zones (Fig. 5a) both show lengthening axis compatible with the normal fault system recognized in those zones (e.g., Cello et al., 1997, 1998; Piccardi et al., 2006). For the last major seismic crisis in the Apennines (Figs. 2d and 6), the computation of the overall post seismic perturbation induced by major shocks in the zones hit by subsequent events allows for the following considerations:  The two 1916 Riminese shocks just occurred when such zone was reached by the highest values of the strain rate that was induced by the 1915 Fucino earthquake (Fig. 7a). Furthermore, it can be noted that when the Riminese events took place the

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style of the induced strain (Fig. 8a) was compatible with the orientation of the transpressional faults recognized in that zone (e.g., Vannoli et al., 2004). The 1917 Upper Tiber Valley earthquake just occurred when that zone was reached by the highest values of the overall strain and strain rate induced by the combined effects of the Fucino and Riminese earthquakes (Fig. 7b). At the time of the Upper Tiber Valley event, the style of the induced strain (Fig. 8b) was compatible with the orientation of the normal faults recognized in that zone (e.g., Delle Donne et al., 2007; Brozzetti et al., 2009). The 1918 Forlivese earthquake took place several months later than the arrival of the highest value of the overall strain rate induced by the Fucino, Riminese and Upper Tiber Valley events (Fig. 7c). To explain this delay, one could tentatively remark that such strain rate peak is associated to a relatively low level of strain, just above the earth tide threshold, while the induced strain was higher when the Forlivese shock took place (Fig. 7c). Anyway, the low values of strain and strain rate that have developed in this case suggest that negative interference of signals from the three events considered occurred. The expected style of the induced strain (Fig. 8c) is compatible with the orientation of the transpressional faults that are tentatively recognized in that zone (Costa, 2003). The 1919 Mugello earthquake occurred when that zone was reached by the highest values of the overall strain and strain rate induced by the Fucino, Riminese, Upper Tiber Valley and Forlivese events (Fig. 7d). At the time, the style of the induced strain (Fig. 8d) was compatible with the kinematics of NE–SW trending active normal faults recognized in the Vicchio fault system of the Mugello basin (e.g., Sani et al., 2009). As for the Forlivese event, the presumably induced 1920 Lunigiana–Garfagnana earthquake occurred some months later than the peak of the strain rate generated by the previous events (Fucino, Riminese, Upper Tiber Valley, Forlivese and Mugello shocks, Fig. 7e). The fact that in this case, as for the Forlivese event, the induced strain is lower than that induced by the other events, could suggest similar considerations. The style of the expected strain at the time of the Lunigiana–Garfagnana earthquake (Fig. 8e) is compatible with the NW–SE normal fault system recognized in that zone (e.g., Di Naccio et al., 2009).

