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Multi-seismic cycle velocity and strain fields for an active normal fault system, central Italy
Earth and Planetary Science Letters 251 (2006) 44 – 51 www.elsevier.com/locate/epsl
Multi-seismic cycle velocity and strain fields for an active normal fault system, central Italy Gerald P. Roberts ⁎ The Research School of Earth Sciences, Birkbeck and University College, University of London, Gower Street, London, WC1E 6BT, United Kingdom Received 26 October 2004; received in revised form 6 April 2005; accepted 28 November 2005 Available online 2 October 2006 Editor: B. Wood
Abstract Vertical offsets of Late Pleistocene–Holocene deposits and landforms are combined with slip-direction indicators on dipping active normal faults in central Italy to produce velocity and strain-rate fields averaged since 12–18 ka for a 5 by 5 km grid. Maximum strain-rates of 0.078 ± 0.0156 ppm/yr, or 0.052 ± 0.01 ppm/yr across a distance of 63,639 m imply maximum NE–SW extension rates of 5.0 ± 1.0 mm/yr or 3.3 ± 0.66 mm/yr depending on whether rates are averaged over 12 ka or 18 ka. The high spatial resolution reveals that strain-rates vary by up to 0.04 ± 0.008 or 0.06 ± 0.012 ppm/yr over a distance of only c. 7 km along the strike of the fault system, again depending on the timescale chosen. For the first time, these data allow a comparison of rates derived over time periods that would contain numerous seismic cycles with those measured with GPS that represent either interseismic strain, or displacements due to a single large earthquake with some interseismic strain. Overall, the rates compare reasonably well with those measured geodetically or derived from seismic moment summations (b 0.12 ppm/yr and generally in the range of 0.10–0.06 ppm/yr; 6.36–3.82 mm/yr). However, the present wide spacing of geodetic sites (30–40 km) means that the small-scale spatial variations in strain-rate since 12–18 ka have not yet been resolved using GPS. This makes it difficult to compare strain measured geodetically with that released during historical earthquakes to assess whether excess strain will be released in an impending earthquake. © 2006 Elsevier B.V. All rights reserved. Keywords: velocity field; Late Pleistocene–Holocene; active normal faults; Italy
1. Introduction Geodetic determinations of velocity and strain fields for areas undergoing active deformation are now relatively common, and have been used to comment on seismic hazards [1,2]. Specifically, it may be that if strain measured geodetically exceeds that implied by seismic moment summations for a given time period, ⁎ Tel.: +44 207 679 7713. E-mail address: [email protected]. 0012-821X/$ - see front matter © 2006 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2005.11.066
then the excess strain may be released during impending earthquakes. Such geodetic data are commonly collected over time periods of a few years using GPS, or exceptionally via re-occupation of old triangulation networks with GPS, allowing deformation to be measured over tens to hundreds of years. However, in areas of relatively slow intra-plate extensional deformation where rates of only a few millimetres per year are experienced, recurrence intervals for large magnitude earthquakes (N Ms 5.5) are hundreds to thousands of years. Geodetic measures of deformation thus reveal
G.P. Roberts / Earth and Planetary Science Letters 251 (2006) 44–51
either interseismic elastic strains (which here we take to include small magnitude earthquakes b c. Ms 5.5 that do not break the ground surface), or contain permanent displacements related to a single large earthquakes (Nc. Ms 5.5) plus some interseismic elastic strain. It would be interesting to know whether strain-rates over such time periods are representative of longer time periods, and whether strain has been sampled at a high enough spatial resolution to recognise discrepancies between geodetic and seismic strains. Unfortunately, such multi-seismic cycle deformation data with sufficient spatial coverage are extremely rare (e.g. [3,4]). This paper uses a deformation rate field derived from faulted Late Pleistocene–Holocene deposits and landforms to produce velocity and strain-rate fields for an active normal fault system in Lazio-Abruzzo, central Italy (Fig. 1). The deformation rates cover a time period from 12–18 ka to the present, a time period that contained numerous seismic cycles [5]. These rates are then compared with rates derived from a GPS re-occupation of an 1875 A.D. triangulation network, a 1994–1999 GPS campaign, and seismic moment summations using historical records and the CMT catalogue [2,6,7]. Results show that strain-rates from the different time periods are similar, but with higher spatial resolution gained through the use of offset geology. 