Architecture and mechanics of an active low-angle normal fault: Alto Tiberina Fault, northern Apennines, Italy

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1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 112,, doi: /2007jb005015, 2007 Architecture and mechanics of an active low-angle normal fault: Alto Tiberina Fault, northern Apennines, Italy L. Chiaraluce, 1 C. Chiarabba, 2 C. Collettini, 3 D. Piccinini, 4 and M. Cocco 4 Received 26 February 2007; revised 28 June 2007; accepted 27 July 2007; published 25 October [1] We present seismological evidence for the existence of an actively slipping low-angle normal fault (15 dip) located in the northern Apennines of Italy. During a temporary seismic experiment, we recorded 2000 earthquakes with M L 3.1.* The microseismicity defines a 500 to 1000 m thick fault zone that crosscuts the upper crust from 4 km down to 16 km depth. The fault coincides with the geometry and location of the Alto Tiberina Fault (ATF) as derived from geological observations and interpretation of depth-converted seismic reflection profiles. In the ATF hanging wall the seismicity distribution highlights minor synthetic and antithetic normal faults (4 5 km long) that sole into the detachment. The ATF-related seismicity shows a nearly constant rate of earthquake production, 3 events per day (M L 2.3), and a higher b value (1.06) with respect to the fault hanging wall (0.85) which is characterized by a higher rate of seismicity. In the ATF zone we also observe the presence of clusters of earthquakes occurring with relatively short time delays and rupturing the same fault patch. To explain movements on the ATF, oriented at high angles (75 ) to the maximum vertical principal stress, we suggest an interpretative model in which crustal extension along the fault is mostly accommodated by aseismic slip in velocity strengthening areas while microearthquakes occur in velocity weakening patches. We propose that these short-lived frictional instabilities are triggered by fluid overpressures related to the buildup of CO 2 -rich fluids as documented by boreholes in the footwall of the ATF. Citation: Chiaraluce, L., C. Chiarabba, C. Collettini, D. Piccinini, and M. Cocco (2007), Architecture and mechanics of an active low-angle normal fault: Alto Tiberina Fault, northern Apennines, Italy, J. Geophys. Res., 112,, doi: /2007jb Introduction: Low-Angle Normal Fault Enigma [2] The possibility that moderate-to-large earthquakes nucleate on low-angle normal faults (LANF, i.e., normal faults dipping less than 30 ) accommodating extension of continental crust is widely debated in the published literature [Wernicke, 1995, and reference therein]. Anderson- Byerlee frictional fault mechanics [e.g., Sibson, 1985] predict no slip on normal faults dipping less than 30 in an extensional tectonic setting characterized by a vertical s 1 if faults have a friction coefficient (m s ) ranging between 0.6 and 0.85 [Byerlee, 1978]. This should mean that mechanically it is easier to form a new fault instead of reactivating a severely misoriented low-angle structure [Sibson, 1985]. This hypothesis is supported by the observation that no 1 Istituto Nazionale di Geofisica e Vulcanologia, Centro Nazionale Terremoti, Grottaminarda Observatory, Rome, Italy. 2 Istituto Nazionale di Geofisica e Vulcanologia, Centro Nazionale Terremoti, Rome, Italy. 3 Geologia Strutturale e Geofisica, Dipartimento di Scienze della Terra, Universita degli Studi di Perugia, Perugia, Italy. 4 Dipartimento Sismologia e Tettonofisica, Istituto Nazionale di Geofisica e Vulcanologia, Rome, Italy. Copyright 2007 by the American Geophysical Union /07/2007JB005015$09.00 moderate-to-large earthquake ruptures have been documented on LANF using positively discriminated focal mechanisms [Jackson and White, 1989; Collettini and Sibson, 2001]. Therefore LANF are believed to be unimportant structures in terms of seismic hazard and in the accommodation of regionally significant amounts of crustal extension. In many cases, the low dip angles observed in the field are explained as being related to passive rotations induced by successive normal fault sets [Proffett, 1977]. [3] Potential but controversial examples of seismic activity on LANF have been proposed for moderate and large earthquakes, triggered subevents and microseismicity. Normal faulting on planes dipping at low angles have been invoked for three 5.7 < M w < 6.8 earthquakes occurring in the Papua New Guinea region [Abers, 1991; Wernicke, 1995]. Source parameters from teleseismic body waves show that the largest event (M w 6.8), dipping at 10 25, is nearly coplanar to the fault bounding the active metamorphic core complex, suggesting slip on the low-angle detachment [Abers, 1991]. Recent studies based on highresolution seismic imaging show that the dip angle of this detachment is about 30 [Floyd et al., 2001]. In general, evidence from body wave focal mechanisms is inconclusive as the uncertainty in dip angle estimates is approximately ±10 on average [Molnar and Chen, 1982] and the discrimination between rupture planes is difficult. *The M L value in line three of the abstract is correct here. The article as originally published appears online. 1of22

2 [4] Small-to-moderate triggered subevents (Dixie Valley, Nevada, 1954; Alasehir, Turkey, 1969; Gediz, Turkey, 1970) have been interpreted as low-angle ruptures that occurred minutes or seconds after nearby events on steeper faults [Axen, 1999]. A similar but controversial interpretation was proposed for the 1980 Irpinia (southern Apennines) earthquake [Pingue and De Natale, 1993; Westaway and Jackson, 1987] and different faulting models were proposed for the Irpinia subevents [Nostro et al., 1997, and references therein]. Although important for seismic hazard, these triggered ruptures are exceptions because they are commonly interpreted to result from the reactivation of preexisting faults caused by local reorientations of the active stress field and main shock-induced stress perturbations [King and Cocco, 2001]. However, the role of coseismic stress changes in modifying the orientation of the stress field is still debated and has not been ultimately proved using robust observations. [5] A two month duration microseismic survey in the Gulf of Corinth (Greece) recorded 5000 microearthquakes (1.0 < M L < 3.0) giving 148 well-constrained focal mechanisms [Rigo et al., 1996]. The hypocenters distribution, imaging a gently northward dipping plane extending from 6 to 11 km together with earthquake nodal planes dipping at in the same direction have been interpreted as evidence of a major detachment gently dipping toward the north, into which active faults mapped at the surface seem to root [Rigo et al. 1996; Rietbrock et al., 1996]. However, Hatzfeld et al. [2000], after reviewing different seismic sequences in the same area, proposed that the main faults in the area have dip angles ranging between 30 and 45 and that the geometry of the nodal planes inferred from fault plane solutions is not consistent with the geometry of the detachment. On the basis of their results they interpreted the detachment-related seismicity to represent the seismicaseismic transition. [6] In stark contrast, the geological evidence for active low-angle normal faulting appears to be overwhelming having been documented in numerous field-based structural studies [Lister and Davis, 1989; Axen, 1999; Sorel, 2000; Hayman et al., 2003; Collettini and Holdsworth, 2004] and interpretation of seismic reflection profiles [Roy and Kenneth, 1992; Barchi et al., 1998; Laigle et al., 2000; Floyd et al., 2001]. In these studies the high-angle and low-angle normal faults appear to contemporarily accommodate deformation with the steeply dipping structures generally soling into the master low-angle fault that acts as a basal detachment (we will use this definition of detachment in this study). Given the general absence of reliable evidence for the existence of moderate-to-large earthquake events along many LANF, one possible explanation is that such detachments are anomalously weak (m s 0.6) and there is some geological evidence to support this suggestion [Holdsworth, 2004; Collettini and Holdsworth, 2004]. [7] In this study we present new seismological evidence for microseismicity associated with active slip along a regional low-angle normal fault, the Alto Tiberina fault (ATF) located in the northern Apennines of Italy (Figure 1) [Barchi et al., 1998]. We first document the geometry and main features of the ATF by integrating seismic reflection profiles with seismological data collected during the seismic survey [Piccinini et al., 2003]. Then, we investigate the seismicity pattern and discuss the role of this active structure in the tectonic development and accommodation of substantial amounts of regional extension in the area. 2. Seismicity and Tectonic Setting [8] The northern Apennines consist of a NE verging thrust-fold belt formed as the result of the collision between the European continental margin (Sardinia-Corsica block) and the Adriatic lithosphere [e.g., Alvarez, 1972; Reutter et al., 1980]. From the Oligocene to the present-day, the area has experienced two phases of eastward migrating deformation: an early compression with eastward directed thrusting and a later phase of extension [e.g., Elter et al., 1975; Pauselli et al., 2006]. The interpretation of the seismic reflection profiles provided by the CROP03-NVR survey (Crosta Profonda Project Near Vertical Reflection [Pialli et al., 1998]) shows that a significant amount of extension within the brittle upper crust is accommodated by a system of east dipping LANF with associated high-angle antithetic structures (Figure 1). Older parts of the extensional system are significantly exhumed to the west in the Tyrrhenian islands (e.g., Elba) and Tuscany [Carmignani and Kligfield, 1990; Keller et al., 1994; Jolivet et al., 1998; Collettini and Holdsworth, 2004] while the ATF [Barchi et al., 1998; Boncio et al., 2000; Collettini and Barchi, 2002], which is the easternmost of these structures, is located in the inner sector of the Umbria-Marche Apennines (Figure 1b) where an extensional stress field is active today. Striated geological fault planes [Lavecchia et al., 1994], focal mechanisms and borehole break outs [Montone et al., 2004] consistently define a regional active stress field with a nearly vertical s 1 and NE trending subhorizontal s 3 with an extensional geodetic strain rate of the order of 2.5 mm/a [Hunstad et al., 2003]. [9] In the Tyrrhenian Sea and Tuscany sector (see Figure 1), the extension has been active for long enough (20 Ma) to affect the geophysical properties of the lithosphere, producing a widespread heat flow anomaly 90 mw/m 2 [Mongelli and Zito, 1991] and a thin crust (20 25 km) [Ponziani et al., 1995]. The region characterized by tectonic extension is also affected by widespread and vigorous CO 2 degassing occurrences [Chiodini et al., 2004]. Carbon mass balance calculations, relating aquifer geochemistry to dissolved inorganic carbon isotopic composition and hydrological data, have been used to demonstrate a deep mantle origin for the CO 2 in the Apennines. The eastern boundary of this degassing region corresponds closely to the investigated area of active extension. Indeed, in the San Donato and Santo Stefano boreholes, which are located in the immediate footwall of the ATF (Figures 2a and 2b), a large flux of deep-seated CO 2 is documented and high-fluid pressures are observed. Thus CO 2 at 85% of the lithostatic pressure has been encountered at depths of 4.8 km and 3.7 km, respectively. Several authors suggest that deep fluids play a key role in triggering earthquakes [Chiodini et al., 2004; Antonioli et al., 2005] and control the spatiotemporal evolution of seismicity [Miller et al., 2004; Antonioli et al., 2005]. 2of22