4. Discussion and conclusions It is widely recognized that each strong earthquake triggers a perturbation of the strain field that propagates in the surrounding regions, influencing the time space distribution of major successive events. This phenomenon could provide the basis for developing a deterministic approach for recognizing the seismic zones most prone to next strong events. However, in most cases this theoretical possibility is practically hampered by the lack of two basic elements, i.e. a detailed knowledge of the tectonic setting and the physical mechanisms by which each major event influences the occurrence of the next earthquakes in surrounding zones. In the case here considered the above difficulties are mitigated by the fact that a detailed reconstruction of the geodynamics and tectonic setting in the central Mediterranean is now available (e.g., Mantovani et al., 2009; Viti et al., 2011) and significant insights have been so far gained about the physical phenomena that control the interaction of seismic sources, interpreted as an effect of post-seismic relaxation (e.g., Viti et al., 2003, 2012; Mantovani et al., 2008, 2010, 2012). In this work, we have applied the above knowledge to the Apennine belt, in order to investigate the possible connection between tectonics and seismic activity. In particular, we argue that the spatio-temporal distribution of major historical earthquakes during

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the four main seismic crises that hit the Apennines since AD 1300 is compatible with the progressive migration of stress and strain that can be expected from the kinematics of the orogenic wedges (MS, ELA, RMU and TE) that form the outer, mobile sector of the Apennine belt. Then, we have investigated the role of postseismic strain perturbations in earthquake triggering by numerical experiments based on the stress diffusion model. Our computation procedure allows us to evaluate the overall effect of multiple triggering events. The results obtained provide considerable support to the hypothesis that the occurrence of major shocks is considerably influenced by the arrival of the highest values of the induced strain rate in the zones implied, when the strain amplitude significantly overcomes the earth tide threshold. The evidence pointed out in this work and other cited papers provides important insights into how seismic sources can be interrelated in the central Mediterranean region and how this phenomenon may be used to recognize the Italian zones most prone to next strong earthquakes. In this regard, it is worth considering that the proposed approach is based on the development through the study area of a well defined physical phenomenon (migrating perturbation of the strain field) that can be monitored by observations, in particular geodetic measures carried out by GPS networks. Previous studies (Viti et al., 2003; Cenni et al., 2012) have shown that the expected amplitude of post-seismic velocity may well exceed the accuracy of GPS measurements. If, for instance, a strong earthquake would strike the southern Apennines again, we could monitor the progressive migration of the related post-seismic perturbation, making thus much easier to predict when the value of strain and strain rate will be highest in the seismic zones of the central and northern Apennines. Acknowledgements We thank two anonymous reviewers for their constructive comments. This research was financially supported by the Regione Toscana and the Italian Civie Protection. References Abers, G.A., 2009. Slip on shallow-dipping normal faults. Geology 37, 767–768. http://dx.doi.org/10.1130/focus 082009.1. Amoruso, A., Crescentini, L., Scarpa, R., 1998. Inversion of source parameters from near- and far field observations: an application to the 1915 Fucino earthquake, central Apennines, Italy. J. Geophys. Res. 103, 29989–29999. Anderson, D.L., 1975. Accelerated plate tectonics. Science 167, 1077–1079. Anelli, L., Gorza, M., Pieri, M., Riva, M., 1994. Subsurface well data in the northern Apennines (Italy). Mem. Soc. Geol. It. 48, 461–471. Ascione, A., Cinque, A., Improta, L., Villani, F., 2003. Late quaternary faulting within the southern Apennines seismic belt: new data from Mt. Marzano area (southern Italy). Quatern. Int. 101–102, 27–41. Ascione, A., Caiazzo, C., Cinque, A., 2007. Recent faulting in southern Apennines (Italy): geomorphic evidence, spatial distribution and implications for rates of activity. Boll. Soc. Geol. Ital. 126, 293–305. Boccaletti, M., Corti, G., Martelli, L., 2010. Recent and active tectonics of the external zone of the northern Apennines (Italy). Int. J. Earth Sci. (Geologische Rundschau). http://dx.doi.org/10.1007/s00531-010-0545-y. Boncio, P., Bracone, V., 2009. Active stress from earthquake focal mechanisms along the padan-adriatic side of the northern Apennines (Italy), with considerations on stress magnitudes and pore fluid pressures. Tectonophysics 476, 180–194. Boncio, P., Lavecchia, G., 2000. A structural model for active extension in central Italy. J. Geodyn. 29, 233–244. Brozzetti, F., Boncio, P., Lavecchia, G., Pace, B., 2009. Present activity and seismogenetic potential of a low-angle normal fault system (Città di Castello, Italy): Constraints from surface geology, seismic reflection data and seismicity. Tectonophysics 463, 31–46. Caiazzo, C., Ascione, A., Cinque, A., 2006. Late tertiary-quaternary tectonics of the southern Apennines (Italy): New evidences from the Tyrrhenian slope. Tectonophysics 421, 23–51. Cello, G., Mazzoli, S., Tondi, E., Turco, E., 1997. Active tectonics in the central Apennines and possible implications for seismic hazard analysis in peninsular Italy. Tectonophysics 272, 43–68. Cello, G., Mazzoli, S., Tondi, E., 1998. The crustal fault structure responsible for the 1703 earthquake sequence of central Italy. J. Geodyn. 26, 443–460.