2. Background Active extension in the Apennines occurs in previously shortened continental crust positioned within the zone of convergence between the Eurasian and African Plates [8–10]. Northward motion of the African plate through the Late Mesozoic–Recent has led to subduction of Tethyan ocean crust and collision of fragments of continental crust which now form the northern margins of the Mediterranean Sea, and the NW–SE striking fold– thrust belt in the Apennines. Thrusting continues to the present day on the Adriatic side of the Apennines, but in general, NE-directed thrusting in central Italy ceased in the Pliocene [11]. Extension in the Apennines began at about 2.5–3.0 Ma [12,13] with earlier extension to the west that produced thinned continental crust and a volcanic province on the western coast of Italy, and ocean crust in the Tyrrhenian Sea. Extension may be due to rollback of the Calabrian subduction zone [10,11]. Focal mechanisms and borehole breakout data indicate a general NE–SW extension in central and southern Italy [14], confirmed by studies of striated active normal faults at outcrop [12,15]. The record of large magnitude earthquakes in the region spans over 1000 years with many written reports of damaging earthquakes that were
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most-likely normal faulting events located in the Apennines [16,17]. The 1915 Fucino Earthquake (Ms 6.9–7.0) produced surface ruptures that exhibited maximum coseismic throws of about 1 m, and 33,000 fatalities [18]. GPS studies have documented an extension rate of 6 ± 2 mm/yr in a thin-corridor crossing the central Apennines [19] with GPS re-occupations of older triangulation networks across larger portions of the Apennines and seismic moment summations agreeing that extension rates are of the order of a few millimetres per year [2,20]. Similar deformation rates have been measured in the southern Apennines [2,7]. For the central Apennines, Roberts and Michetti [12] provide a database that allows extension rates averaged over 12–18 ka to be compared to those from geodesy and seismic moment summations (Fig. 1). They have measured the directions and amounts of slip that have occurred across active fault scarps in central Italy since the Late Pleistocene to Early Holocene (12–18 ka) for 71 sites. They measured the orientations of striations and corrugations on exposed active normal fault planes along the scarps to reveal the plunge and plunge directions of the motions. They measured the faultedoffsets of landforms and deposits that formed due to the demise of the last glacial maximum to derive the rates of fault motion. Ice cover at the last glacial maximum was restricted to small mountain–valley glaciers so the measured offsets are tectonic and are not produced by glacial unloading. Indeed, 36Cl cosmogenic exposure dating has revealed that the fault planes associated with these scarps have formed since about 12 ka, and have grown incrementally through 1.5–3.0 m slip events, probably associated with repeated Mw 6.7–7.0 earthquakes ([5] and references therein). Measured vertical offsets since the last glacial maximum are 4–20 m, indicating typical rates of vertical motion (throw-rates) of 0.2–1.7 mm/yr across individual faults [12]. These vertical motions are produced on fault planes that mostly dip towards the SW. However, the faulting is not pure dip-slip, with fault-slip directions varying from N–S to E–W due to the converging nature of normal fault hangingwall displacements (e.g. [12,15,21–23]). Additional rates and slip-directions are also available from the vicinity of palaeoseismological trench sites reported in the literature (e.g. [12,24–26]). With both the magnitude of vertical offset, and the plunge and plunge direction of the slip-vector known for each of the localities in the database of Roberts and Michetti [12] (Fig. 1), it is possible to use trigonometry to calculate the horizontal projection of the magnitude and orientation of the fault-slip vectors, which is done for the first time herein. These vectors can be applied to a grid of
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G.P. Roberts / Earth and Planetary Science Letters 251 (2006) 44–51
G.P. Roberts / Earth and Planetary Science Letters 251 (2006) 44–51
