Contemporary crustal extension in the Umbria–Marche Apennines from regional CGPS networks and comparison between geodetic and seismic deformation

Tectonophysics 476 (2009) 3–12

Contents lists available at ScienceDirect

Tectonophysics j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / t e c t o

Contemporary crustal extension in the Umbria–Marche Apennines from regional CGPS networks and comparison between geodetic and seismic deformation N. D'Agostino a,⁎, S. Mantenuto a, E. D'Anastasio a, A. Avallone a, M. Barchi b, C. Collettini b, F. Radicioni c, A. Stoppini c, G. Fastellini c a b c

Istituto Nazionale Geofisica e Vulcanologia, Roma, Italy Dipartimento Scienze Geologiche, Università di Perugia, Italy Dipartimento Ingegneria Civile ed Ambientale, Università di Perugia, Italy

a r t i c l e

i n f o

Article history: Received 8 February 2008 Received in revised form 26 June 2008 Accepted 11 September 2008 Available online 30 September 2008 Keywords: GPS Crustal extension Northern Apennines

a b s t r a c t Here we report the results of the analysis of a GPS velocity field in the Umbria–Marche Apennines (central Italy) obtained from the integration of diverse geodetic networks. The velocity field obtained shows a high degree of consistency both spatially and in terms of comparison with independent information, despite the limited time span of some GPS stations. Starting from the velocity field we derive a continuous strain rate field applying a spline interpolation technique which provide a smooth estimate of the deformation field. The main feature of the resulting strain rate field is a continuous high (N 50 nanostrain/year) strain rate belt coincident with the area of largest historical and instrumental seismic release. The model directions of the principal axes agree with geological and seismological information indicating NE–SW extension. We transform the strain rate field into geodetic moment rate using the Kostrov formula to evaluate the potential seismic activity of the region and compare it with actual seismic release in the last 720 years from Mw N 5.5 earthquakes. This comparison highlights a large possible deficit in the seismic release with respect to the overall potential seismic activity, particularly concentrated in the northern part of the study area. This discrepancy can be resolved with either a large amount of seismicity to be released in the near future or significant aseismic slip and deformation. © 2008 Elsevier B.V. All rights reserved.

1. Introduction Regional and local networks of continuous GPS (CGPS) stations are spreading all around the world in most technically advanced countries, covering the territory with a growing density of stations and offering a wide variety of applications, many of which were unpredictable in the early times of satellite positioning. The high concentration of stations in small areas and the high accuracy of the GPS positioning are elements of great interest for the study of local geodynamics phenomena, particularly regional crustal deformation and associated seismic activity. One aim of this study is to integrate local regional networks developed for civil and real-time positioning applications with geodynamics-oriented GPS networks to demonstrate that as little as a few years of GPS observations can be used to characterize the tectonic deformation of seismically active regions providing important information for assessment of seismic hazard. In the past years a rapid development of GPS networks for both scientific and civil applications occurred in the Umbria and Marche

⁎ Corresponding author. Istituto Nazionale di Geofisica e Vulcanologia, Via Vigna Murata 605, 00143 Roma, Italy. Tel.: +39 06 51860537. E-mail address: [email protected] (N. D'Agostino). 0040-1951/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2008.09.033

regions (Central Italy). The Umbria–Marche Apennines are an arcshaped, NE-verging, thrust-and-fold belt corresponding to the easternmost part of the Northern Apennines. A comprehensive review of the regional geology has recently been provided by Barchi et al. (2001). The area is characterized by the presence of a complex pattern of thrusts, folds and normal faults, reflecting the superposition of two main tectonic phases: an upper Miocene-lower Pliocene compressional phase forming E–NE verging thrusts and folds, and a superimposed Quaternary extensional phase, forming intramountain basins bounded by NNW–SSE trending normal faults which offset earlier fabrics. The main geodynamic mechanism driving active extension in the Apennines is still a subject of lively debate. Some authors associate the active extension with the continuing subduction of the Adriatic lithosphere beneath the Apennines through a roll-back mechanism which would result in paired, parallel contractional (in the Adriatic) and extensional (in the Apennines) belts (Reutter et al., 1980; Doglioni, 1991; Frepoli and Amato, 1997; Jolivet et al., 1998; Basili and Barba, 2007). Other authors (Anderson and Jackson, 1987; Calais et al., 2002; D'Agostino et al., 2005, 2008) instead relate the extension in the Apennines mainly to the NE-directed motion of the Adriatic block relative to Eurasia around a pole of rotation located in the western part of the Po Plain. This rotation is consistent with the styles of active deformation observed in the orogenic belts around the Adriatic block,

4

N. D'Agostino et al. / Tectonophysics 476 (2009) 3–12

e.g. extension in the Apennines and shortening in the Eastern Alps and the Dinarides. At present, the actively extending 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 b M b 6.0) earthquakes are located (Fig. 1). The historical and instrumental seismicity do not follow the arc shape structures inherited from the previous compressional tectonic phase, but cluster along a ≈30 km wide belt following the main crest of the Apennines (Fig. 1). Furthermore the present-day extensional strain in the northern Apennines inferred from geodetic data (Hunstad et al., 2003) is concentrated across a 30–40 km wide zone that coincides with the area struck by the strongest earthquakes (Selvaggi, 1998). In the past 20 years, three main seismic sequences have occurred in the study region (Pondrelli et al., 2006): the 1979 Norcia earthquake (Mw 5.8), the 1984 Gubbio sequence (Mw 5.6), and the 1997 Colfiorito multiple main shock sequence (Mw 6.0, 5.7 and 5.6) (Fig. 1). All the main shocks have been associated to SW-dipping normal faults (Boncio et al., 2004). The current seismic hazard map in Italy is based on the historical catalogue and the subdivision into seismogenic zones characterized by homogeneous seismotectonic behaviour (Meletti et al., 2008). The evaluation of the overall active tectonic deformation and the fraction of it which is released aseismically is clearly of great importance for an improved assessment of seismic hazard.

In this paper we present the preliminary results of the analysis of a dense GPS velocity field resulting from the integration of different GNSS networks. Firstly we derive a continuous strain rate field which is used to map the distribution of active deformation. Then we define the geodetic moment rate to quantify the potential seismic activity in the region. Finally we compare the geodetically measured strain with the strain released by Mw N 5.5 earthquakes in the last 720 years. 2. GPS data 2.1. Local networks in Umbria and Central Italy The DICA (Dipartimento Ingegneria Civile Ambientale) Department at Perugia University is currently operating two local GNSS permanent stations networks (Fig. 2). The station distribution is quite dense, with an average distance of about 40 km between the stations. So far, these networks have been mostly used by surveyors, consultancies and mapping agencies, taking advantage of the different positioning services in applications related to mapping, engineering and cadastre. In addition, we are convinced of the great potential of such dense networks for the analysis of local crustal deformations, when analysed together with the monitoring of seismic activity in Umbria and surrounding areas (Fastellini et al., 2008). These two GNSS permanent stations networks are the following:

Fig. 1. Seismicity distribution and focal mechanisms of the largest instrumentally recorded earthquakes in the Northern Apennines. Instrumental seismicity is from the CSI catalogue (Castello et al., 2006). Historical seismicity is from CPTI04 catalogue (Gruppo di Lavoro CPTI04, 2004) and focal mechanisms are from the RCMT Centroid Moment Tensor catalogue (Pondrelli et al., 2006). The box outlined with a yellow dashed line limits the area used for the calculation of geodetic and seismic moment rates. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

N. D'Agostino et al. / Tectonophysics 476 (2009) 3–12

5

Fig. 2. Site locations and affiliation of the continuous GPS stations used in this work.