34

M. Viti et al. / Physics and Chemistry of the Earth 63 (2013) 25–35

Cenni, N., D’Onza, F., Viti, M., Mantovani, E., Albarello, D., Babbucci, D., 2002. Postseismic relaxation processes in the Aegean-Anatolian system: insights from space geodetic data (GPS) and geological/geophysical evidence. Bollettino Geofisica Teorica Applicata 43, 23–36. Cenni, N., Viti, M., Baldi, P., Mantovani, E., Ferrini, M., D’Intinosante, V., Babbucci, D., Albarello, D., 2008. Short-term (geodetic) and long-term (geological) velocity fields in the northern Apennines. Boll. Soc. Geol. Ital. 127, 93–104. Cenni, N., Mantovani, E., Baldi, P., Viti, M., 2012. Present kinematics of central and northern italy from continuous GPS measurements. J. Geodyn. 58, 62–72. http:// dx.doi.org/10.1016/j.jog.2012.02.004. Ciarapica, G., Passeri, L., 2002. The paleogeographic duplicity of the Apennines. Boll. Soc. Geol. It. Spec. 1, 67–75. Ciarapica, G., Passeri, L., 2005. Late triassic and early jurassic sedimentary evolution of the northern Apennines: an overview. Boll. Soc. Geol. It. 124, 189–201. Collettini, C., Niemeijer, A., Viti, C., Marone, C., 2009a. Fault zone fabric and fault weakness. Nature 462, 907–910. http://dx.doi.org/10.1038/nature08585. Collettini, C., Viti, C., Smith, S.A.F., Holdsworth, R.E., 2009b. Development of interconnected talc networks and weakening of continental low-angle normal faults. Geology 37, 567–570. http://dx.doi.org/10.1130/G25645A.1. Costa, M., 2003. The buried, apenninic arcs of the po plain and northern adriatic sea (Italy): a new model. Boll. Soc. Geol. Ital. 122, 3–23. De Paola, N., Collettini, C., Faulkner, D.R., Trippetta, F., 2008. Fault zone architecture and deformation processes within evaporitic rocks in the upper crust. Tectonics 27, TC4017. http://dx.doi.org/10.1029/2007TC002230. Delle Donne, D., Piccardi, L., Odum, J.K., Stephenson, W.J., Williams, R.A., 2007. High resolution shallow reflection seismic image and surface evidence of the upper tiber basin active faults (northern Apennines, Italy). Boll. Soc. Geol. Ital. 126, 323–331. Di Bucci, D., Massa, B., Tornaghi, M., Zuppetta, A., 2005. Structural setting of the 1688 Sannio earthquake epicentral area (southern Italy) from surface and subsurface data. J. Geodyn. 40, 294–315. Di Naccio, D., Boncio, P., Brozzetti, F., Pazzaglia, F.J., 2009. Morphotectonics of the lunigiana–garfagnana plio-quaternary grabens (northern Apennines). Rendiconti online Società Geologica Italiana 5, 70–72. Elsasser, W.M., 1969. Convection and stress propagation in the upper mantle. In: Runcorn, S.K. (Ed.), The Application of Modern Physics to the Earth and Planetary Interiors, Wiley-Interscience, New York, pp. 223–246. Elter, F.M., Elter, P., Eva, C., Eva, E., Kraus, R.K., Padovano, M., Solarino, S., 2012. An alternative model for the recent evolution of the northern-central Apennines (Italy). J. Geodyn. 