3 Figure 1. (a) Distribution of seismicity in the northern Apennines. Recent seismicity is distributed along the active extensional inner sector of the Apennines, where the strongest historical earthquakes are also located (white stars). Red symbols show the epicenters of the earthquakes recorded during the seismic survey (this study). Orange and blue symbols indicate the aftershocks of the 1984 Gubbio (M w 5.1) and the 1998 Gualdo Tadino (M w 5.1) earthquakes, respectively. Green symbols indicate the seismicity during the 1997 Colfiorito sequence, consisting of several normal faulting main shocks (we have plotted the focal mechanisms of the three largest shocks: M w 6.0, M w 5.7, and M w 5.6 from NW to SE, respectively). (b) Crustal-scale cross section interpretation of the CROP03 seismic profile running from the Tyrrhenian to the Adriatic coasts [Barchi et al., 1998; Collettini and Barchi, 2004]. The ATF is drawn in red, while other low-angle normal faults in the Tyrrhenian and Tuscany sectors are shown in blue. The brittle-ductile transition [Pauselli and Federico, 2002] is indicated by the light grey curve. Extension migrated with time from west to east and is accommodated by a set of east dipping low-angle normal faults and antithetic structures. Color appears in back of the print issue. [10] At present, the active extension region is concentrated in the inner zone of the Umbria-Marche Apennines where the strongest historical (intensity XI [Boschi et al., 1998]) and instrumental (5.0 < M < 6.0) earthquakes are located (Figure 1a). The seismicity does not follow the arc shape structures inherited from the previous compressional tectonic phase but clusters along a 30 km wide longitudinal zone [Chiaraluce et al., 2004; Chiarabba et al., 2005] where the historical earthquakes are also located (see Figure 1a). Furthermore the present-day extensional strain in the northern Apennines inferred from geodetic data [Hunstad et al., 2003] is concentrated across a km wide zone that coincides with the area struck by the strongest earthquakes. In the past 20 years, three main seismic sequences have occurred in the study region: the 1984 Gubbio sequence (M w 5.1), the 1997 Colfiorito multiple main shock sequence (M w 6.0, 5.7 and 5.6), and the 1998 Gualdo Tadino (M w 5.1) sequence (Figure 1a, orange, green, and blue symbols, respectively [see also Chiaraluce et al., 2004]. All the main shocks are related to SW dipping (40 ) normal faults, with fault plane ruptures dipping in the opposite direction to the ATF. Most of the recent seismic sequences in the study area show a temporal evolution of seismicity consistent with a standard main shock-aftershock behavior (see Chiaraluce et al. [2004] for a discussion on historical seismicity in the area). The Colfiorito earthquake sequence is the only one that shows a peculiar spatial and temporal evolution characterized by repeated main shocks spread over a relatively short time interval (nearly a month). This prolonged seismic activity observed for the Colfiorito sequence and the nucleation of aftershocks on the hanging 3of22

4 Figure 2. Geometry of the ATF constrained by surface geology, interpretation of seismic reflection profiles and borehole data. (a) Structural map of the study area. (b) Vertical geological cross section (see its location in Figure 2a) integrating surface geology, borehole data, and seismic reflection profiles (modified after Collettini and Barchi [2002]). (c) Geometry of the Tiber basin in a seismic profile (see its location in Figures 2a and 2b). The asymmetrical half graben shape of the basin is consistent with the presence of an east dipping normal fault. (d) ATF imaged on a seismic profile (for location, see Figures 2a and 2b). The seismic reflectors in the hanging wall block are truncated abruptly by the east dipping signals representing the ATF. Mf, Marne a Fucoidi reflector; te, top of the Triassic evaporites. Color appears in back of the print issue. wall of the main normal fault segments [Chiaraluce et al., 2003, 2004] have been interpreted in terms of fluid flow and pore pressure relaxation caused by a high fluid pressure due to CO 2 trapped at depth [Miller et al., 2004; Antonioli et al., 2005]. The presence of pressurized fluids within the crustal volume at seismogenic depths is corroborated by the observed V p and V p /V s structure of the Colfiorito normal faulting system [Ripepe et al., 2000; Chiarabba and Amato, 2003]. 3. Imaging the Alto Tiberina Fault System 3.1. Surface Geology, Seismic Reflection Profiles, and Borehole Data [11] The ATF is located in the inner region of the Umbria- Marche Apennines (Figure 1). Geologically, the region 4of22 comprises a cover sequence of continental margin sedimentary rocks deposited upon a Paleozoic metamorphic basement. The base of the cover succession is represented by late Triassic Evaporites, overlain by Liassic platform carbonates, Jurassic-Oligocene pelagic sediments and Miocene- Pliocene turbidites. Pliocene-Quaternary fluvio-lacustrine sediments fill a series of syntectonic basins formed during the extensional phase. In the Perugia Mountains area, the ATF is intersected by two boreholes (Figure 2) where it is observed to juxtapose Miocene turbidites over the Triassic Evaporites, completely omitting the intervening Umbria- Marche carbonates. Near to the surface, the fault is represented by a splay of domino-like normal faults that detach downward onto the ATF at depth (Figure 2b) [Boncio et al., 2000]. The easternmost splay of this system bounds the Tiber

5 Figure 3 5of22

6 Figure 4. Map view and vertical cross section of the seismicity (a) before and (b) after the relocation using the DD method (see Appendix A). The grey triangles represent seismic stations. Initial earthquake locations are taken from Piccinini et al. [2003]. basin, which is infilled by upper Pliocene-Quaternary sediments [Ambrosetti et al., 1978]. The asymmetrical shape of the basin is imaged clearly on seismic reflection profiles (e.g., Figure 2c) and is consistent with movements along an east dipping extensional fault. In the hanging wall block, commercial seismic reflection profiles [Barchi et al., 1998; Collettini and Barchi 2002; Mirabella et al., 2004] highlight the almost flat attitude of the seismic reflectors that abruptly terminate against the fault trace (Figure 2d). The seismic reflector corresponding to the Marne a Fucoidi formation (Mf on Figure 2d) is located at 2 km depth in the ATF hanging wall and its offset equivalent is exposed at the surface in the footwall of the basin bounding fault. The age of syntectonic sediments forming the Tiber basin and the displacement of the basin-bounding fault, define a minimum time-averaged long-term slip rate of 1 mm/a in the last 2 Ma [Collettini and Holdsworth, 2004]. [12] In the CROP03 profile [Barchi et al., 1998], the trace of the ATF is represented by the eastward dipping alignments of reflectors that can be traced down to 13 km below the Apenninic chain (Figure 3a); the same alignments emerge at the surface at the western margin of the Tiber basin. Along the ATF trace ( s two-way traveltime in Figure 3a), a lens-shaped east dipping package of reflectors appear to correspond to an extensional duplex at depth. In the hanging wall block, the previously developed compressional structures are clearly imaged and are offset by ATF-related faulting. A 5 km displacement along the fault is estimated using the observed offset of the seismic reflection corresponding to the top of the Phyllitic basement (tp in Figure 3a) [Barchi et al., 1998; Boncio et al., 2000]. [13] We have reconstructed the three dimensional geometry of the ATF reflector by depth converting the seismic reflection profiles (see location in Figures 2a and 3b) using seismic interval velocities obtained from borehole data in the area [Bally et al., 1986; Barchi et al., 1998]. The ATF was picked using 209 points along the depth-converted profiles and these points have been used to build up the three-dimensional (3-D) shape of the fault (Figure 3c). We are able to image the ATF continuously for about 60 km longitudinally. At depth, the attitude of the fault surface is almost constant with a NNW-SSE strike direction and an average dip of 15 20, while toward the surface (depth <2 km) the fault becomes slightly steeper In section 3.2 we integrate the inferred geometry of the fault with the relocated seismicity. Figure 3. (a) Geometry of the ATF along a portion of the CROP03 seismic reflection profile (for location, see Figure 2a). The fault is represented by the east dipping alignment of reflectors; tp indicates the top of the phyllitic basement. (b) Geometry of the ATF (modified after Collettini and Barchi [2002]). Dots represent the 209 points collected along the depth converted seismic reflection profiles and used for the 3-D reconstruction. (c) A 3-D image of the ATF fault. Color appears in back of the print issue. 6of22