(1) GPSUMBRIA, the GNSS official network of the Umbria (administrative) Region, offering services both in post-processing and real time. It comprises 10 stations, covering the whole territory of Umbria. Information about this network and its data (publicly available at the moment) can be collected on the web site http://www.gpsumbria.it/. (2) LABTOPO, a local network set up for research purposes, putting together a set of GNSS permanent stations from different Universities, schools, public administrations and private companies. It presently comprises 21 stations (including all GPSUMBRIA sites), covering a wide area in central Italy. At the present time, this network only provides post-processing services. Information and data (mostly free) are available through the web site http:// labtopo.ing.unipg.it/labtopo/index.php. All antennas are installed on buildings and two GPSUMBRIA stations (UNPG— Perugia and UNTR—Terni) are part of the EUREF EPN network and the EUREF-IP real-time project. 2.2. RING network To improve understanding of current deformation and seismogenesis in Italy the INGV (Istituto Nazionale Geofisica e Vulcanologia) started in 2004 the deployment of a permanent, integrated and real-time monitoring GPS network (RING) throughout the country. RING now consists of about 120 stations (http://ring.gm.ingv.it). The largest concentrations are Southern Italy and Sicily where station spacing is of the order of 20–25 km in some target regions. Fewer stations have been installed in northern Italy because of the existing networks there. The CGPS sites, acquiring at 1 Hz and 30 s sampling rates, are integrated either with broad-band and very-broad-band seismometers with accelerometers for improved definition of the seismically active regions. Most of the sites are connected to the acquisition centers (in Rome and duplicated elsewhere) through a satellite system (VSAT), while the remaining sites transmit data by Internet and classical phone connections.

2.3. Other networks To increase station density outside the study area we included available sites from other networks. To align our network solutions in a global reference frame we included sites from IGS (http://igscb.jpl.nasa.gov), EUREF (http://www.epncb.oma.be). The Agenzia Spaziale Italiana (ASI) currently distributes about 40 stations in Italy (http://geodaf.mt.asi.it), some of which have observation intervals longer than 10 years and are included in the IGS and EUREF networks. We also include two sites (RIET, VITE) from the ItalPos network (http://www.italpos.it), which have similar characteristics to the LABTOPO and GPSUMBRIA sites. 3. Data processing Code and phase data for the time interval 1996.0–2008.0 have been reduced with the Gipsy-Oasis II software. Precise orbits and clocks from the Jet Propulsion Laboratory (ftp://sideshow.jpl.nasa.gov) were used to analyse GPS data. GPS data are processed using the precise point positioning strategy (PPP) (Zumberge et al., 1997). PPP network solutions can be further improved by the application of ambiguity resolution, however, the processing time for this step generally scales quadratically (Blewitt, 1989). Furthermore, ambiguity resolution introduces inter-station correlations that cause subsequent network kinematic analysis to scale quadratically rather than linearly. Thus some of the practical advantages of PPP are subsequently lost. Here we use the Ambizap algorithm (Blewitt, 2008) which attempts to resolve the ambiguities of n-1 baselines connecting n stations and produce a solution whose coordinates agree to bb1 mm with fullnetwork analysis (Blewitt, 2008). This approach is useful for network processing strategies that scale linearly with the number of stations. The approach allows for very rapid, multiple reanalysis of extremely large networks to assess various models and makes trivial the addition of extra stations or subnetworks to an existing solution. To avoid the approximately six-month latency of JPL transformation

6

N. D'Agostino et al. / Tectonophysics 476 (2009) 3–12

Table 1 GPS site coordinates, velocities in ITRF2005 and Eurasian reference frames and associated uncertainties (in mm/year), correlation between East and North component (CorrEN) and observation time (DT) interval (in years). GPS sites used to realize the Eurasian frame: Lon

Lat

58.560 21.035 17.073 355.503 4.359 353.658 350.582 33.991 13.730 0.492 2.052 351.411 30.497 0.336 21.032 8.411 5.810 20.670 358.781 0.155 37.224 31.973 2.587 36.569 17.274 3.865 11.280 11.280 34.543 13.066 10.460 3.879 24.059 1.481 1.720 356.048 25.299

56.430 52.476 52.277 48.380 50.798 39.479 38.693 44.413 51.030 40.821 45.403 41.106 50.364 50.867 52.097 49.011 52.178 53.892 46.159 48.019 56.028 46.973 48.841 55.115 48.373 43.637 48.086 48.086 49.603 52.379 52.296 45.044 56.949 43.561 47.294 40.444 54.653

ITRF2005

Eurasia

SigE SigN CorrEN STA

DT

Ve

Vn

Ve

Vn

24.84 20.64 19.64 16.12 17.21 18.37 17.75 24.36 19.58 19.65 19.04 17.68 22.13 18.71 20.59 19.06 17.84 20.45 18.03 17.18 21.88 23.38 17.99 22.79 20.98 19.67 20.13 19.86 22.77 18.94 18.57 18.90 19.85 19.36 18.28 18.96 20.74

6.33 14.12 14.68 16.76 15.72 17.13 16.55 12.46 15.80 16.09 15.81 16.38 12.98 15.91 14.33 15.81 16.42 13.13 16.11 16.66 11.54 12.44 15.92 12.03 15.36 15.91 15.97 15.90 12.73 15.07 15.98 16.30 13.44 15.89 15.54 16.84 13.68

− 0.68 − 0.02 − 0.28 − 0.29 − 0.42 − 0.12 − 0.39 − 0.12 0.04 0.24 0.45 0.03 − 0.75 1.98 − 0.16 0.10 0.29 0.21 0.30 − 0.32 − 0.93 − 0.38 0.20 − 0.11 0.10 0.29 0.36 0.09 − 0.93 − 0.12 0.04 − 0.14 − 0.30 0.42 0.26 0.28 − 0.23