54, 55–63. Ergintav, S., McClusky, S., Hearn, E., Reilinger, R., Cakmak, R., Herring, T., Ozener, H., Lenk, O., Tari, E., 2009. Seven years of post-seismic deformation following the 1999, M = 7.4 and M = 7.2, Izmit-Duzce, Turkey earthquake sequence. J. Geophys. Res. 114, B07403. http://dx.doi.org/10.1029/2008JB006021. Feigl, K.L., Thatcher, W., 2006. Geodetic observations of post-seismic transients in the context of the earthquake deformation cycle. C. R. Geosci. 338, 1012– 1028. Finetti, I.R., Boccaletti, M., Bonini, M., Del Ben, A., Pipan, M., Prizzon, A., Sani, F., 2005. Lithospheric tectono-stratigraphic setting of the ligurian sea-northern apennines-adriatic foreland from integrated CROP seismic data. In: Finetti, I.R. (Ed.), Deep Seismic Exploration of the Central Mediterranean and Italy, CROP PROJECT. Elsevier, Amsterdam, pp. 119–158. Freed, A.M., Ali, S.T., Burgmann, R., 2007. Evolution of stress in southern california for the past 200 years from coseismic, post-seismic and interseismic stress changes. Geophys. J. Int. 169, 1164–1179. Galadini, F., Poli, M.E., Zanferrari, A., 2005. Seismogenic sources potentially responsible for earthquakes with M P 6 in the Eastern Southern Alps (Thiene-Udine sector, NE Italy). Geophys. J. Int. 161, 739–762. Galli, P.A.C., Naso, J.A., 2009. Unmasking the 1349 earthquake source (southern Italy): paleoseismological and archaeoseismological indications from the Aquae Iuliae fault. J. Struct. Geol. 31, 128–149. Gasperini, P., Bernardini, G., Valensise, G., Boschi, E., 1999. Defining seismogenic sources from historical earthquake felt reports. Bull. Seismol. Soc. Am. 89, 94– 110. Guidoboni, E., Comastri, A., 2005. Catalogue of earthquakes and tsunamis in the Mediterranean area from the 11th to the 15th century. Istituto Nazionale di Geofisica e Vulcanologia, Roma, Italy, pp. 1037. Lay, T., Kanamori, H., Ammon, C.J., Hutko, A.R., Furlong, K., Rivera, L., 2009. The 2006–2007 Kuril Islands great earthquake sequence. J. Geophys. Res. 114, B11308. http://dx.doi.org/10.1029/2008JB006280. Lorenzo-Martin, F., Roth, F., Wang, R., 2006. Elastic and inelastic triggering of earthquakes in the North Anatolian Fault zone. Tectonophysics 424, 271– 289. Mantovani, E., Viti, M., Cenni, N., Albarello, D., Babbucci, D., 2001. Short and longterm deformation patterns in the aegean-anatolian systems: insights from space geodetic data (GPS). Geophys. Res. Lett. 28, 2325–2328. Mantovani, E., Viti, M., Babbucci, D., Albarello, D., 2007. Nubia-eurasia kinematics: an alternative interpretation from mediterranean and north atlantic evidence. Ann. Geophys. 50, 311–336. Mantovani, E., Viti, M., Babbucci, D., Vannnucchi, A., 2008. Long-term prediction of major earthquakes in the calabrian arc. Environ Semeiotics 1, 190–207. http:// dx.doi.org/10.3383/es.1.2.3. Mantovani, E., Babbucci, D., Tamburelli, C., Viti, M., 2009. A review on the driving mechanism of the tyrrhenian-apennines system: implications for the present seismotectonic setting in the central-northern Apennines. Tectonophysics 476, 22–40. http://dx.doi.org/10.1016/j.tecto.2008.10.032.