7 Table 1. One-Dimensional P Wave Velocity Model Top of Layer, km Vp Velocity, km s Relocation of the Seismicity [14] In order to investigate independently the geometry of the ATF using background seismicity, we relocated the earthquakes collected during a seismic survey performed between October 2000 and May 2001 around the Tiber basin (Figure 1a), in the Città di Castello area [Piccinini et al., 2003]. The temporary local network was deployed in the study area during a project funded by the Italian Civil Protection [Amato et al., 2000; Cocco et al. 2000]. The main objective of the study was to provide the location and the geometry of seismogenic sources located along this active portion of the northern Apennines characterized by the lack of historical earthquakes as well as to determine possible ground shaking scenarios. A dense network of 33 three-component seismic stations was installed across the ATF system and maintained for 8 months, operating for 211 days until 15 May The network geometry shown in Figure 4 recorded nearly 2000 earthquakes with M L < 3.1 which represents the background seismicity of the area (red dots in Figure 1a). To obtain high-resolution hypocenter locations, we applied the double-difference relocation algorithm using the code hypodd [Waldhauser and Ellsworth, 2000; Waldhauser, 2001]. We used ordinary P and S wave arrival times picked from three component seismograms and high-precision differential traveltimes from waveform cross correlations. The 1-D P wave velocity model (see Table 1) used to compute partial derivatives of theoretical traveltimes and the adopted V p /V s ratio (1.84) were obtained from seismic data [Moretti, 2005] and rock velocities inferred from borehole data [Bally et al., 1986]. The initial hypocentral locations are those computed by Piccinini et al. [2003]. The details of the performed relocation procedure are set out in Appendix A. [15] In Figure 4 we show the earthquakes before and after relocation. The final catalogue consists of 1416 events with relative hypocentral location errors (see Appendix A), mainly constrained by cross-correlation data, that are a few meters between nearby events. Conversely, horizontal and vertical formal errors in hypocentral locations, which have been estimated between events predominantly constrained by phase picks, are of the order of 50 m and 70 m, respectively. These errors reflect the relative location uncertainty, while the absolute location uncertainty might be larger [see Waldhauser and Ellsworth, 2000]. We decreased the error related to the routine locations by more than an order of magnitude and the RMS residuals decrease after relocation from 114 ms to 8 ms and 34 ms for cross correlation and phase pick data, respectively. [16] Because of the relocation process it is possible to appreciate in Figure 4 both in map view and in the NE trending vertical cross section the improvement in earthquake distribution. The double-difference procedure allows us to define with a higher-resolution both the geometry of the ATF fault and the secondary structures that we will describe in section 3.3 together with the kinematics of this normal fault system Analysis of Hypocenters and Fault Plane Solutions [17] Figure 5 illustrates the distribution of the 1416 relocated earthquakes (M L < 3.1). In order to unravel the detailed geometry of the ATF system and its kinematics, we plotted in six NE trending vertical cross sections (whose position and width are shown in Figure 5a), the available focal mechanism solutions and traces of the intercept of each cross section with the geometry of the ATF as imaged by seismic reflection profiles (grey lines in the cross sections). The thickness of these heavy grey lines representing the ATF is scaled to show the nearly 400 m width of the fault zone. [18] The focal mechanisms have been computed using the first motion polarity method and to better constrain the solutions, we used takeoff angles derived from a local threedimensional velocity model [Moretti, 2005]. The black fault plane solutions are associated with earthquakes located in the hanging wall of the ATF, while grey and red solutions indicate those events that nucleated on the low-angle fault zone itself. Unfortunately, we were only able to compute one focal mechanism for the microearthquakes located on or near to the ATF surface (grey solution in cross section 6). This is due to both the small aperture of the local seismic network (see station distribution in Figure 4) and the low magnitude of the ATF-related earthquakes (M L 2.3). This solution is consistent with normal faulting and shows an east dipping low-angle plane. The two red solutions are composite focal mechanisms computed for a cluster of 12 events occurred in a small area of a 500 m radius (solution in cross section 1) and for a doublet (solution in cross section 6). We show in Figure 6 the focal sphere with the polarities for these two composite focal mechanisms. Because of the difficulties in computing reliable focal mechanisms for the ATF-related seismicity, these composite solutions are extremely useful to better constrain the geometry and kinematics of the events nucleating within the fault Figure 5. (a) Structural map of the study area. Black dots indicate the relocated seismicity (this study), while orange open circles show the aftershocks of the 1984 Gubbio (M w 5.1) earthquake [Collettini et al., 2003]. The inset shows the stress field obtained by inverting the focal mechanisms computed in this study. (b) Vertical cross section perpendicular to the Apenninic chain showing both the relocated composite seismicity in this study and the 1984 Gubbio sequence. (c) Six vertical cross sections showing the seismicity distribution and the available fault plane solutions computed for the data. Their positions are shown in Figure 5a together with the width used to plot hypocenters. The heavy grey lines plotted in each cross section represent the trace of the ATF fault as imaged on the depth-converted seismic reflection profiles (see Figure 3). Color appears in back of the print issue. 7of22

8 Figure 5 8of22

9 Figure 6. Two composite fault plane solutions computed using the first motion polarity method. (a) Solution computed for a cluster of 12 events occurring in a small area of a 500 m radius (red solution in cross section 1 of Figure 5). (b) Solution computed for a doublet (red solution in cross section 6 of Figure 5). See text for explanation. zone. Both solutions show a gently east dipping fault plane (NNW trending) consistent with the ATF geometry inferred by both geology and seismicity. [19] The two independent data sets (seismicity and seismic reflection profiles) are in excellent agreement, imaging the same regional-scale eastward dipping low-angle fault. In map view (e.g., Figure 5a) the seismicity is elongated in the Apenninic direction (NNW) for nearly 60 km parallel to the strike of the normal faults in the area, while in cross section (Figure 5b), the earthquake distribution defines a volume confined at depth by a nearly planar surface dipping E-NE at low angles (15 ). This plane separates a seismically active hanging wall block from a completely aseismic footwall (see Figure 5a). The major NNW trending structure is associated with microseismic activity from 3 4 km to 16 km depth (see Figure 5a and the six vertical cross sections). The detachment-related earthquakes and east dipping reflectors evident in the seismic profiles define a fault zone with a thickness of between 500 m (see earthquake alignments in cross sections 2 and 4) and 1 km (see the other vertical cross sections). The latter might be associated with the observed thickening of the east dipping reflectors evident in the seismic profiles (see also Figure 3a). [20] The alignment of the hypocenters in the hanging wall of the ATF highlight details of minor synthetic and antithetic faults linked at depth with the low-angle detachment. Two minor NW trending faults are especially clear: the first structure (F1) is a 5 km long, gently eastward dipping fault that soles into the detachment (cross sections 2 and 3 in Figure 5). The second structure (F2) is a high-angle west dipping fault which merges at the surface with the mapped Gubbio normal fault (cross section 4). The dip and the strike angles of the nodal planes of the earthquakes located in the ATF hanging wall are consistent with the geometry of the two minor faults imaged by hypocenter alignments (F1 and F2, see cross sections 2 3 and 4, respectively), revealing normal faulting on these high-angle faults. The P axis distributions (see Figure 7a) consistently show a peak plunge angle between 70 and 90, while the T axes (Figure 7b) have a nearly horizontal plunge and a direction trending perpendicular to the regional Apennine fault trend. We also computed the stress tensor by inverting the available fault Figure 7. Distribution of azimuth and plunge angles for (a) P and (b) T axes derived from focal mechanisms (black and open symbols). 9of22

10 Figure 8. Map view distribution of the 621 earthquakes (black open circles) located within 500 m of the ATF detachment surface represented by the isobaths (solid black lines). The red circles represent the multiplets (see text for explanation). Major grey squares labeled A and B indicate the locations of the two multiplets used to compute composite focal mechanisms shown in Figure 6. Minor grey squares labeled 1, 2, and 3 show the locations of the repeaters shown in Figure 10. Color appears in back of the print issue. plane solutions (using the method of Michael [1984]). The results are illustrated in the inset of Figure 5a and suggest a subvertical s 1 and subhorizontal, NE trending s 3 consistent with the present-day regional stress field inferred from other data [Lavecchia et al., 1994; Frepoli and Amato, 1997; Mariucci et al., 1999; Boncio et al., 2000; Montone et al., 2004]. [21] In Figures 5a and 5b, we have also plotted all of the relocated seismicity together with the aftershocks (orange circles) of the 1984 Gubbio earthquake Figure 9. (a) Cumulative number of events that occurred along the Alto Tiberina fault (black line) and in the ATF hanging wall (grey line) versus time for the entire monitoring period (211 days). (b) and (c) Same data divided into two time intervals: the first 142 and the last 69 days, respectively. 10 of 22

11 Table 2. Location, Depth, Origin Date and Time, Magnitude, Identification Number, Magnitude, and Cluster Number of the 72 Peculiar Events Located Along the Alto Tiberina Fault Zone a Event Longitude Latitude Depth Year Month Day Hour Min M L ID Cluster of 22