− 0.10 − 0.14 − 0.10 0.44 − 0.24 0.78 0.20 0.36 0.64 − 0.07 − 0.28 0.02 0.23 − 0.26 0.07 0.14 0.56 − 1.18 − 0.12 0.48 0.08 − 0.04 − 0.14 0.44 0.60 − 0.08 0.56 0.49 0.73 − 0.16 0.49 0.31 − 0.38 − 0.23 − 0.57 0.53 0.05

0.25 0.09 0.08 0.38 0.12 0.25 0.19 0.23 0.18 0.08 0.17 0.24 0.16 0.41 0.10 0.16 0.10 0.19 0.27 0.18 0.47 0.24 0.15 0.23 0.12 0.08 0.24 0.25 0.19 0.06 0.15 0.24 0.18 0.14 0.15 0.10 0.26

0.19 0.09 0.07 0.36 0.08 0.17 0.13 0.17 0.13 0.09 0.17 0.16 0.12 0.20 0.10 0.14 0.07 0.19 0.22 0.13 0.32 0.15 0.12 0.20 0.14 0.07 0.17 0.45 0.13 0.05 0.12 0.16 0.15 0.12 0.12 0.08 0.21

− 0.12 0.07 0.11 0.01 0.00 0.10 0.18 − 0.14 0.23 0.05 − 0.15 0.15 0.12 − 0.68 − 0.05 0.16 0.01 − 0.48 0.29 0.01 0.01 − 0.05 0.04 − 0.06 − 0.13 0.03 − 0.08 0.10 − 0.04 0.00 − 0.00 0.10 − 0.00 0.04 0.06 0.12 − 0.09

ARTU BOGO BOR1 BRST BRUS CACE CASC CRAO DRES EBRE EGLT GAIA GLSV HERS JOZE KARL KOSG LAMA LROC MANS MDVO MIKL MLVL MOBN MOPI MTPL OBE2 OBER POLV POTS PTBB PUYV RIGA TLSE VFCH VILL VLNS

8.4 9.7 12.0 9.0 12.0 6.1 8.8 7.7 8.0 10.7 6.2 7.0 9.7 11.0 12.0 9.1 12.0 12.0 6.1 9.7 7.1 5.6 7.1 6.8 9.7 8.7 6.3 4.5 6.5 12.0 7.8 3.3 7.2 6.8 6.3 12.0 5.3

2.39 3.32 3.15 2.17 0.00 3.39 3.90 2.81 1.68 3.26 2.91 2.15 1.48 1.86 3.16 2.15 1.86 2.09 1.71 2.03 1.44 0.63 2.89 1.49 2.50 1.50 0.77 1.82 3.79 2.72 1.27 3.67

0.34 0.44 0.58 0.21 1.26 0.32 0.38 0.42 0.31 0.28 0.49 0.42 0.22 0.40 0.38 0.40 0.40 0.29 0.35 0.09 0.20 0.14 0.20 0.21 0.19 0.11 0.34 0.36 0.31 0.16 0.38 0.42

0.70 0.48 0.43 0.17 1.28 0.26 0.24 0.32 0.47 0.29 0.36 0.29 0.25 0.24 0.53 0.20 0.24 0.32 0.34 0.08 0.16 0.10 0.27 0.17 0.15 0.18 0.20 0.22 0.22 0.15 0.36 0.34

0.04 0.08 0.28 0.06 0.20 0.13 0.15 0.01 − 0.08 0.10 0.21 − 0.00 0.10 0.09 0.01 0.01 0.09 − 0.00 0.03 0.13 − 0.10 − 0.12 0.08 0.17 0.51 0.09 0.07 0.02 0.17 − 0.04 − 0.25 0.00

IEMO UNUB REMO CAIE COIE ITFA MOIE RENO UNTR ITRA ITGT REPI UNOV RETO CSGP REFO RETO VITE RIET PRAT UNPG ELBA CAME M0SE AQUI BRAS MAON TOLF INGP RSTO MURB RSMN

2.5 2.2 2.1 4.8 1.4 2.5 2.8 2.8 3.5 2.8 3.5 2.6 3.3 2.6 2.3 2.7 2.6 2.5 2.5 9.7 9.6 7.2 7.7 5.9 8.6 7.6 2.7 3.8 4.3 6.3 3.5 3.1

GPS sites in the study area and others site: 12.053 12.640 12.226 12.248 11.996 12.926 13.123 13.093 12.674 14.002 12.782 12.002 12.113 12.407 13.592 12.704 12.407 12.120 12.857 11.099 12.356 10.211 13.124 12.493 13.350 11.113 11.131 12.000 13.316 14.002 12.525 12.451

43.592 43.700 43.453 43.467 43.279 43.344 43.503 42.793 42.559 42.658 43.234 42.952 42.716 42.782 42.855 42.956 42.782 42.418 42.408 43.886 43.119 42.753 43.112 41.893 42.368 44.122 42.428 42.064 42.383 42.658 43.263 43.934

20.81 22.12 22.61 21.53 21.22 21.65 22.63 23.10 21.75 22.84 24.61 21.06 21.65 21.12 22.98 21.87 21.12 21.12 21.49 21.30 20.28 20.48 22.74 21.15 21.83 21.16 20.64 20.07 23.64 23.00 23.80 21.86

17.73 18.60 18.47 17.49 18.11 18.64 19.13 18.04 16.95 18.39 18.17 17.49 16.81 17.16 18.34 17.42 17.16 17.42 16.96 17.46 16.75 16.14 18.12 16.78 17.70 16.93 16.20 17.16 19.00 17.85 16.56 18.97

− 0.13 1.10 1.61 0.53 0.00 0.50 1.48 1.81 0.49 1.36 3.46 0.00 0.52 − 0.05 1.61 0.68 − 0.05 − 0.07 0.16 0.60 − 0.81 − 0.30 1.51 − 0.22 0.41 0.51 − 0.37 − 1.17 2.22 1.52 2.71 0.92

Table 1 (continued) used realize theand Eurasian GPS sites in thetostudy area others frame: site: Lon 12.606 16.704 8.973 7.661 8.921 14.990 11.878 12.332 7.465 6.604

Lat 35.500 40.649 39.136 45.063 44.419 36.876 45.407 45.437 46.877 52.915

ITRF2005

Eurasia

Ve

Vn

Ve

19.74 23.33 21.56 20.02 20.47 21.16 20.62 21.53 19.45 17.43

18.50 − 2.78 19.19 1.00 15.89 0.27 15.83 0.24 15.74 0.30 19.78 − 1.51 17.58 0.11 17.16 0.94 16.14 0.15 16.24 − 0.09