hypothetical geodetic stations to produce a velocity field for the deformation: strains and strain-rates can be calculated by comparing the relative velocities of the stations and then dividing by appropriate time intervals. 3. Method The database of Roberts and Michetti [12] was updated with new trench site data [27,28] and data from an exposure age dating site [5]. Trigonometry was then used to convert vertical offsets into horizontal vectors using both the plunge and plunge directions of the striations and corrugations on the fault planes (Fig. 2). The rates of motion were calculated by dividing the displacements by both 18 ka and 12 ka because the demise of the last glacial maximum may have been as early as 18 ka and the scarps may have begun to form as late as 12 ka [5,29]. This brackets the rates between their likely maximum and minimum values. The locations of these vectors were recorded in the field via UTM coordinates using a hand-held GPS in 1998, so the locations are within c. 100 m of their actual positions. This poor spatial resolution makes no difference to the calculations herein because the sites where offsets are recorded are larger than 100 m across; in other words, all points within the geographic area of uncertainty from the hand-held GPS measurements have experienced the same deformation. A 5 by 5 km grid of hypothetical geodetic stations was constructed in a UTM map projection to capture the spatial variations in fault-slip which occur over only a few kilometres along the strikes of faults (Fig. 1b). Vectors representing the motions across the faults were applied to the hypothetical geodetic stations in a stepwise manner from NE to SW, that is, the regional extension direction [12,14]. This produced a second grid showing new station locations. Comparison of the two grids produces velocity vectors showing the motion of each station. A restoration of the motions was chosen to display the results. In other words, the present positions of the stations were restored to their pre-deformation positions (i.e. their positions at 18 ka or 12 ka). This provides an 18 kyrs or 12 kyrs velocity field for sites in the hangingwalls of the faults located furthest to the NE in Fig. 2; the footwalls of the faults in the NE remain fixed in space. As it is generally agreed that no other active
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normal faults exist between these faults and the Adriatic coast [12,13,19], this can be thought of as a fixed Adriatic coast reference frame. However, note that stations move in a direction generally towards the Adriatic because a restoration has been chosen. To measure the strain, distances in a direction of 225° between equivalent stations located on two NW–SE oriented lines (A–B and C–D on Fig. 2) were calculated before and after the deformation. As these lines are located on opposite sides of, and generally parallel to the active fault system, the strain across the fault system can be derived by comparing the original and deformed distances between these lines. Strain-rates can then be derived by dividing by either 12 kyrs or 18 kyrs. These rates can then be compared with those from geodesy and seismology. Errors have been propagated through the calculations for directions, velocities and strain-rate. Errors for faultslip directions are shown as ± 5 times the average 99% confidence cone radius about the mean value (5°; N = 71 locations supported by 1706 measurements of the slipdirection), calculated using Fisher statistics (i.e. ± 25°) [12]; 5 times the value was used because up to 5 faults may be crossed on a NE–SW transect. Thus the error will be smaller than that shown for the majority of stations where less faults are crossed. Errors for velocities and strain-rate are shown for ± 20% measurement error of the vertical offsets of Late-Pleistocene and Holocene features, as described in the original paper containing the data [12]. 4. Results The orientations of lines representing the velocities of hypothetical geodetic stations reflect the converging patterns of slip across the individual faults and their summation between faults located across strike from one another (Fig. 2). In places this has produced herringbone patterns where velocity vectors from the NW tip of one fault (c. N–S slip) and the SE tip of another fault (c. E–W slip) combine (e.g. where the L'Aquila and Fucino Faults overlap). The lengths of the lines represent the magnitudes of slip across individual faults and their summation across neighbouring faults in a fixed Adriatic coast reference frame. The restorations imply an extension generally towards the SW in accordance with determinations of the active stress
Fig. 1. (a) Location map showing active normal faults in Italy and geodetic stations used to monitor them. (b) Fault map showing the plunge directions of striae/corrugations for specific localities along the faults. (c) Scarp height variation converted into throw-rate profiles assuming an 18 ka age for the scarps (see Discussion), augmented by throw-rates from palaeoseismological trench sites. Data from (b) and (c) are used to make Fig. 2. (d) Cumulative values for throw across the scarps summed along transects shown in Fig. 1b and c.