12 Table 2. (continued) Event Longitude Latitude Depth Year Month Day Hour Min M L ID Cluster a See text for details. (M w 5.1), which is the closest and most recent main shock (see Figure 1). The 1984 aftershock sequence was also relocated using the double-difference algorithm [see Collettini et al., 2003] and has comparable final location errors. The Gubbio sequence is clearly located in the hanging wall of the ATF and does not crosscut the major fault. This observation is consistent with the interpretation of the ATF as a basal detachment that drives the deformation and probably constrains the size of the synthetic and antithetic seismogenic faults confining their related seismicity at depth. We will further discuss this issue in sections 4 and Seismicity Patterns [22] Having documented seismic activity associated with the ATF, we now investigate the seismicity pattern and the details of earthquake clustering. To isolate those earthquakes which are possibly located within the ATF detachment zone, we selected the seismic events that nucleated at a distance less than 500 m from the detachment as imaged by the seismic profiles. This selection provides us with 621 microearthquakes with M L 2.3 from the 1461 relocated ones that we show in map view in Figure 8, together with the isolines imaging the shape of the ATF at depth. In Figure 8 we plotted the earthquake slip patches inferred from the moment magnitude assuming a constant stress drop of 3 MPa. Given their very small source dimensions compared to the ATF, we multiplied the fault patch dimensions by a factor of five. The results show that the ATF-related seismicity is almost uniformly distributed over the entire fault plane. [23] We analyzed the rate of earthquake production for the events that nucleated in the hanging wall and those located along the ATF detachment zone for the entire 211 day experiment. We plotted in Figure 9a the cumulative number of earthquakes as a function of time for the two groups of events. For the entire duration of our experiment, we measured a seismicity rate of 2.9 events per day along the ATF and 3.8 events per day in the hanging wall of the ATF. However, for the seismicity located in the ATF hanging wall, there is a clear change in the rate of earthquake production in March 2001 (see black line in Figure 9a). This variation is due to a change of the network geometry. At that time of the experiment, three stations were moved from the peripheral area toward the central sector, where the two minor faults (F1, F2) are located. This caused a decrease in the detection threshold for the shallower seismicity that is mainly located in the hanging wall volume of the ATF, while not significantly affecting the seismicity rate along the ATF. For the ATF-related seismicity, the rate varies from 2.8 events per day in the first time interval (142 days before 7 March 2001, grey line in Figure 9b) to 3.2 events per day in the second time period (69 days after 7 March 2001, grey line in Figure 9c). On the contrary, in the hanging wall of the ATF, the seismicity rate changes from 2.7 events per day (black line in Figure 9b), to 6.0 events per day in the second time period (after 7 March 2001, black line in Figure 9c). [24] These results highlight two main findings. The first is that the rate of earthquake production in the hanging wall block is higher in comparison with that observed along the ATF, the latter being nearly constant. The second is that despite the possibility to constrain a precise value for the seismicity rate with the available data, background seismicity is characterized by a nearly constant rate of earthquake production that does not show an Omori-like temporal decay typical of aftershock sequences. This implies that the earthquakes investigated in this study are mainly due to regional tectonic stress and local mechanical properties of the fault zones. As expected, to investigate the level of background seismicity in areas characterized by relatively low strain rates, it is necessary to deploy dense seismic networks and to reduce the minimum detection magnitude. This is the case of most of Italian seismogenic areas. [25] The earthquakes located on the ATF detachment show several interesting features. We identified 28 groups of seismic events, comprising 72 separate earthquakes (red circles in Figure 8, listed in Table 2) out of 621 (12% of the ATF-related seismicity), that nucleated within 250 m of each other. Each is typically composed of two to three earthquakes, with the exception of one larger group composed of 12 events (cluster 17 in Table 2). The events in each group also nucleate within a very short time delay, and we therefore consider them to be clusters. As expected, the seismic waveforms of these earthquakes are almost identical (e.g., see Figure A2 in Appendix A, which refers to cluster 17 in Table 2). According to our location resolution, four pairs of these events (belonging to clusters 25, 14, and 3 in Table 2) have their slipping patches partially overlapping (Figures 10a, 10b, and 10c). Thus they behave as repeating earthquakes (repeaters): earthquakes of similar size located in the same area which repeatedly rupture the same fault patch [Schaff et al., 2002; Waldhauser et al., 2004]. [26] The time delays between earthquakes belonging to each cluster are on the order of minutes to hours: 60% of them occur within 1 h, 75% within 3 hours and 100% within 1 d. A few (15%) occur within 1 min. We have plotted in Figure 11 the time delays for each earthquake pair belonging to a cluster as a function of their subsequent occurrence (see Table 2). In our data set each cluster begins when the seismic activity in the previous one is terminated (in other words, there is no temporal overlap between two subsequent clusters). A short elapse time might suggest the existence of interaction between these subsequent earthquakes, although they do not show any clear relation with their magnitude: the magnitude of the first event appears randomly smaller or bigger compared to the subsequent one (Table 2). In order to investigate the interaction between individual clusters occurring at different times, we computed the time delay and distance between the occurrence of two subsequent clusters (from the origin time and the location of the first 12 of 22

13 Figure 11. Histogram of the delay times between the occurrence of subsequent earthquakes belonging to the same cluster. See text for explanation. shock of each cluster) and we estimate an equivalent migration velocity (Figure 12). The inferred equivalent migration velocity is quite high, in the order of 0.1 km/h (2.4 km/d). We also report the inferred distance between subsequent clusters as a function of azimuth (rose diagram in Figure 12). The results shown in Figure 12 reveal that there is no evidence for a preferred migration direction, thus suggesting that these clusters are most likely controlled by local rheological properties within the entire ATF fault zone. The absence of prevalent migration directions is also consistent with the lack of fault slip parallel alignments of hypocenters along the fault plane (e.g., streaks) and the Figure 10. Location on the detachment plane of three groups of repeating earthquakes. They can be identified by their ID numbers in Table 2. The circles represent the slipping patch obtained from moment magnitude and assuming a stress drop value of 3 MPa (see text for explanation), while the bars represent formal errors (see Appendix A). Figure 12. Histogram showing the number of multiple events versus the velocity migration (see text for explanation). The rose diagram shows the azimuth of the distance vector between the subsequent clusters. 13 of 22

14 Figure 13. Gutenberg-Richter distribution of the seismicity located (a) in the hanging wall and (b) on the ATF detachment plane. M c is the completeness magnitude calculated using the ZMAP code [Wiemer, 2001]. The straight lines represent the theoretical frequency-magnitude distribution resulting from the two b values estimated in this study: 1.06 ± 0.07 for the ATF-related seismicity and 0.85 ± 0.03 for the seismicity located in the hanging wall block. uniform distribution of seismicity on the ATF, although it should be realized that the monitoring time may be too short to preclude the existence of preferred directions for seismicity migration. [27] We then examined the cumulative earthquake size distribution (G-R relation [Gutenberg and Richter, 1944]) for the two groups of relocated earthquakes: 840 from the hanging wall of the ATF and 621 located on the ATF. The results are shown in Figure 13; the b value and the completeness magnitude (M c ) have been computed using the ZMAP code [Wiemer, 2001]. It is well known that the estimate of the b value depends critically on the completeness of the catalogue. As seen in Figure 13, the frequencymagnitude distribution deviates from a linear power law for smaller magnitudes; this is due to the lack of precisely located events with magnitudes smaller than the M c [see Wiemer and Wyss, 2002]. This magnitude is smaller for the shallower sampled volume: M c = 0.5 and M c = 1 for the hanging wall and the ATF-related seismicity, respectively (Figure 13). The b values estimated for the two groups of earthquakes are quite different: b is equal to 1.06 ± 0.07 for the ATF and 0.85 ± 0.03 for the hanging wall block. We performed a Z test [Habermann, 1987] to verify the statistical significance of the two values using the following Z parameter: Z ¼ q ð a 1 a 2 Þ ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi s 2 1 þ ; s2 2 where a 1 and a 2 are the two b values and s 1 and s 2 the related errors computed for each catalogue. The resulting value of the Z parameter is 2.7, which means that the computed b values have a significance close to 99%. [28] Despite the fact that a constant value of b equals 1 in large sampled volumes has been considered as a universal constant [Gerstenberger et al., 2001], recent papers [see Wiemer and Wyss, 2002, and references therein] have shown statistically significant variations of this value: these difference have been observed in laboratory experiments (i.e., acoustic emissions), from real seismicity in mines or distinct tectonic regimes (e.g., subducting slabs or crustal fault zones), in volcanic areas (magma chambers) as well as in aftershock sequences. Schorlemmer et al. [2005] found that the b value also varies systematically for different styles of faulting. Gerstenberger et al. [2001] observed in different tectonic areas that the b value decreases with depth and explained this observation by invoking different rheological properties for distinct seismogenic volumes. Our results show that the rate of earthquake production is higher and b value is smaller in the hanging wall of the ATF, suggesting a higher probability of moderate magnitude earthquakes (5.0 < M < 6.0) occurring in this region. By contrast, along the ATF detachment zone, the seismicity rate is smaller and the b value is larger which can indicate a tendency to have more microearthquakes than moderate magnitude events. The occurrence of the Gubbio 1984 seismic sequence in the ATF hanging wall and the lack of historical earthquakes clearly nucleating along the ATF may support this hypothesis. These interpretations have important implications for the assessment of seismic hazard in the study area. 5. Discussion [29] The integration of surface geology, borehole data, seismic reflection profiles and seismological data consistently demonstrate that crustal extension is accommodated by an east dipping LANF and associated more steeply dipping synthetic and antithetic structures in a tectonic regime characterized by a nearly vertical s 1. Microearthquakes define a fault zone 500 to 1000 m thick and the available fault plane solutions have a nodal plane consistent with the geometry inferred for the LANF. These findings have profound implications for fault rheology and seismic hazard in extensional tectonic settings and therefore deserve an in-depth discussion. 14 of 22