SigE SigN CorrEN STA

DT

0.17 0.10 0.11 0.13 0.08 0.24 0.14 0.25 0.09 0.12

8.8 12.0 11.8 10.8 9.4 7.3 5.7 10.9 12.0 10.5

Vn 3.22 4.37 0.27 0.10 0.12 4.76 2.23 1.85 0.40 0.43

0.09 0.09 0.10 0.10 0.09 0.12 0.12 0.30 0.06 0.10

0.10 − 0.01 − 0.03 − 0.02 − 0.05 0.04 − 0.01 − 0.08 − 0.04 − 0.05

LAMP MATE CAGL TORI GENO NOT1 UPAD VENE ZIMM WSRT

parameters (x-files) between the daily satellites-fix and the ITRF2005 (Altamimi et al., 2007) reference frames, we constructed a “regional” ITRF2005 coordinate and velocities solution for the 100 older and more stable sites in our network using weekly averages and ITRF2005 x-files from JPL. This solution has been used to align the final daily ambiguityfixed solutions from Ambizap to the ITRF2005 reference frame using a 7-parameter Helmert transformation. Time series in the ITRF2005 reference frame are locally rotated to north–east-up components and cleaned from outliers using the strategy described in Nikolaidis (2002). The time series are analyzed for their noise properties, linear velocities, periodic signals and antenna jumps using maximum likelihood estimation (MLE). This technique allows simultaneous estimation of the noise properties structure together with the parameters of a time dependent model of the data. Quantities estimated in the MLE analysis are linear trend, offset at designated times, annual and semi-annual periodic signals, power law noise index and amplitude, and white noise amplitude (Williams, 2003). The MLE analysis has been performed using the CATS software (Williams, 2007) and all parameters estimated with full white noise plus flicker noise covariance, which previous studies suggest to be the more common noise model in GPS time series (Williams, 2003). The uncertainties in parameters can thus be considered realistic estimates based on analysis of the noise at individual stations. Station positions, horizontal linear velocities in various reference frames and associated uncertainties, time span of the observations and number of occupations are listed in Table 1. Fig. 3 shows the distribution of scatter of postfit residuals for GPSUMBRIA and LABTOPO sites compared with the other sites used in this work. Average postfit residuals are 1–2 mm for north and east components and about 5 mm for the vertical component. No significant difference appears between the two sets of sites. The Eurasia reference frame has been realized by minimizing the horizontal velocities of 36 sites located on the stable part of the plate. The RMS of the horizontal velocities after the least-squares linear inversion are 0.49 and 0.41 mm/year for the East and North components respectively, approximately at the same level of uncertainty of horizontal velocities suggesting the robustness of the Eurasian frame realization (Table 2). 4. GPS velocity field The GPS velocity field in the Eurasia reference frame is shown in Fig. 4, with velocities for all stations with over 2 years of data coverage. At present our results should be interpreted with caution. Some of the presented site velocities have been determined from data spanning ~2 years. Ideally, all velocities should be derived from at least 2.5 years of data in order for seasonal signals not to significantly affect the estimation of velocities for the linear part of the signal (Blewitt and Lavallée, 2002). Nevertheless, the spatial consistency of the velocity field and the similarity in velocities from nearby GPS sites with different observation intervals (ITRA, RSTO) is encouraging. Significant differences exist between the monument types within the area. Antennas from the Labtopo, ASI and Italpos networks are generally

N. D'Agostino et al. / Tectonophysics 476 (2009) 3–12

7

Fig. 3. Histograms of scatter of velocity postfit residuals for GPSUMBRIA and LABTOPO sites (black bars) compared to the other sites used in this work (grey bars).

monumented on steel bars or light concrete pillars on buildings. Antennas from the RING network are installed on concrete pillars on bedrock. Despite the differences in monumentation no significant difference appears in the precision of velocity estimates. The main feature of the velocity field is given by a general N-directed motion on the Tyrrhenian (southwestern) side at 1–2 mm/year and a NE-directed motion at 3–4 mm/year on the Adriatic (northeastern) side of the Apennines. A sharp divergence is seen along the crest of the Apennines occurring in a narrow region which follows the Apennines orographic divide and results in net extensional strain localized in the region of largest historical and instrumental seismic release. This pattern confirms early analyses based on survey-type GPS results (D'Agostino et al., 2001) and remeasurement of the historical triangulation network Hunstad et al. (2003). Velocity profiles constructed across the Umbria–Marche Apennines (Fig. 5) show that the net rate of extension ranges between 2–3 mm/ year in agreement with the previous 2.5 mm/year estimate of Hunstad et al. (2003) and Serpelloni et al. (2005). This extension is similar in direction and magnitude to the motion predicted by the rotation of the Adria block using GPS sites in the Po Plain (D'Agostino et al., 2005). On the Adriatic side the easternmost site velocities align with the velocity predicted by the rigid rotation of the Adriatic microplate relative to Eurasia (D'Agostino et al., 2008) consistent with the interpretation that the northeastern motion of the Adriatic block imposes the boundary conditions responsible for the Apennines extension. 5. Strain rate distribution Calculation of the strain rate tensor provides a valuable tool to map the distribution of active deformation and to define the areas of highest strain accumulation. The strain rates evaluated from GPS velocity fields, in the absence of significant earthquakes, is generally interpreted as reflecting the interseismic elastic loading on faults in the upper seismogenic crust, which may be potentially released seismically in the future (Holt et al., 2000; Pancha et al., 2006). Various approaches to derive a continuous strain rate fields have been proposed in the literature. In order to relate the geodetic velocities to crustal deformation rates we model the crustal horizontal velocity field under the assumption that the crust deforms as a continuum (Haines and Holt, 1993; England and Molnar, 1997) and a smooth estimate of the Table 2 Parameters of the ITRF2005-Eurasia Eulerian pole. Lat

Lon

Ω (deg/My)

σmax

σmin

θ

σΩ

− 55.160

81.717

− 0.257

0.3

0.1

− 46.0

0.001

RMS (mm/year) East

North

0.49

0.41

Ω is the rotation rate (°/My), σmax and σmin are the semi-axes of the 1-σ error ellipse, θ is the azimuth of σmax (clockwise from North), σΩ is the 1-σ rotation rate uncertainty. Also shown are the RMS of the residuals of horizontal velocities.