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Fig. 2. (a) Velocity field for Lazio-Abruzzo (located in b) overlain on an elevation map with a UTM grid, in an Adriatic Coast fixed reference frame showing the motions of points on a 5 by 5 km grid back in time from the present to 12–18 ka. Note this is a restoration back to 12–18 ka, so points move generally towards the Adriatic in a fixed Adriatic reference frame (e.g. see the white and black dots for an example). Data input are the throwrates from Roberts and Michetti [12] (Fig. 1), which have been converted into 2 D horizontal vectors using the plunge, plunge-direction and displacement along measured fault-slip directions. Errors for vectors are shown for ±20% error on the displacement and ±5 times the average 99% confidence cone radii on the slip direction from [12]. Errors are shown only for a strip parallel to the line A–B; errors further SW are the same whilst errors further NE are smaller. Line A–B is the transect along which the velocity field has been sampled (see Fig. 3). Strain is calculated between lines A–B and C–D (see Fig. 4). The total length of both A–B and C–D is only partially shown on this view. The velocity calculations were performed by adding the displacement vectors from the faults to the UTM coordinates of points on the grid in a stepwise fashion from NE to SW. The strain calculations (Fig. 4) were performed by finding the change in distance between equivalent points on the deformed and undeformed grids. (b) Location map showing known active normal faults in Italy [12]. Lines 1 and 2 show transects which cross the Apennines, but no known active normal faults, explaining why strain decreases to zero at X UTM coordinates 300,000 and 420,000 on Figs. 3 and 4.
field from borehole breakouts, focal mechanisms and striated faults (e.g. [14]). Fig. 3 shows a sample of the velocity field along line A–B located in Fig. 2. The displacements achieve a maximum horizontal magnitude of 60.2 ± 12.04 m in a direction of 039.3° ± 25° in a fixed Adriatic coast reference frame, implying a velocity of 3.35 ± 0.67 mm/yr or 5.02 ± 1.0 mm/yr depending on whether the displace-
ment is averaged over 18 kyrs or 12 kyrs respectively. These values decrease along strike to zero at the lateral terminations of the fault system noted by Roberts and Michetti [12]. In these locations, they drew serial crosssections across the geology and showed that it is possible to cross the Apennines in a NE–SW direction either without crossing a known active normal fault or by crossing very close to lateral fault tips (Fig. 1); hence
G.P. Roberts / Earth and Planetary Science Letters 251 (2006) 44–51
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Fig. 3. Sample of the 18–12 ka velocity field along a NW–SE transect located SW of the active faults (line A–B), in an Adriatic coast fixed reference frame.
the rates decrease to close to zero. Other active faults exist further NE and SW (Fig. 1), but these must form part of separate fault systems in the sense of Cowie and Roberts [30] as they lie outboard of the displacement minima shown in Fig. 1d. The strain in a direction of 225° reaches a maximum where there is 59.92 ± 11.86 m horizontal difference between the positions of stations on the original and deformed grids whose original separation was 63,639.61 m. This implies strain-rates of 0.052 ±
0.01 ppm/yr or 0.078 ± 0.0156 ppm/yr depending on whether the displacement is averaged over 18 ka or 12 ka respectively (Fig. 4). Again, the strain and strainrates decrease along strike to zero at the lateral terminations of the fault system noted by Roberts and Michetti [12]. By comparison (see Fig. 4), the upper bound to the strain-rate across the Apennines was given as 0.12 ppm/yr by Hunstad and England [6] and revised to between 0.06–0.10 ppm/yr by Hunstad et al. [2] who re-occupied an 1875 A.D. triangulation
Fig. 4. Strain over 12–18 ka in the direction 225° between lines A–B and C–D compared to published values from geodesy and seismology. For the 12–18 ka strains, 63,639.61 m is the undeformed distance between hypothetical geodetic stations, so 0.1 ppm/yr = 6.36 mm/yr across this distance.