15 [30] The CROP03 seismic reflection profile (Figure 1b) [Barchi et al., 1998] suggests that a substantial amount of regional extension is accommodated by low-angle east dipping normal faults in the Apennines. Ancient and exhumed examples of these faults have been mapped at the surface and interpreted as being previously active low-angle detachments [e.g., Hayman et al., 2003; Axen 2004; Collettini and Holdsworth, 2004]. The ATF represents the most recent and currently active detachment that accommodates crustal extension. It extends 60 km along the Apennine belt and shows a minimum total displacement of 5 km, with 2 km accumulated in the last 2 Ma. Geological and seismological data show that the ATF is an active structure that separates a seismically active hanging wall block from an aseismic footwall. Microseismicity is distributed along the east dipping detachment at depths ranging between 3 4 and 16 km. The low-angle surface defined by the alignments of seismicity and seismic reflectors cannot be interpreted as the brittle-ductile transition (compare Hatzfeld et al. [2000] in the Gulf of Corinth [see also Rigo et al., 1996; Rietbrock et al., 1996]) because in an extending continental crust the transition cannot reach 3 4 km of depth and the modeling of heat flow data in the northern Apennines requires that this transition is situated well below the LANF [Pauselli and Federico, 2002] (see also Figure 1b). In the ATF hanging wall, we observe minor synthetic and antithetic high-angle normal faults that root down into the detachment suggesting the simultaneous activity of the whole normal fault system. The ATF acts as a detachment over which high-angle normal faults in the hanging wall detach and slip seismically, generating both microseismicity (e.g., F1 and F2 in Figure 5) and moderate magnitude earthquakes (5 < M < 6) such as the 1984 Gubbio main shock. The results emerging from these independent and multidisciplinary data sets agree with the interpretation that crustal extension in the study area is mainly accommodated by LANF. Moreover, in the present-day active stress field the ATF plays a major role in localizing tectonic strain into its hanging wall and in generating microearthquake within the fault zone. [31] The long-term regional displacement observed across the LANF from surface geology and seismic reflection profiles and the microearthquakes located along the ATF with compatible focal mechanisms, raise interesting questions concerning the fault zone rheology and the mechanical properties of the active structures. If we assume Anderson- Byerlee frictional fault mechanics (i.e., a constant friction coefficient along a fault plane governed by a brittle rheology), the high angle between the ATF and the maximum principal stress (75, Figure 7) implies that this active structure is severely misoriented with respect to the presentday regional stress field. The presence of earthquakes and coseismic slip on the ATF detachment means that the fault has to be weak in an absolute sense in order to generate earthquakes during slip. In this case we have implicitly assumed that all the seismicity located on the ATF ruptures the nodal plane with low dip angle consistent with the detachment geometry. This interpretation leaves open the possibility of having moderate-to-large earthquakes occurring on the ATF. Different physical processes have been proposed to explain fault weakness such fluid overpressures [Axen, 1992] or low friction coefficient [Townend and Zoback, 2000]. In this framework, our results could represent striking evidence for seismic activity on LANF. [32] An alternative scenario assumes that the ATF accommodates the crustal extension solely by ductile deformation (aseismic slip and creeping) and that the related seismicity occurs close to the ATF, but above the actual low-angle fault plane, as a brittle response to strain localization. In other words, earthquake slip events do not rupture the ATF detachment and correspond to the high dip angle nodal plane of the computed focal solutions (i.e., the conjugate plane to that considered in the previous scenario). In this case, the ATF accommodates deformation by a ductile process not following frictional fault mechanics and therefore the LANF enigma would not exist. However, there is no independent geological evidence to suggest that such ductile deformation occurs along the LANF in the Apennines. [33] A third possibility is related to the complex architecture of some faults where discrete planes of slip are distributed within a wide, m thick fault zone possessing different rheological properties [Faulkner et al., 2003; Wibberley and Shimamoto, 2003; Sibson, 2003] and following frictional behavior described by rate- and statedependent constitutive laws [Dieterich, 1979; Ruina, 1983; Marone, 1998]. Our results suggest that the ATF has a finite thickness ranging between 500 and 1000m. In this context, we propose that microseismicity occurs within the fault zone in velocity weakening patches which are loaded by a combination of (1) tectonic stress, (2) stress redistribution caused by aseismic slip or creep in adjacent volumes, (3) coseismic slip of nearby earthquakes, (4) fluid pressure fluctuations, and (5) pore pressure relaxation. Most of the ATF fault zone should slip aseismically or creep, because it is assumed to be velocity strengthening (see Boatwright and Cocco [1996] for a discussion on the different frictional regimes characterizing faulting and earthquake mechanics). Because microearthquakes occur in a volume, they can rupture fault planes with different geometry depending on the orientation of the local stress field [Rice, 1992; Faulkner et al., 2006]. In cases of this kind, we do not have to pay attention to which nodal plane ruptures during single events. This interpretation of the rheology and mechanical properties of the low-angle normal fault, is consistent with a fault that accommodates extension and it is also consistent with the presence of microseismicity within the fault zone as well as with the existence of clusters and repeaters [Boatwright and Cocco, 1996; Gomberg, 2001]. In this framework, the friction coefficient is variable and depends on the local rheological properties. The model is also consistent with the observed spatial variation of b values. [34] We emphasize that all the three scenarios described above rely on the assumption of fault zone weakness. In the short term, the origin of fault weakening is most likely to be due to fluid overpressure. Since the onset of extension, the northern Apennines have been affected by deep-seated CO 2 degassing [Chiodini et al., 2004]. The boreholes located in the footwall of the ATF (Figure 2), reveal the presence of CO 2 at 85% of the lithostatic pressure at depth around km. It is important to note that according to the first scenario proposed above, the fluid pressure condition P f > s 3, required to reactivate a severely misoriented fault such 15 of 22

16 as the ATF, cannot be maintained for a long time, due to the increase of permeability under low effective stress [Collettini and Barchi, 2002]. Therefore it is more likely that fluid overpressures only create local, short-lived frictional instabilities, lasting from minutes to hours, responsible for the nucleation of microseismicity. This is consistent with our third scenario and with the observed short delay and continuous occurrence of the multiple events within the whole fault zone (Figure 10). [35] In the long term, the fluid influx into the fault zone might alter the load bearing mineral phases resulting in finegrained aggregates of interconnected weak, platy minerals (micas, clay) that reduce the friction coefficient to sub- Byerlee values, m s < 0.3 [Bos and Spiers, 2002; Holdsworth, 2004]. The occurrence of these weakening mechanisms has already been inferred from textural studies made in fault rocks from an exhumed analogue of the ATF located within the older, extended portion of the Apennines (the Zuccale fault [see Collettini and Holdsworth, 2004]) (Figure 1b). This implies that chemical and mineralogical transitions during the interseismic period are likely to contribute to changing the rheological properties of the fault zone. [36] The last issue that deserves to be discussed concerns the seismogenic potential of the ATF and the likelihood that it will generate moderate-to-large magnitude earthquakes. The differences in b values documented for the ATF (1.06 ± 0.07) and for hanging wall faults (0.85 ± 0.03) might suggest different fault zone properties for the two seismogenic volumes. The ATF shows a tendency to generate more microearthquakes than moderate-magnitude seismic events, contrary to what is thought for the hanging wall faults. The difference in the b values proposed is consistent with what is observed in the San Andreas fault: the creeping portion has a higher b value (1.2) than that observed for the locked portion (b = ) [Amelung and King, 1997]. Wyss and Wiemer [2000] correlate high b value with asperity locations. Their observations are consistent with the frictional behavior of crustal faulting proposed by Boatwright and Cocco [1996]. However, the different b values do not prove that the ATF is creeping while the hanging wall faults are locked. Thus this evidence cannot be used unambiguously to assess the seismic hazard for damaging earthquakes on the ATF. [37] More convincing evidence for the low probability of a large earthquake on the ATF arises from the evaluation of the slip rate value. The ATF has a long-term slip rate of 1 mm/a in the last 2 Ma, and it is located in an area characterized by extensional geodetic strain rate of roughly 2.5 mm/a. By considering the dimension of the fault and its slip rate, the ATF should be capable of generating a M 7.0 event with a recurrence interval of roughly 400 years. However, in the area where the ATF is located (Figure 1b) there have been no historical earthquakes and, for this sector of the Apennines, the Italian catalogue of large historical earthquakes [Boschi et al., 1998] is complete for the past 2000 years. Importantly, the observed microseismicity is not enough to achieve the observed slip rate. These two observations suggest that the fault is most likely to accommodate most of the observed displacement aseismically with fluid-assisted hydrofracturing processes responsible for the microseismicity. Although we are aware that only a dense GPS local network can give information to constrain the ATF style of faulting, several lines of evidence point to significant amounts of aseismic behavior. First, for earthquakes occurring less than 500 m away from the ATF, as imaged on the seismic reflection profiles, the microseismicity is homogeneously distributed over the entire fault area with an absence of obvious areas devoid of microseismicity. The ATF seismicity rate is high (3 events per day) and mainly constant with time. These observations suggest that the fault is incapable of accumulating significant stress to generate a large earthquake but that it instead releases stress continuously with time. Secondly, the higher b value for the ATF-related microseismicity suggests that the fault plane has a different rheology compared to the contemporaneous faults in the hanging wall; this would also be consistent with the creeping behavior of the detachment. Finally, field studies on the Zuccale fault, the exhumed (3 6 km) analogue of the ATF, have documented the presence of a phyllosilicate-rich fault core accommodating deformation by stress-induced pressure solution [Collettini and Holdsworth, 2004]. Applying the frictional viscous creep model of Bos and Spiers [2002] and Niemeijer and Spiers [2005], this would require the pressure solution accommodating deformation along the phyllosilicate foliae at low slip rates to be a velocity strengthening process. Since most of the ATF fault trace lies within the Phyllitic basement (Figure 3a) the same deformation process may operate along the fault being a possible indication of aseismic slip. This implies that microseismicity occurs in velocity weakening patches, which might be caused by local short-lived attainment of CO 2 overpressure as testified by the two boreholes located in the footwall of the ATF. Therefore we believe that our third case scenario represents the most general and reliable explanations for our observations. We are aware that low-angle normal faults do not occur only in phyllitic rocks, since other LANF have been documented in quartzo-feldspathic rocks [e.g., Axen, 2004]. The latter may behave differently in terms of frictional regime from phyllosillicates-rich fault zones. 6. Conclusions [38] In this study we have presented evidence for seismicity associated with an active low-angle normal fault (dip 15 ) situated in the northern Apennines, Italy. The 2000 earthquakes investigated in this study were recorded by a local network deployed in the study area for 8 months. We relocated the events using the DD algorithm employing cross correlations of recorded waveforms. Our integrated geological observations, interpretation of seismic reflection profiles and active seismicity data sets, independently define an east dipping low-angle normal fault 60 km long and 40 km wide, the ATF. In the last 2 Ma, this structure has accumulated 2 km of displacement and it is currently microseismically active. The fault is the major structure of the area that separates an active hanging wall block from an aseismic footwall. In the hanging wall block, seismic reflection profiles and seismological data reveal the presence of moderately to steeply inclined minor faults soling into the detachment. The computed focal mechanisms are in agreement with the geometry of the faults highlighted by aftershock distributions and seismic reflection profiles and are consistent with a stress field characterized by a nearly 16 of 22