Mantovani, E., Viti, M., Babbucci, D., Albarello, D., Cenni, N., Vannucchi, A., 2010. Long-term earthquake triggering in the southern and northern Apennines. J. Seismolog. 14, 53–65. http://dx.doi.org/10.1007/s10950-008-9141-z. Mantovani, E., Viti, M., Babbucci, D., Cenni, N., Tamburelli, C., Vannucchi, A., 2012. Middle term prediction of earthquakes in Italy: some remarks on empirical and deterministic approaches. Bollettino Geofisica Teorica Applicata 53, 89–111. Marsan, D., Bean, C.J., 2003. Seismicity response to stress perturbations, analysed for a world-wide catalogue. Geophys. J. Int. 154, 179–195. http://dx.doi.org/ 10.1046/j.1365-246X.2003.01963.x. Martinis, B., Pieri, M., 1964. Alcune notizie sulla formazione evaporitica del Triassico superiore nell’Italia centrale e meridionale. Mem. Soc. Geol. It. 4, 649–678. Mikumo, T., Yagi, Y., Singh, S.K., Santoyo, M.A., 2002. Coseismic and post-seismic stress changes in a subducting plate: possible stress interactions between large interplate thrust and intraplate normal-faulting earthquakes. J. Geophys. Res. 107, 2023. http://dx.doi.org/10.1029/ 2001JB000446. Mirabella, F., Barchi, M., Lupattelli, A., Stucchi, E., Ciaccio, M.G., 2008. Insights on the seismogenic layer thickness from the upper crust structure of the umbriamarche Apennines (central Italy). Tectonics 27, TC1010. http://dx.doi.org/ 10.1029/2007TC002134. Morewood, N.C., Roberts, G.P., 2000. The geometry, kinematics and rates of deformation within an en echelon normal fault segment boundary, central Italy. J. Struct. Geol. 22, 1027–1047. Ozawa, S., Nishimura, T., Suito, H., Kobayashi, T., Tobita, M., Imakiire, T., 2011. Coseismic and post-seismic slip of the 2011 magnitude-9 tohoku-oki earthquake. Nature 475, 373–376. http://dx.doi.org/10.1038/nature10227. Pace, B., Boncio, P., Lavecchia, G., 2002. The 1984 Abruzzo earthquake (Italy): an example of seismogenic process controlled by interaction between differently oriented synkinematic faults. Tectonophysics 350, 237–254. Panet, I., Mikhailov, V., Diament, M., Pollitz, F., King, G., de Viron, O., Holschneider, M., Biancal, R., Lemoine, J.-M., 2007. Coseismic and post-seismic signatures of the Sumatra 2004 December and 2005 March earthquakes in GRACE satellite gravity. Geophys. J. Int.. http://dx.doi.org/10.1111/j.1365-246X.2007.03525.x. Piccardi, L., Gaudemer, Y., Tapponnier, P., Boccaletti, M., 1999. Active oblique extension in the central Apennines (Italy): evidence from the Fucino region. Geophys. J. Int. 139, 499–530. Piccardi, L., Tondi, G., Cello, G., 2006. Geo-structural evidence for active oblique estension in South-Central Italy. In: Pinter, N., Grenerczy, G., Weber, J., Stein, S., Medak, D. (Eds.), The Adria Microplate: GPS Geodesy, Tectonics and Hazard, NATO Science Series IV – Earth and Environmental Sciences. Springer, pp. 95– 108. Pollitz, F.F., Sacks, I.S., 1995. Consequences of stress changes following the 1891 Nobi earthquake, Japan. Bull. Seismol. Soc. Am. 85, 796–807. Pollitz, F.F., Burgmann, R., Romanowicz, B., 1998. Viscosity of oceanic asthenosphere inferred from remote triggering of earthquakes. Science 280, 1245–1249. Pollitz, F., Bakun, W.H., Nyst, M., 2004. A physical model for strain accumulation in the San Francisco Bay region: stress evolution since 1838. J. Geophys. Res. 109, B11408. http://dx.doi.org/10.1029/2004JB003003. Pollitz, F.F., Burgmann, R., Banerjee, P., 