deformation field is provided. To derive continuous velocity gradient tensor field we apply a spline interpolation technique modified from the approach of Haines and Holt (1993). In order to get a continuous strain rate map of the region, we use a “spline in tension” technique (Wessel and Bercovici, 1998) to interpolate the GPS velocities on 0.1°× 0.1° longitude and latitude grids, and then calculate the horizontal velocity gradient tensor δvi/δxj. The solution curvature is controlled by the tension factor (0≤T ≤ 1). Fig. 5 shows the observed velocities together the velocity interpolation obtained with values of the tension parameter T ranging between 0 and 1. T = 0.5 appears to offer the optimal value to minimise short wavelength noise. However some steep velocity gradients in profile C seem to be related to noise in site velocities (UNPG, ITFA). From the components of the velocity gradient tensor we derive the strain rate tensor (0.5 ⁎ (δvi/δxj + δvj/δxi)), its principal axes and the second invariant of the strain rate tensor. This method produces a reliable strain rate field, especially for the area where the distribution of observed velocities is dense. Fig. 6 shows the distribution of the second invariant of the strain rate tensor and the direction and magnitude of its principal axes using a tension factor T = 0.5. The main feature of the resulting strain rate field is a continuous band of high values (N50× 10− 9 year− 1) parallel to the Apennines. This belt of active deformation corresponds with the clear divergence observed in the velocity field (Fig. 4) between N–NW oriented vectors on the western side and NNE oriented vectors on the eastern side of the Apennines, and appears to correlate with the distribution of historical and instrumental seismicity (Fig. 1) as well as the distribution of known active faults (Boncio et al., 2004). The principal axes of the strain rate tensor indicate almost uni-axial extension which is also shown by focal mechanisms and geological data (Boncio et al., 2004). Observed strain rates on the Adriatic side of the Apennines remain below 20 × 10− 9 year −1 and do not show clear evidence of active shortening on land as predicted by the seismotectonic zoning of Lavecchia et al. (1994) and by the geomorphological study of Vannoli et al. (2004). The lack of significant shortening on the Adriatic side may also result from a paucity of GPS sites on the coast and associated reduced resolution. It is also possible that significant shortening now occurs off-shore. Further investigations aimed at the evaluation of seismic hazard related to active shortening in this area are in progress and will require a denser station density. Some short wavelength features in the strain rate field seem to be related to noise in the velocity field. In particular the high local strain centered at 13.0°E 43.25°N seems to be related to higher station density which enhances short wavelengths in the strain rate field. A large smoothing factor in the interpolation scheme would reduce this short wavelength noise but would spatially spread the strain rate in the regions with lower station density. The maximum strain rate is observed north of Perugia in the AltoTiberina Valley. Here a NE ∼20°-dipping normal fault of regional importance (ATF, Alto-Tiberina Fault) has been proposed on the basis of surface geology, seismic reflection lines and microseismicity

8

N. D'Agostino et al. / Tectonophysics 476 (2009) 3–12

Fig. 4. Geodetic velocities in the Eurasian reference frame and associated 95% confidence error ellipse. Colored dots for each site relate to the data time span as indicated by the bar on the right. The dashed lines show the trace of the velocity profiles in Fig. 5.

(Boncio et al., 2000; Chiaraluce et al., 2007). Current aseismic slip on this structure has been proposed by Collettini and Holdsworth (2004) and by Chiaraluce et al. (2007).

To compare geodetic deformation with the seismically released deformation (see next section) we defined a box which includes the area with optimal station density and geometry (boxes in Figs.1 and 6).

Fig. 5. GPS velocities projected along three profiles whose traces are shown in Fig. 4. The dashed line is the model velocity obtained from continuous interpolation of observed GPS velocities using different values of the tension factor T.

N. D'Agostino et al. / Tectonophysics 476 (2009) 3–12

9

Where μ = 3 × 1010 N/m2, A c and H c are the area and the seismogenic thickness of the 0.1° × 0.1° cell respectively, and ε1̇ i and ε̇2i are the eigenvalues of the strain rate tensor of the ith cell. The range of values so obtained are listed in Table 3 and varies from 1.35 × 1020 to 3.35 × 1020 Nm. 6. Seismic deformation and comparison with geodetic strain rates

Fig. 6. Map of the second invariant and principal axes of the model strain rate tensor obtained from the interpolation of the velocity field of Fig. 3 using a tension factor T = 0.5. Observed GPS velocities are shown as red arrows with 95% CI ellipse errors. The box shown with a dashed line is the same of that shown in Fig. 1 and includes the area used for summation of scalar seismic moment for earthquakes in the last 720 years.

Strain rates ε̇ measured over an area A can be translated into rates of potential moment release Ṁ 0 by Kostrov's (1974) formula Ṁ 0 = 2μAHε̇, where μ is rigidity (3 × 1010 N/m2) and H is the seismogenic thickness. Note that the tensor form of the equation is reduced here to scalar form. Acknowledging the nonuniqueness and uncertainty involved with converting surface strain rate to a scalar moment rate (Savage and Simpson, 1997), this study utilizes two methods to help quantify the moment rate from geodesy. The first uses the summation of the largest eigenvalue of the strain rate tensor (Ward, 1994) while the second uses the formulation of Savage and Simpson (1997). A critical factor is the parameter H, the depth of the seismogenic layer. A possible approach is the analysis of depth distributions of earthquakes from the Italian instrumental catalogue (Castello et al., 2006) analysed in the same box used for the strain rate calculation. To increase the reliability of the selected dataset we only used events with formal error in depth less than 5 km, at least ten P-waves arrival times, an azimuth gap below 180°, and RMS of the solution below 0.8 s. The cumulative frequency plot of depth distribution of crustal earthquakes (Fig. 7) shows that approximately 90% of events occur at depths less than 12.5 km. This depth is greater than the depths of the largest instrumentally recorded earthquakes in the area (Haessler et al., 1988; Amato et al., 1998). We therefore evaluate the geodetic moment rate for a range of seismogenic thicknesses (7.5 km, 10 km and 12.5 km). Considering the impact of the tension factor in the velocity interpolation scheme we also varied T from 0.0 to 1.0. To obtain the geodetic moment rates summed over 720 years we summed the strain rate contribution of each of the n 0.1° × 0.1° cells within the box of Fig. 6 using the equations: n  : : X :i  i M0 = 2μHc Ac Max je1 j ; je2 j

The seismic deformation is evaluated here by summing the contribution of the earthquakes contained within the same box used for the evaluation of the geodetic moment rate. We select the contributing earthquakes from the catalogue CPTI04 (Gruppo di Lavoro CPTI, 2004). To quantify the contribution of every earthquake we used the estimates of moment magnitude from macroseismic intensity distributions given in the CPTI04 catalogue. Details of the estimation of moment magnitudes from macroseismic information are given in Gruppo di Lavoro (2004) and Gasperini and Ferrari (2000). To increase the completeness of the extracted dataset we selected (Table 4) those earthquakes with estimated Mw N 5.5 only in the 720 years time interval (1277–1997). Historical and statistical analyses (Gruppo di Lavoro CPTI, 2004) suggest, however, that in the study area, the historical catalogue is probably complete for Mw in the range 5.4–5.7 from A.D. 1500–1700. Fig. 8 shows that the cumulative frequency distribution of the selected earthquakes as a function of magnitude is well reproduced by a Gutenberg–Richter function with a b value close to unity (b = 1.02). Note, however, that possible incompleteness in the lower magnitude range was not taken into account. We converted the moment magnitude Mw provided by the catalogue using the formula M0 = 10(1.5Mw + 9.1) (Hanks and Kanamori, 1979) to obtain scalar seismic moments (in Nm). Mw magnitudes for earthquakes after 1979 were taken from the Regional Centroid Moment Tensor Catalogue (RCMT) of Pondrelli et al. (2006). Fig. 9 shows the cumulative seismic moment release with time over the study region. The cumulative moment is clearly dominated by the two events of the 1703 seismic sequence which ruptured a large part of the active fault

Ward (1994) ð1Þ

i=1

n  : : X :i :i :i  i M 0 = 2μHc Ac Max je1 j ; je2 j ; je1 + e2 j Savage and Simpson (1997) ð2Þ i=1

Fig. 7. Cumulative frequency plot of depth distribution of crustal earthquakes (CSI catalogue; Castello et al., 2006) selected from the box used for the calculation of geodetic and seismic moment rates (Fig. 1). Selection criteria include a location RMS of less than 0.8 s, a vertical formal error b 5 km, at least 10 P-wave travel time arrivals and an azimuthal gap less than 200°. About 90% of the seismicity occurs above 12.5 km.