17 Table A1. Picking, Reading, and Weighting Scheme Pick Quality Reading Errors, s Weights (Double Difference) 0 < vertical s 1 and a NE trending s 3, perpendicular to the strike of the detachment, which has also been inferred from independent data. The focal mechanisms calculated for the ATF-related seismicity show one nodal plane that is consistent with a gently east dipping geometry for the fault and its normal faulting kinematics. [39] During 8 months of temporary monitoring, the ATF fault zone showed a high and nearly constant rate of earthquake production of 3 events per day. Our data set comprising 621 microearthquakes with M L < 2.3 located within the Alto Tiberina fault zone and 840 events with M L < 3.1 located on normal faults in its hanging wall. This microseismicity is uniformly distributed over the ATF plane and the earthquake distribution in the downdip direction reveals a fault zone thickness of between 500 and 1000 m. We observe the occurrence of repeating earthquakes in very small slip patches with dimensions of the order of m. These multiple events occur within a very short time delay from each other (i.e., minutes/hours) and without any evident correlation to their size. Moreover, we observe that the ATF-related seismicity is characterized by higher b values (1.06 ± 0.07) than that inferred for the seismicity located in the contemporaneously deforming hanging wall block (b = 0.85 ± 0.03). [40] We interpret these results, together with the absence of historical earthquakes associated with the ATF, the presence of a source of overpressurized fluids located in the fault hanging wall, and the likely development of a phyllosilicaterich fault core at depth (compare the Zuccale fault) as being consistent with the observed aseismic behavior for this misoriented fault. In this case, microseismicity occurs in distinct velocity weakening patches and could be generated by local, short-lived build ups in fluid pressure during regional-scale degassing of the mantle during regional tectonic extension. Appendix A A1. Relocation Method [41] The double-difference (hereinafter DD) algorithm minimizes the residuals between observed and calculated traveltime differences for pairs of earthquakes at common stations by iteratively adjusting the vector difference between the hypocenters. The DD technique can be applied and yields very accurate locations if the hypocentral separation between two earthquakes is small compared to the source receiver separation. In this case, seismic rays from two events to a common station can be considered similar along almost the entire path [Fréchet, 1985; Got et al., 1994] and relative traveltimes difference only depends on the spatial offset between the events [Poupinet et al., 1984] that can be calculated by differencing Geiger s equation for earthquake location (see Waldhauser and Ellsworth [2000] for a comprehensive review). Thus the DD method minimizes the effects of unknown Earth structure, without the need for station corrections. [42] The algorithm allows us to use ordinary phase picks from earthquake catalogues and high-precision differential traveltimes from waveform correlation. These data are combined together into a system of linear equations with each event pair (k, l) at each station (i), forming the ik rk r l ¼ dtik obs dt cal ik dt obs il dt cal il ða1þ where r is the hypocentral adjustment vector and n is the four-component vector of Cartesian coordinates and origin time. The hypodd code [Waldhauser, 2001] computes the solution to equation (A1) by weighted least squares (LSQR by Paige and Saunders [1982] using the conjugate gradients method. Higher-resolution hypocenters are found by iteratively adjusting the vector difference between hypocentral pairs. The hypocentral parameters and their related new partial derivatives are updated after each iteration. [43] P and S wave differential traveltimes derived from waveform cross correlation and P and S wave catalog traveltime differences can be combined and the same equation (A1) is used for all the data. The reliable recognition of small interevent distances within clusters of correlated events (multiplets) is ensured by the high accuracy of the cross-correlation data, while relative locations between multiplets and uncorrelated events are determined to the accuracy of the arrival time data. A2. Application to the Data Set [44] In order to constrain the geometry of the ATF using seismicity distributions, we have relocated the earthquakes collected during a seismic survey performed around the Tiber basin in the Città di Castello area in [Piccinini et al., 2003]. A dense seismic network composed of 30 three component seismic stations was installed across the ATF system and maintained for 8 months. The network (geometry shown in Figure 4) recorded about 2000 earthquakes with M L 3.1, which represent the background seismicity of the area. To obtain high-resolution hypocenter locations, we have used the code hypodd [Waldhauser, 2001] based on the DD algorithm (see the details in the work by Waldhauser and Ellsworth [2000]), and here we describe the details of the performed relocation. [45] We have used phase picks from the three components of the recorded seismograms and high-precision differential traveltimes from waveform correlation. Phase pick data from five stations of the Italian National Network (RSNC) located within 60 km from the area were added to the data derived from the 30 local stations. The accuracy of the earthquake catalogue phase readings span from 0.03 s for the best hand-picked data discarding those events with reading error 0.1 s as shown in Table A1. [46] To compute partial derivatives of theoretical traveltimes, we used a 1-D P wave velocity model (see Table 1) obtained from seismic data [Moretti, 2005] and the velocities of the rocks inferred from borehole data [Bally et al., 1986]. The V p /V s ratio adopted is equal to [47] The starting hypocentral locations are those computed by Piccinini et al. [2003]. From this initial data set of 17 of 22

18 Figure A1. Catalog of traveltime differences for event pairs as a function of the distance between the events [after Waldhauser and Ellsworth, 2002], computed for 1679 events recorded in [Piccinini et al., 2003]. The straight line represents traveltime differences between two events computed with a constant source region velocity of 5 km/s. The line is shifted by 0.5 s along the time axis to account for possible uncertainties in the catalog locations. Traveltime differences above the straight line are considered outliers events, we selected 1679 with a minimum number of 10 P and S phase picks. Traveltimes routinely computed from P and S wave arrival times of the selected events recorded by the local network have been differenced for pairs of earthquakes at each station that observed the couple of events. Double-difference equations are built to link each event to several neighbors, so that all the events are connected and the solution for the adjustment to each hypocenter can simultaneously be determined. To keep the effect of raypath differences outside the source region [Waldhauser and Ellsworth, 2000], only event pairs with hypocentral separation <10 km have been considered. [48] Figure A1 shows 56,000 selected catalogue P and S wave traveltime differences as a function of separation distance obtained from the 1679 events; 84% and 86% of the P and S wave catalogue traveltime differences, respectively, are smaller than the maximum traveltime difference expected for a given event pair offset when the station lies on the line connecting both events. Traveltime differences are close to 0 s at zero event pair offset and generally smaller than 2 s at offsets of 10 km, indicating an average source region velocity of 5 km/s. Data above the indicated straight line in Figure A1 are considered to be outliers and are removed before relocation. The concentration of outliers near the zero offset is due to the increased number of observations for nearby events. [49] In addition to the catalogue data, we measured traveltime differentials for each event pair with waveforms that correlated at a common station by using the time domain cross correlation method described by Schaff [2001] and Schaff et al. [2004]. We used the digital waveforms recorded on vertical and horizontal component seismometers filtered using a 4 pole-zero phase band-pass Butterworth filter (1 20 Hz). Because of the low magnitude of the recorded events and to the high-frequency content, we evaluated two waveforms recorded at a specific station to be similar when half of the squared coherency values, in the frequency range 1 20 Hz, of a tapered s (128 samples) window containing the P wave (or S wave) train exceed 0.8. Figure A2 (top) shows an example of similar waveforms, with similarity values exceeding 0.82, which are related to seven earthquakes closely located (see Figure A2, bottom). For sufficiently similar signals, a precision of 1 ms can be achieved for the measurement of time differences from data digitized at 10 ms intervals [Poupinet et al., 1984; Fréchet, 1985]. The distribution of coherency values for and 8678 measured P and S waves, respectively, are shown in Figure A3. Both distributions feature a peak at coherency while we consider waveforms to be similar showing values greater than 0.8. Outliers are difficult to detect prior to relocation, but we can assess the consistency between the P and S wave data. However, coherency values of these data are typically low, and the measurements will therefore be downweighted during relocation. [50] We performed 25 iterations in the relocation procedure optimizing the damping parameter to obtain good conditioning and convergence of the solution. The iterations were grouped into five sets for a better control of the inversion procedure. During the first set of iterations, the algorithm runs only using the full weight on catalogue phase picks in order to remove the influence of velocity 18 of 22

19 Figure A2. (top) Example of similar waveforms recorded on vertical components at the station D002, containing P and S wave trains. (bottom) Map with the location of the earthquakes. model error over long distance ranges. As the locations improve and interevent distances become more accurate, the data derived from waveform correlations have been increasingly weighted because of their higher measurement precision at shorter distance ranges. Finally, the catalogue data are downweighted by a factor of 100 relative to the crosscorrelation data while equal weights are used for P and S wave cross-correlation data. During the iterations, the residuals are reweighted according to the misfit and the distance between the events following the functions proposed by Waldhauser and Ellsworth, 2000, Figure 4]. Catalogue and cross-correlation data are removed/ reweighted for event pairs with separation distances larger than/smaller than 10 and 2 km, respectively. [51] Hypocentral errors have been estimated by using the singular valued decomposition (SVD) to solve equation (A1) (see Waldhauser and Ellsworth [2000] for further details) for earthquake subsets, because error estimates obtained with LSQR (formal errors) are too optimistic. This analysis indicates relative location errors of a few meters between nearby events, which are constrained mainly by crosscorrelation data. Conversely, horizontal and vertical formal errors in hypocentral locations, which have been estimated between events predominantly constrained by phase picks, are of the order of 50 m and 70 m, respectively. These errors reflect the relative location uncertainty, while the absolute location uncertainty may be larger [see Waldhauser and Ellsworth, 2000]. The final catalogue consists of 1416 events having a substantial reduction of final RMS and 19 of 22