2006. Post-seismic relaxation following the great 2004 Sumatra–Andaman earthquake on a compressible self-gravitating Earth. Geophys. J. Int. 167, 397–420. http://dx.doi.org/10.1111/j.1365246X.2006.03018.x. Rovida, A., Camassi, R., Gasperini, P., Stucchi, M. (Eds.), 2011. CPTI11, The 2011 version of the parametric catalogue of italian earthquakes. Milano, Bologna . Rydelek, P.A., Sacks, I.S., 1990. Asthenospheric viscosity and stress diffusion: a mechanism to explain correlated earthquakes and surface deformation in NE Japan. Geophys. J. Int. 100, 39–58. Rydelek, P.A., Sacks, I.S., 2001. Migration of large earthquakes along the San Jacinto fault; stress diffusion from the 1857 Ft. Tejon earthquake Geophys. Res. Lett. 28, 3079–3082. Rydelek, P.A., Sacks, I.S., 2003. Triggering and inhibition of great Japanese earthquakes: the effect of Nobi 1891 on Tonankai 1944, Nankaido 1946 and Tokai. Earth Planet. Sci. Lett. 206, 289–296. Ryder, I., Parsons, B., Wright, T.J., Funning, G.J., 2007. Post-seismic motion following the 1997 Manyi (Tibet) earthquake: InSAR observations and modelling. Geophys. J. Int.. http://dx.doi.org/10.1111/j.1365-246X.2006.03312.x. Sani, F., Bonini, M., Piccardi, L., Vannucci, G., Delle Donne, D., Benvenuti, M., Moratti, G., Corti, G., Montanari, D., Sedda, L., Tanini, C., 2009. Late pliocene-quaternary evolution of outermost hinterland basins of the northern Apennines (Italy), and their relevance to active tectonics. Tectonophysics 476, 336–356. Steacy, S., Gomberg, J., Cocco, M., 2005. Introduction to special section: Stress transfer, earthquake triggering, and time-dependent seismic hazard. J. Geophys. Res. 110, B05S01. http://dx.doi.org/10.1029/2005JB003692. Vannoli, P., Basili, R., Valensise, G., 2004. New geomorphic evidence for anticlinal growth driven by blind-thrust faulting along the northern marche coastal belt (central Italy). J. Seismolog. 8, 297–312. Viti, M., D’Onza, F., Mantovani, E., Albarello, D., Cenni, N., 2003. Post-seismic relaxation and earthquake triggering in the southern Adriatic region. Geophys. J. Int. 153, 645–657. Viti, M., De Luca, J., Babbucci, D., Mantovani, E., Albarello, D., D’Onza, F., 2004. Driving mechanism of tectonic activity in the northern Apennines: quantitative insights from numerical modelling. Tectonics 23, TC4003. http://dx.doi.org/ 10.1029/2004TC001623. Viti, M., Mantovani, E., Babbucci, D., Tamburelli, C., 2006. Quaternary geodynamics and deformation pattern in the southern Apennines: implications for seismic activity. Boll. Soc. Geol. Ital. 125, 273–291.

M. Viti et al. / Physics and Chemistry of the Earth 63 (2013) 25–35 Viti, M., Mantovani, E., Babbucci, D., Tamburelli, C., 2011. Plate kinematics and geodynamics in the central Mediterranean. J. Geodyn. 51, 190–204. http:// dx.doi.org/10.1016/j.jog.2010.02.006. Viti, M., Mantovani, E., Cenni, N., Vannucchi, A., 2012. Post-seismic relaxation: an example of earthquake triggering in the Apennine belt (1915–1920). J. Geodyn. 61, 57–67, http://dx.doi.org/10.1016/j.jog.2012.07.002.

35

Wang, K., 1995. Coupling of tectonic loading and earthquake fault slips at subduction zones. Pure Appl. Geophys. 145, 537–559. Wells, D.L., Coppersmith, K.J., 1994. New empirical relationships among magnitude, rupture length, rupture width, rupture area and surface displacement. Bull. Seismol. Soc. Am. 84, 974–1002.