10

N. D'Agostino et al. / Tectonophysics 476 (2009) 3–12

Table 3 Geodetic moment rates cumulated over 720 years (Nm × 1020) from the method of Ward (1994) and Savage and Simpson (1997). Method

H = 7.5 km

H = 10.0 km

H = 12.5 km

T = 0.0 T = 1.0 T = 0.0 T = 1.0 T = 0.0 T = 1.0 Ward (1994) : M 0 = 2μHA

Pn i= 1

 :i :i  Max je1 j; je2 j

Savage and Simpson (1997) : M 0 = 2μHA

Pn

i =1

:i :i :i :i  Max je1 j; je2 j ; je1 + e2 j

1.92

1.35

2.56

1.80

3.20

2.25

2.01

1.43

2.68

1.90

3.35

2.38

A is the area of a 0.1° × 0.1° cells, H is the seismogenic thickness; T is the tension factor used in the interpolation of the horizontal velocity field; μ = 3 × 1010 N/m2, and ε1̇ i and ε̇2i are the eigenvalues of the strain rate tensor.

system along this section of the Northern-Central Apennines and released an estimated scalar seismic moment of 0.33 × 1020 Nm. Allowing a conservative 30% uncertainty in the estimate of M0 derived from historical sources, the seismic moment cumulated in the last 720 years (in the range 5.5b Mwb 6.81) is 0.65 ± 0.20 × 1020 Nm. To take into account the seismic moment released by Mwb 5.5 earthquakes we follow Molnar (1979) and combine the frequency-magnitude relation (log N(M) = a − bM) with the moment–magnitude relation (log M0 =cM+ d) to derive the seismic moment released by earthquakes with a seismic moment M0 ≤Mmax : 0 Σ

M0 =

α max1 − β M 1− β 0

ð3Þ

where a α = 10ð

+

bd c

Þ

and β = b = c

ð4Þ

We use Eq. (3) to estimate the moment we are missing from failing to account for Mw b 5.5 earthquakes. Assuming c = 3/2 (Hanks and Kanamori, 1979) and b = 1 we estimate that the proportion of the total seismic moment we are missing by failing to account for the events Mw b 5.5 is 22%, providing a corrected estimate of the seismic moment cumulated in the last 720 years (0.83 ± 0.27 × 1020 Nm). Considering

Table 4 Earthquake selected from the CPTI04 catalogue (Gruppo di Lavoro CPTI04, 2004) for the sum of seismic moment tensors. Year

Lon

Lat

Mw

M0 (Nm)

1277 1279 1298 1328 1352 1389 1458 1599 1639 1703 1703 1730 1747 1751 1781 1789 1799 1832 1838 1859 1878 1917 1979 1984 1997 1997 1997

12.73 12.87 12.83 13.01 12.12 12.29 12.23 13.01 13.25 13.12 13.20 13.11 12.82 12.73 12.50 12.20 13.12 12.65 12.88 13.09 12.67 12.12 13.06 12.57 12.88 12.81 12.92

42.73 43.09 42.55 42.85 43.46 43.52 43.45 42.71 42.63 42.68 42.47 42.75 43.20 43.22 43.59 43.50 43.14 42.96 42.87 42.82 42.85 43.46 42.81 43.27 43.05 43.08 42.93

5.57 6.33 5.93 6.44 6.00 6.00 5.87 5.82 6.26 6.81 6.65 5.85 5.93 6.30 6.23 5.80 5.93 5.80 5.63 5.70 5.55 5.80 5.80 5.60 5.70 6.00 5.60

2.85e+ 17 3.94e+ 18 9.89e+ 17 5.75e+ 18 1.26e+ 18 1.26e+ 18 8.04e+ 17 6.76e+ 17 3.09e+ 18 2.07e+ 19 1.19e+ 19 7.50e+ 17 9.89e+ 17 3.55e+ 18 2.79e+ 18 6.31e+ 17 9.89e+ 17 6.31e+ 17 3.51e+ 17 4.47e+ 17 2.66e+ 17 6.31e+ 17 6.31e+ 17 3.16e+ 17 4.47e+ 17 1.26e+ 18 3.16e+ 17

Starting from 1979 values of Mw are taken from Pondrelli et al. (2006).

Fig. 8. Cumulative magnitude-frequency distribution of seismicity in the last 720 years. The dashed line shows the Gutenberg–Richer distribution best-fit to the observations and its b value.

the evidence reported above that the Mw N 5.5 earthquake population in the last 720 years can be incomplete (Gruppo di Lavoro, 2004), this value of corrected seismic deformation can still be underestimated. Considering the associated uncertainties in the evaluation of both seismic and geodetic moment rates there appears to be a significant difference between the two. However the lower bound of the geodetic moment overlaps with the 95% confidence interval upper bound of the seismic moment (Fig. 10) implying no deficit in seismic release. Alternatively the upper bound of the geodetic moment and the lower 95% confidence interval of the seismic moment define a large (3.06 × 1020 Nm) deficit in the seismic release. The upper bound of geodetic deformation results from the combination of the largest seismogenic thickness (H = 12.5 km) with the lower tension factor (T = 0.0). The complete seismic release of this discrepancy would need about 330 Mw 5.5 and 33 Mw 6.5 earthquake. This value does not

Fig. 9. Plot of cumulative seismic moment release with time over the area defined by the box in Fig. 1. The cumulative value is dominated by the moment release of the 1703 earthquake events (3.26 × 1019 Nm) in the southern part of the selected box.