20 Figure A3. Quality of cross-correlation data. Histogram of coherency values of (top) P (black) and (bottom) S (grey) phases used to determine traveltime differences between events at common stations. formal location errors. The initial RMS residuals (114 ms) and the related location errors after the relocation procedure are decreased by more than 60% for the phase pick data (34 ms) and by 93% from for cross correlation data (8 ms). [52] It is possible to appreciate in Figure 4 both in map view and in vertical cross section the difference in earthquake distributions due to the relocation process. The doubledifference procedure allows us to substantially enhance the hypocentral locations leading us to infer the geometry of the ATF fault and associated hanging wall structures. [53] Acknowledgments. We wish to thank Bill Ellsworth, Nicola D Agostino, Darrel Cowen, David Schaff, Eiichi Fukuyama, and Massimiliano Barchi for helpful discussions and for reviewing a preliminary version of the manuscript. We are indebted to Sandy Steacy, Robert Holdsworth, and Gary Axen for their helpful reviews and detailed comments which allowed us to substantially improve the paper. We thank Enzo Boschi for his continuous encouragements in completing this study. Lauro Chiaraluce has been supported by the PROSIS project funded by the Italian Ministry of University and Research. References Abers, G. A. (1991), Possible seismogenic shallow-dipping normal faults in the Woodlark-D Entrecasteaux extensional province, Papua New Guinea, Geology, 19, Alvarez, W. (1972), Rotation of the Corsica-Sardinia microplate, Nature, 248, Amato, A., et al. (2000), Terremoti probabili in Italia tra l anno 2000 e il 2030: Elementi per la definizione di priorità e gli interventi di riduzione del rischio sismico. Attività Scientifica GNDT, report, Dep. of the Civ. Protect., Rome. Ambrosetti, P., et al. (1978), Evoluzione paleogeografica e tettonica dei bacini tosco-umbro-laziali nel Pliocene e nel Pleistocene inferiore, Mem. Soc. Geol Ital., 19, Amelung, F., and G. King (1997), Earthquake scaling laws for creeping and non-creeping faults, Geophys. Res. Lett., 24, 5, Antonioli, A., D. Piccinini, L. Chiaraluce, and M. Cocco (2005), Fluid flow and seismicity pattern: Evidence from the 1997 Umbria- Marche (central Italy) seismic sequence, Geophys. Res. Lett., 32, L10311, doi: / 2004GL Axen, G. J. (1992), Pore pressure, stress increase, and fault weakening in low-angle normal faulting, J. Geophys. Res., 97, Axen, G. J. (1999), Low-angle normal fault earthquakes and triggering, Geophys. Res. Lett., 26, Axen, G. J. (2004), Mechanics of low-angle normal faults, in Rheology and Deformation of the Lithosphere at Continental Margins, edited by G. D. Karner et al., pp 46 91, Columbia Univ. Press, New York. Bally, A. W., L. Burbi, C. Cooper, and L. Ghelardoni (1986), Balanced sections and seismic reflection profiles across the central Apennines, Mem. Soc. Geol. Ital., 35, Barchi, M., R. Minelli, and G. Pialli (1998), The crop 03 profile: A synthesis of results on deep structures of the northern Apennines, Mem. Soc. Geol. Ital., 52, Boatwright, J., and M. Cocco (1996), Frictional constraints on crustal faulting, J. Geophys. Res., 101, 13,895 13,909. Boncio, P., F. Brozzetti, and G. Lavecchia (2000), Architecture and seismotectonics of a regional low-angle normal fault zone in central Italy, Tectonics, 19, Bos, B., and C. J. Spiers (2002), Frictional-viscous flow of phyllosilicatebearing fault rock: Microphysical model and implications for crustal strength profiles, J. Geophys. Res., 107(B2), 2028, doi: / 2001JB Boschi, E., E. Guidoboni, G. Ferrari, and G. Valensise (1998), I terremoti dell Appennino Umbro-Marchigiano (area sud orientale dal 99 a.c. al 1984)-ING-SGA, 267 pp., Compositori, Bologna, Italy. Byerlee, J. D. (1978), Friction of rocks, Pure Appl. Geophys., 116, Carmignani, L., and R. Kligfield (1990), Crustal extension in the northern Apennines: The transition from compression to extension in the Alpi Apuane core complex, Tectonics, 9, Chiarabba, C., and A. Amato (2003), Vp and Vp/Vs images in the Mw 6.0 Colfiorito fault region (central Italy): A contribution to the understanding of seismotectonic and seismogenic processes, J. Geophys. Res., 108(B5), 2248, doi: /2001jb Chiarabba, C., L. Jovane, and R. Di Stefano (2005), A new view of Italian seismicity using 20 years of instrumental recordings, Tectonophysics, 395, Chiaraluce,L.,W.L.Ellsworth,C.Chiarabba,andM.Cocco(2003), Imaging the complexity of an active normal fault system: The 1997 Colfiorito (central Italy) case study, J. Geophys. Res., 108(B6), 2294, doi: /2002jb Chiaraluce, L., et al. (2004), Complex normal faulting in the Apennines Thrust and fold belt: The seismic sequence in central Italy, Bull. Seismol. Soc. Am., 94(1), Chiodini, G., C. Cardellini, A. Amato, E. Boschi, S. Caliro, F. Frondini, and G. Ventura (2004), Carbon dioxide Earth degassing and seismogenesis in central and southern Italy, Geophys. Res. Lett., 31, L07615, doi: / 2004GL Cocco, M., et al. (2000), Sviluppo e confronto di metodologie per la valutazione della pericolosità sismica in aree sismogenetiche: Applicazione all Appennino Centrale e Meridionale. Attività Scientifica GNDT, report, Dep. of the Civ. Protect., Rome. Collettini, C., and M. R. Barchi (2002), A low angle normal fault in the Umbria region (central Italy): A mechanical model for the related microseismicity, Tectonophysics, 359, Collettini, C. and M. R. Barchi (2004), A comparison of structural data and seismic images for low-angle normal faults in the northern Apennines (central Italy): Constraints on Activity, in Flow Processes in Faults and Shear Zones, edited by G. I. Alsop and R. E. Holdsworth, Geol. Soc. Spec. Publ., 224, Collettini, C., and R. E. Holdsworth (2004), Fault zone weakening processes along low-angle normal faults: insights from the Zuccale Fault, Isle of Elba, Italy, J. Geol. Soc., 161, Collettini, C., and R. H. Sibson (2001), Normal faults normal friction?, Geology, 29, of 22

21 Collettini, C., M. R. Barchi, L. Chiaraluce, F. Mirabella, and S. Pucci (2003), The Gubbio fault: Can different methods give pictures of the same object, J. Geodyn., 36, Dieterich, J. H. (1979), Modeling of rock friction: 1. Experimental results and constitutive equations, J. Geophys. Res., 84, Elter, P., G. Giglia, M. Tongiorgi, and L. Trevisan (1975), Tensional and contractional areas in recent Tortonian to Present evolution of the northern Apennines, Boll. Geofis. Teorica Appl., 17, Faulkner, D. R., A. C. Lewis, and E. H. Rutter (2003), On the internal structure and mechanics of large strike-slip faults: Field observations from the Carboneras fault, southeastern Spain, Tectonophysics, 367, Faulkner, D. R., T. M. Mitchell, D. Healy, and M. J. Heap (2006), Slip on weak faults by the rotation of regional stress in the fracture damage zone, Nature, 444, Floyd, J. S., J. C. Mutter, A. M. Goodliffe, and B. Taylor (2001), Evidence for fault weakness and fluid flow within active low-angle normal fault, Nature, 411, Fréchet, J. (1985), Sismogénèse et doublets sismiques, thése d Etat, 206 pp., Univ. Sci. et Méd. de Grenoble, Grenoble, France. Frepoli, A., and A. Amato (1997), Contemporaneous extension and compression in the northern Apennines from earthquake fault-plane solutions, Geophys. J. Int., 129, Gerstenberger, M., S. Wiemer, and D. Giardini (2001), A systematic test of the hypothesis that the b value varies with depth in California, Geophys. Res. Lett., 28(1), Gomberg, J. (2001), The failure of earthquake failure models, J. Geophys. Res., 106, 16,253 16,264. Got, J. L., J. Fréchet, and F. W. Klein (1994), Deep fault plane geometry inferred from multiplet relative relocation beneath the south flank of Kilauea, J. Geophys. Res., 99, 15,375 15,386. Gutenberg, R., and C. F. Richter (1944), Frequency of earthquakes in California, Bull. Seismol. Soc. Am., 34, Habermann, R. E. (1987), Man-made changes of seismicity rates, Bull. Seismol. Soc. Am., 77, Hatzfeld, D., V. Karakostas, M. Ziazia, I. Kassaras, E. Papadimitriou, K. Makropoulos, N. Voulgaris, and C. Papaioannou (2000), Microseismicity and faulting geometry in the Gulf of Corinth (Greece), Geophys. J. Int., 141, Hayman, N. W., J. R. Knott, D. S. Cowan, E. Nemser, and A. M. Sarna-Wojcicki (2003), Quaternary low-angle slip on detachment faults in Death Valley, California, Geology, 31, Holdsworth, R. E. (2004), Weak faults-rotten cores, Science, 303, Hunstad, I., G. Selvaggi, N. D Agostino, P. England, P. Clarke, and M. Pierozzi (2003), Geodetic strain in peninsular Italy between 1875 and 2001, Geophys. Res. Lett., 30(4), 1181, doi: /2002gl Jackson, J. A., and N. J. White (1989), Normal faulting in the upper continental crust: Observations from regions of active extension, J. Struct. Geol., 11, Jolivet, L., et al. (1998), Midcrustal shear zones in postorogenic extension: Example from the northern Tyrrhenian Sea, J. Geophys. Res., 103, 12,123 12,160. Keller, J. V. A., G. Minelli, and G. Pialli (1994), Anatomy of late orogenic extension: The northern Apennines case, Tectonophysics, 238, King, G. C. P., and M. Cocco (2001), Fault interaction by elastic stress changes: New clues from earthquake sequences, Adv. Geophys., 44, Laigle, M., A. Hirn, M. Sachpazi, and N. Roussos (2000), North Aegean crustal deformation: An active fault imaged to 10 km depth by reflection seismic data, Geology, 28, Lavecchia, G., F. Brozzetti, M. R. Barchi, M. Menichetti, and J. V. Keller (1994), Seismotectonic zoning in east-central Italy deduced from analysis of the Neogene to present deformations and related stress fields, Geol. Soc. Am. Bull., 106, Lister, G. S., and G. A. Davis (1989), The origin of metamorphic core complexes and detachment faults formed during Tertiary continental extension in the northern Colorado River region, USA, J. Struct. Geol., 11, Mariucci, M. T., A. Amato, and P. Montone (1999), Recent tectonic evolution and present day stress in the northern Apennines (Italy), Tectonics, 18, Marone, C. J. (1998), Laboratory-derived friction laws and their application to seismic faulting, Annu. Rev. Earth Planet. Sci., 26, Michael, A. (1984), Determination of stress from slip data: Faults and folds, J. Geophys. Res., 89, 11,517 11,526. Miller, S. A., C. Collettini, L. Chiaraluce, M. Cocco, M. R. Barchi, and J. B. Kraus (2004), Aftershocks driven by a high-pressure CO 2 source at depth, Nature, 427, Mirabella, F., M. G. Ciaccio, M. R. Barchi, and S. Merlini (2004), The Gubbio normal fault (central Italy): Geometry, displacement distribution and tectonic evolution, J. Struct. Geol., 26, Molnar, P. and W.-P. Chen (1982), Seismicity and mountain building, in Mountain Building Processes, edited by K. J. Hsu, pp , Academic, London. Mongelli, F., and G. Zito (1991), Flusso di calore nella regione Toscana, Studi Geol. Camerti, 1, Montone, P., M. T. Mariucci, S. Pondrelli, and A. Amato (2004), An improved stress map for Italy and surrounding regions (central Mediterranean), J. Geophys. Res., 109, B10410, doi: /2003jb Moretti, M. (2005), Analisi della struttura crostale in zone di lacuna sismica tramite tomografia sismica: Applicazione all Appennino Umbro- Marchigiano e Abruzzese, Ph.D. thesis, Univ. degli Studi di Bologna, Bologna, Italy. Niemeijer, A. R. and C. J. Spiers (2005), Influence of phyllosilicates on fault strength in the brittle-ductile transition: Insights from rock analogue experiments, in Microstructural Evolution and Physical Properties in High Strain Zones, edited by D. Bruhn and L. Burlini, Geol. Soc. Spec. Publ., 245, Nostro, C., M. Cocco, and M. E. Belardinelli (1997), Static stress changes in extensional regimes: An application to southern Apennines (Italy), Bull. Seismol. Soc. Am., 87, Paige, C. C., and M. A. Saunders (1982), LSQR: An algorithm for sparse linear equations and sparse least squares, Trans Math. Software, 8, Pauselli, C., and C. Federico (2002), The brittle/ductile transition along the Crop03 seismic profile: Relationship with the geological features, Boll. Soc. Geol. Ital., 121(1), Pauselli, C., M. R. Barchi, C. Federico, B. M. Magnani, and G. Minelli (2006), The crustal structure of the northern Apennines (central Italy): An insight by the CROP03 seismic line, Am. J. Sci., 306, Pialli, G., M. Barchi, and G. Minelli (1998), Results of the CROP03 deep seismic reflection profile, Mem. 52., 657 pp., Soc. Geol. Ital., Rome. Piccinini, D., et al. (2003), Microseismic study in a low seismicity area of Italy: The Città di Castello experiment, Ann. Geophys., 46(6), Pingue, F., and G. De Natale (1993), Fault mechanism of the 49 seconds subevent of the Irpinia (southern Italy) earthquake from levelling data, Geophys. Res. Lett., 20, Ponziani, F., R. De Franco, G. Minelli, G. Biella, C. Federico, and G. Pialli (1995), Crustal shortening and duplication of the Moho in the northern Apennines: A view from seismic refraction data, Tectonophysics, 252, Poupinet, G., W. L. Ellsworth, and J. Fréchet (1984), Monitoring velocity variations in the crust using earthquake doublets: An application to the Calaveras Fault, California, J. Geophys. Res., 89, Proffett, J. M. (1977), Cenozoic geology of the Yerington district, Nevada, and implications for the nature and origin of basin and range faulting, Geol. Soc. Am. Bull., 88, Reutter, K. J., P. Giese, and H. Closs (1980), Lithospheric split in the descending plate: Observations from the northern Apennines, Tectonophysics, 64, T1 T9. Rice, J. R. (1992), Fault stress state, pore pressure distribution, and the weakness of the San Andreas Fault, in Fault Mechanics and Transport Properties of the Rocks, A Festschrift in Honour of W. F. Brace, edited by B. Evans and T. F. Wong, pp , Academic, San Diego, Calif. Rietbrock, A., C. Tiberi, F. Scherbaum, and H. Lyon-Caen (1996), Seismic slip on a low-angle normal fault in the Gulf of Corinth: Evidence from high-resolution cluster analysis of microearthquakes, Geophys. Res. Lett., 23, Rigo, A., H. Lyon-Caen, R. Armijo, A. Deschamps, D. Hatzfeld, K. Makropoulos, P. Papadimitriou, and I. Kassaras (1996), A microseismic study in the western part of the Gulf of Corinth (Greece): Implications for large-scale normal faulting mechanisms, Geophys. J. Int., 126, Ripepe, M., D. Piccinini, and L. Chiaraluce (2000), Foreshock sequence of the September 26th, 1997 Umbria-Marche earthquakes, J. Seismol., 4, Roy, A. J., and L. L. Kenneth (1992), Seismic reflection evidence for seismogenic low-angle faulting in south-eastern Arizona, Geology, 20, Ruina, A. L. (1983), Slip instability and state variable friction laws, J. Geophys. Res., 88, 10,359 10,370. Schaff, D. P. (2001), 4D high resolution seismology: Repeating events and large scale relocation, Ph.D. thesis, 115 pp., Stanford Univ., Stanford, Calif. Schaff, D. P., H. R. Bokelmann, G. C. Beroza, F. Waldhauser, and W. L. Ellsworth (2002), High-resolution image of Calaveras Fault seismicity, J. Geophys. Res., 107(B9), 2186, doi: /2001jb Schaff, D. P., G. H. R. Bokelmann, W. L. Ellsworth, E. Zanzerkia, F. Waldhauser, and G. C. Beroza (2004), Optimizing correlation techniques for improved earthquake location, Bull. Seismol. Soc. Am., 94, of 22

22 Schorlemmer, D., S. Wiemer, and M. Wyss (2005), Variations in earthquake-size distribution across different stress regimes, Nature, 437, Sibson, R. H. (1985), A note on fault reactivation, J. Struct. Geol., 7, Sibson, R. H. (2003), Thickness of the seismic slip zone, Bull. Seismol. Soc. Am., 93, Sorel, D. (2000), A Pleistocene and still-active detachment fault and the origin of the Corinth-Patras rift, Greece, Geology, 28, Townend, J., and M. D. Zoback (2000), How faulting keeps the crust strong, Geology, 28, Waldhauser, F. (2001), HypoDD: A computer program to compute double-difference earthquake locations, U.S. Geol. Surv. Open File Rep., Waldhauser, F., and W. L. Ellsworth (2000), A double-difference earthquake location algorithm: Method and application to the northern Hayward Fault, California, Bull. Seismol. Soc. Am., 90, Waldhauser, F., and W. L. Ellsworth (2002), Fault structure and mechanics of the Hayward Fault, California, from double-difference earthquake locations, J. Geophys. Res., 107(B3), 2054, doi: /2000jb Waldhauser, F., W. L. Ellsworth, D. P. Schaff, and A. Cole (2004), Streaks, multiplets, and holes: High-resolution spatiotemporal behavior of Parkfield seismicity, Geophys. Res. Lett., 31, L18608, doi: / 2004GL Wernicke, B. (1995), Low-angle normal faults and seismicity: A review, J. Geophys. Res., 100, 20,159 20,174. Westaway, R., and J. Jackson (1987), The earthquake of the 1980 November 23 in Campania-Basilicata (southern Italy), Geophys. J. R. Astron. Soc., 90, Wibberley, C. A. J., and T. Shimamoto (2003), Internal structure and permeability of major strike-slip zones: the Median Tectonic Line in Mie Prefecture, southwest Japan, J. Struct. Geol., 25, Wiemer, S. (2001), A software package to analyze seismicity: ZMAP, Seismol. Res. Lett., 72, 2. Wiemer, S., and M. Wyss (2002), Mapping spatial variability of the frequencymagnitude distribution of earthquakes, Adv. Geophys., 45, Wyss, M., and S. Wiemer (2000), Change in the probability for earthquakes in southern California due to the Landers magnitude 7.3 earthquake, Science, 290, C. Chiarabba, Istituto Nazionale di Geofisica e Vulcanologia, Centro Nazionale Terremoti, Via di Vigna Murata, I Rome, Italy. (chiarabba@ingv.it) L. Chiaraluce, Istituto Nazionale di Geofisica e Vulcanologia, Centro Nazionale Terremoti, Via di Vigna Murata, I Rome, Italy. (chiaraluce@ingv.it) M. Cocco and D. Piccinini, Dipartimento Sismologia e Tettonofisica, Istituto Nazionale di Geofisica e Vulcanologia, Via di Vigna Murata, I Rome, Italy. (cocco@invg.it; piccinini@ingv.it) C. Collettini, Geologia Strutturale e Geofisica, Dipartimento di Scienze della Terra, Universita degli Studi di Perugia, Piazza dell Università, I Perugia, Italy. (colle@unipg.it) 22 of 22

23 Figure 1. (a) Distribution of seismicity in the northern Apennines. Recent seismicity is distributed along the active extensional inner sector of the Apennines, where the strongest historical earthquakes are also located (white stars). Red symbols show the epicenters of the earthquakes recorded during the seismic survey (this study). Orange and blue symbols indicate the aftershocks of the 1984 Gubbio (M w 5.1) and the 1998 Gualdo Tadino (M w 5.1) earthquakes, respectively. Green symbols indicate the seismicity during the 1997 Colfiorito sequence, consisting of several normal faulting main shocks (we have plotted the focal mechanisms of the three largest shocks: M w 6.0, M w 5.7, and M w 5.6 from NW to SE, respectively). (b) Crustal-scale cross section interpretation of the CROP03 seismic profile running from the Tyrrhenian to the Adriatic coasts [Barchi et al., 1998; Collettini and Barchi, 2004]. The ATF is drawn in red, while other low-angle normal faults in the Tyrrhenian and Tuscany sectors are shown in blue. The brittle-ductile transition [Pauselli and Federico, 2002] is indicated by the light grey curve. Extension migrated with time from west to east and is accommodated by a set of east dipping low-angle normal faults and antithetic structures. 3of22

24 Figure 2. Geometry of the ATF constrained by surface geology, interpretation of seismic reflection profiles and borehole data. (a) Structural map of the study area. (b) Vertical geological cross section (see its location in Figure 2a) integrating surface geology, borehole data, and seismic reflection profiles (modified after Collettini and Barchi [2002]). (c) Geometry of the Tiber basin in a seismic profile (see its location in Figures 2a and 2b). The asymmetrical half graben shape of the basin is consistent with the presence of an east dipping normal fault. (d) ATF imaged on a seismic profile (for location, see Figures 2a and 2b). The seismic reflectors in the hanging wall block are truncated abruptly by the east dipping signals representing the ATF. Mf, Marne a Fucoidi reflector; te, top of the Triassic evaporites. 4of22

25 Figure 3 5of22