N. D'Agostino et al. / Tectonophysics 476 (2009) 3–12

Fig. 10. Comparison of the range of geodetic and seismic moment in the last 720 years. The seismic moment is obtained summing the contribution of Mw N 5.5 earthquakes and correcting for Mw b 5.5 earthquakes (see details in the text). The upper and lower bounds represent the 95% confidence interval. The range of geodetic moment (Table 3) is obtained summing the geodetic strain rates from Fig. 6 varying the seismogenic thickness H and the tension factor T used for velocity interpolation.

appear realistic considering the past release of seismicity (the seismicity should have been anomalously low in the last 720 years ) and the lack of paleoseismological evidence for such high rate of Mw 6.5 events. We thus consider this upper bound of geodetic deformation unlikely. Reducing the discrepancy to 1.47 × 1020 Nm by considering the median value of both seismic and geodetic intervals, decreases the number of required Mw 6.5 and 5.5 earthquakes to 16 and 160 respectively. Considering that the 1703 events occurred in the southern part of the box (Fig. 1) it is possible to hypothesize an along strike variation of seismic release properties of this part of the Apennines with a larger fraction of accumulated deformation released aseismically in the northern part of the box of Fig. 1. In this area previous authors have proposed aseismic slip on the Alto-Tiberina fault (Chiaraluce et al., 2007). The high strain rates measured in this area are consistent with a shallow locking depth and aseismic slip on this structure, but further investigations are needed to put firmer constraints on the activity of this fault and its seismogenic behaviour. A significant discrepancy between seismically-released and geodetic moment has already been documented by Hunstad et al. (2003) for the whole Apennines using the strain rates derived from remeasurements of the Italian first-order triangulation network. In that work a general consistency in orientation between the geodetic and seismic estimates of principal strain rates was observed, but the combined uncertainties in the magnitudes of seismic and geodetic strains were narrow enough to indicate significant discrepancies in strain magnitude all along the Apennine chain. 7. Conclusion We present geodetic velocities resulting from a combination of different GPS networks in the Umbria–Marche Apennines. Although our results should be interpreted with caution because of the short time span of some of the data, the velocity field and the derived strain rates are consistent with independent seismic and geologic information on the contemporary distribution of active deformation. We show that integration of GPS networks developed with different purposes is a powerful approach to obtain a dense GPS velocity field and map the distribution of the tectonic strain rate with few years of observations. The velocity and strain rate results confirm several known features of the Umbria–Marche Apennines active deformation, such as the dominant uni-axial extensional character of the deformation regime,

11

the limited width of the actively extending region and the minor role of active shortening in the on-shore area of the Adriatic side of the Apennines. Scalar seismic moment released by Mw N 5.5 earthquakes over the last 720 years amounts to 0.65 ± 0.20 × 1020 Nm with a dominant contribution from the 1703 earthquake events. Allowing for Mw b 5.5 earthquakes increases this value to 0.83 ± 0.27 × 1020 Nm. This overall budget highlights a significant deficit in seismic release, which may be resolved with (1) the incomplete catalogue, (2) a large number of 5.5 b Mw b 6.5 earthquakes in the future, (3) a significant fraction of the geodetic moment rate being released aseismically or a combination of the previous factors. Considering that the 1703 events occurred in the southern part of the geographic box used to compare geodetic and seismic moment rates, a larger fraction of aseismically released deformation may occur in the northern part of the box. Various sources of uncertainty still affect the estimation of both geodetic and seismic moment rates. In regions of high spatial station density our approach maps the distribution of active deformation emphasising the need for dense GPS network crossing the Apennines for high-resolution imaging of strain rate concentrations. This calls for continuing efforts to integrate GPS networks of various types in actively deforming regions. Acknowledgments We thank the Guest Editor Gerald Roberts, James Jackson, an anonymous reviewer and John Haines for the exhaustive reviews of our manuscript and for their helpful comments. References Altamimi, Z., Collilieux, X., Legrand, J., Garayt, B., Boucher, C., 2007. ITRF2005: a new release of the International Terrestrial Reference Frame based on time series of station positions and Earth Orientation Parameters. J. Geophys. Res. 112, B09401. doi:10.1029/2007JB004949. Amato, A., et al., 1998. The 1997 Umbria–Marche, Italy, earthquake sequence; a first look at the main shocks and aftershocks. Geophys. Res. Lett. 25, 2861–2864. Anderson, H., Jackson, J., 1987. Active tectonics of the Adriatic region. Geophys. J. R. Astron. Soc. 91, 937–983. Barchi, M., Landuzzi, A., Minelli, G., Pialli, G., 2001. Outer northern Apennines, in anatomy of an orogen: the Apennines and Adjacent Mediterranean Basins. In: Vai, G., Martini, I. (Eds.), Kluwer Academic Publishers, Great Britain, pp. 215–254. Basili, R., Barba, S., 2007. Migration and shortening rates in the northern Apennines, Italy: implications for seismic hazard. Terra Nova 19, 462–468. doi:10.1111/j.13653121.2007.00772.x 2007. Blewitt, G., Lavallée, D., 2002. Effect of annual signals on geodetic velocity. J. Geophys. Res. 107 (B7). doi:10.1029/2001JB000570. Blewitt, G., 1989. Carrier phase ambiguity resolution for the global positioning system applied to geodetic baselines up to 2000 km. J. Geophys. Res. 94 (B8), 10187–10203. Blewitt, G., 2008. Fixed-Point Theorems of GPS Carrier Phase Ambiguity Resolution and Their Application to Massive Network Processing: “Ambizap”. J. Geophys. Res. 113, B12410. Boncio, P., Brozzetti, F., Lavecchia, G., 2000. Architecture and seismotectonics of a regional low-angle normal fault zone in central Italy. Tectonics 19 (6), 1038–1055. Boncio, P., Lavecchia, G., Pace, G., 2004. Defining a model of 3-D seismogenic sources for seismic hazard assessment applications: the case of central Apennines. J. Seismol. 8, 407–425. Boschi, E., Guidoboni, E., Ferrari, G., Valensise, G., 1998. I terremoti dell'Appennino Umbro-Marchigiano (area sud orientale dal 99 a.C. al 1984)-ING-SGA. Compositori, Bologna, Italy. 267 pp. Calais, E., Nocquet, J.M., Jouanne, F., Tardy, M., 2002. Current strain regime in the western Alps from continuous Global Positioning System measurements, 1996– 2001. Geology 7, 651–654. Castello, B., Selvaggi, G., Chiarabba, C., Amato, A., 2006. CSI Catalogo della sismicità italiana 1981–2002, ver.1.1. INGV-CNT, Roma. http://legacy.ingv.it/CSI/. Chiaraluce, L., Chiarabba, C., Collettini, C., Piccinini, D., Cocco, M., 2007. Architecture and mechanics of an active low-angle normal fault: Alto Tiberina Fault, northern Apennines, Italy. J. Geophys. Res. 112, B10310. doi:10.1029/2007JB005015. Collettini, C., Holdsworth, R.E., 2004. Fault zone weakening processes along low-angle normal faults: insights from the Zuccale Fault, Isle of Elba, Italy. J. Geol. Soc. 161, 1039–1051. D'Agostino, N., Cheloni, D., Mantenuto, S., Selvaggi, G., Michelini, A., Zuliani, D., 2005. Strain accumulation in the southern Alps (NE Italy) and deformation at the northeastern boundary of Adria observed by CGPS measurements. Geophys. Res. Lett. 32, L19306. doi:10.1029/2005GL024266. D’Agostino, N., Avallone, A., Cheloni, D., D’Anastasio, E., Mantenuto, S., Selvaggi, G., 2008. Active tectonics of the Adriatic region from GPS and earthquake slip vectors. J. Geophys. Res. 113, B12413.

12

N. D'Agostino et al. / Tectonophysics 476 (2009) 3–12

D'Agostino, N., Giuliani, R., Mattone, M., Bonci, L., 2001. Active crustal extension in the Central Apennines (Italy) inferred from GPS measurements in the interval 1994– 1999. Geophys. Res. Lett. 28 (10), 2121–2124. Doglioni, C., 1991. Proposal for the kinematic modelling of W-dipping subduction, possible applications to the Tyrrhenian–Apennines system. Terra Nova 3, 423–434. England, P., Molnar, P., 1997. The field of crustal velocity in Asia calculated from Quaternary rates of slip on faults. Geophys. J. Int. 130, 551–582. Fastellini, G., Radicioni, F., Stoppini, A., 2008. mpact of Local GNSS Permanent Networks in the Study of Geodynamics in Central Italy. XXIV IUGG General Assembly, Perugia, Luglio 2007. IAG Symposia 133, 549–555. Frepoli, A., Amato, A., 1997. Contemporaneous extension and compression in the northern Apennines from earthquake fault-plane solutions. Geophys. J. Int. 129, 368–388. Gasperini, P., Ferrari, G., 2000. Deriving numerical estimates from descriptive information: the computation of earthquake parameters. Ann. Geofis. 43 (4), 729–746. Gruppo di Lavoro CPTI, 2004. Catalogo Parametrico dei Terremoti Italiani, versione 2004 (CPTI04), INGV, Bologna. (http://emidius.mi.ingv.it/CPTI04). Gruppo di Lavoro, 2004. Redazione della mappa di pericolosità sismica prevista dall'Ordinanza PCM 3274 20/03/2004. Rapporto Conclusivo, INGV, Milano-Roma. 65 pp., Rapporto Conclusivo (http://zonesismiche.mi.ingv.it). Haessler, H., Gaulon, R., Rivera, L., Console, R., Frogneux, M., Gasparini, G., Martel, L., Patau, G., Siciliano, M., Cisternas, A., 1988. The Perugia (Italy) earthquake of 29, April 1984: a microearthquake survey. Bull. Seismol. Soc. Am. 78, 1948–1964. Hanks, T.C., Kanamori, H., 1979. A moment magnitude scale. J. Geophys, Res. 84 (5), 2348–2350. Haines, A.J., Holt, W.E., 1993. A procedure for obtaining the complete horizontal motions within zones of distributed deformation from the inversion of strain rate data. J. Geophys. Res. 98, 12057–12082. Holt, W.E., Shen-Tu, B., Haines, J., Jackson, J., 2000. On the determination of selfconsistent strain rate fields within zones of distributed deformation. In: Richards, M.A., Gordon, R.G., van der Hilst, R.D. (Eds.), The History and Dynamics of Global Plate Motions. AGU, Washington, D.C., pp. 113–141. Hunstad, I., Selvaggi, G., D'Agostino, N., England, P., Clarke, P., Pierozzi, M., 2003. Geodetic strain in peninsular Italy between 1875 and 2001. Geophys. Res. Lett. 30 (4), 1181. doi:10.1029/2002GL016447 2003. Jolivet, L., et al., 1998. Midcrustal shear zones in postorogenic extension: example from the northern Tyrrhenian Sea. J. Geophys. Res. 103 (B6), 12,123–12,160. Kostrov, V.V., 1974. Seismic moment and energy of earthquakes, and seismic flow of rock. Earth Phys. 1, 23–40.

Lavecchia, G., Brozzetti, R., Barchi, M.R., Menichetti, M., Keller, J.V., 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, 1107–1120. Meletti, C., Galadini, F., Valensise, G., Stucchi, M., Basili, R., Barba, R., Vannucci, G., Boschi, E., 2008. A seismic source zone model for the seismic hazard assessment of the Italian territory. Tectonophysics 450 (1–4), 85–108. Molnar, P., 1979. Earthquake recurrence intervals and plate tectonics. Bull. Seismol. Soc. Am. 69, 115–133. Nikolaidis, R., 2002, Observation of Geodetic and Seismic Deformation with the Global Positioning System, Ph.D. Thesis, University of California, San Diego. Pancha, A., Anderson, J.G., Kreemer, C., 2006, Comparison of seismic and geodetic scalar moment rates across the Basin and Range Province, BSSA, 96(1), 11–32. doi:10.1785/0120040166. Pondrelli, S., Salimbeni, S., Ekström, G., Morelli, A., Gasperini, P., Vannucci, G., 2006. The Italian CMT dataset from 1977 to the present. Phys. Earth Planet. Int. 286–303. doi10.1016/j.pepi.2006.07.008,159/3-4. Reutter, K.J., Gieseand, P., Closs, H., 1980. Lithospheric split in the descending plate: observations from the Northern Apennines. Tectonophysics 64, T1–T9. Savage, J.C., Simpson, R.W., 1997. Surface strain accumulation and the seismic moment tensor. Bull. Seismol. Soc. Am. 87, 1345–1353. Selvaggi, G., 1998. Spatial distribution of horizontal seismic strain in the Apennines from historical earthquakes. Ann. Geof. 41 (2), 241–251. Serpelloni, E., Anzidei, M., Baldi, P., Casula, G., Galvani, A., 2005. Crustal velocity and strain-rate fields in Italy and surrounding regions: new results from the analysis of permanent and non-permanent GPS networks. Geophys. J. Int. 161, 861–880. doi:10.1111/j.1365-246X.2006.06218.x. 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). Journal of Seismology 8, 297–312 2004. Ward, S.N., 1994. A multidisciplinary approach to seismic hazard in southern California. Bull. Seismol. Soc. Am. 84, 1293–1309. Wessel, P., Bercovici, D., 1998. Interpolation with splines in tension: a Green's function approach. J. Math. Geol. 30, 77–93. Williams, S.D.P., 2003. The effect of coloured noise on the uncertainties of rates estimated from geodetic time series. Journal of Geodesy 76 (9–10), 483–494. Williams, S.D.P., 2007. CATS: GPS Coordinate Time Series Analysis Software, GPS Solut. doi:10.1007/s10291-007-0086-4. Zumberge, J.F., Heflin, M.B., Jefferson, D.C., Watkins, M.M., Webb, F.H., 1997. Precise point positioning for the efficient and robust analysis of GPS data from large networks. J. Geophys. Res. 102, 5005–5501.