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network with GPS. Anzidei et al. [7], studying an area further south, but still across the axis of the Apennines, reported strain-rates of 0.021 ppm/yr from a GPS campaign spanning 1994–1999, and strain-rates from seismic moment summations of 0.005–0.06 depending on whether the Harvard CMT catalogue or 100 years of historical seismicity are used. 5. Discussion Overall, there is a reasonable agreement between strain-rates measured over (1) a few years from GPS and seismology, (2) about 100 years from re-occupation of triangulation networks and historical seismicity, and (3) the Late Pleistocene and Holocene; all provide strainrates that are less than 0.12 ppm/yr, implying rates of extension of less than about 6 mm/yr over a distance of 63,639.61 m (the distance between lines A–B and C–D in a direction of 225°). However, the fact that strains over different time intervals are measured at different spatial resolutions (Fig. 1), makes it difficult to extract information pertaining to the seismic cycle on specific faults. Specifically, one would like to compare the strain measured geodetically with that released in historical earthquakes to see if any impending earthquakes are implied. However, the problem is that because the strain-rates derived from geodetic and seismological data are averages across areas of tens of kilometres, due to the large distances between geodetic stations (30–40 km in Hunstad et al. [2]), they do not resolve the small-scale spatial variations revealed by the 12–18 ka strain-rate field. These variations are significant as their magnitude is as great as 0.04 ± 0.08, or 0.06 ± 0.012 ppm/yr over a distance of only 7 km along the strike of the Apennines (calculated using the difference between sites located close to X UTM coordinate 380,000 on Fig. 4). The geodetic and seismological data have also not resolved the fact that active extension on large faults dies to zero at the NW and SE ends of the fault system described by Roberts and Michetti [12] (Fig. 1d). It is thus difficult to be sure that differences between strains determined geodetically and those determined through seismic moment summations can be viewed simply as missing seismic energy release, that is, impending earthquakes. This is because the 12–18 ka data suggest that one cannot be sure that one is comparing the geodetic strain across an individual fault with the seismic energy released on the same fault due to the poor spatial resolution of present geodetic networks. This problem was noted by Hunstad et al. [2], and its resolution must await geodetic strain measurements at higher spatial resolution.
6. Conclusions Velocity and strain fields can be constructed from measurements of striated faults offsetting Late Pleistocene and Holocene features. These allow a view of the multi-seismic cycle deformation for comparison with geodetic and seismological data that by necessity cover only periods of interseismic elastic strain or a period containing a large magnitude earthquake and some interseismic elastic strain. Overall the velocities and strain-rates compare well, but the geodetic data appear to lack sufficient spatial resolution to resolve small-scale spatial variations in deformation rates. This lack of spatial resolution makes it difficult to decide whether strain measured geodetically in excess of that implied by earthquake moment release suggests an impending earthquake. There is a need for more studies of multiseismic cycle deformation fields to establish whether the patterns described herein are the rule or the exception. Acknowledgements The author thanks Ioannis Papanikolaou, Alessandro Michetti, Patience Cowie, Greg Tucker, Alex Whittaker, Nigel Morewood and Kerry Sieh for discussions that have led to this study. The work was funded by NERC GR9/02995, NERC NE/B504165/1, and Birkbeck, University of London. The Benfield Greig Hazard Research Centre at UCL is thanked for support. Iain Stewart, Andy Nicol and Bernie Wood are thanked for their referees' and Editor's comments. References [1] P. Clarke, R. Davies, P. England, B. Parsons, H. Billiris, D. Paradissis, G. Veis, P. Denys, P. Cross, V. Ashkenazi, R. Bingley, Geodetic estimate of seismic hazard in the Gulf of Korinthos, Geophys. Res. Lett. 24 (1997) 1303–1306. [2] I. Hunstad, G. Selvaggi, N. D'Agostino, P. England, P. Clarke, M. Pierozzi, Geodetic strain in peninsular Italy between 1875 and 2001, Geophys. Res. Lett. 30 (4) (2003) 1181, doi:10.1029/2002GL016447. [3] P. England, P. Molnar, The field of crustal velocity in Asia calculated from Quaternary rates of slip on faults, Geophys. J. Int. 130 (1997) 551–582. [4] S. Lamb, Vertical axis rotations in the Bolivian orocline, South America 2. Kinematic and dynamical implications, J. Geophys. Res. 106 (2001) 26,633–26,653. [5] L. Palumbo, L. Benedetti, D. Bourles, A. Cinque, R. Finkel, Slip history of the Magnola fault (Apennines, Central Italy) from 36Cl surface exposure dating: evidence for strong earthquakes over the Holocene, Earth Planet. Sci. Lett. 225 (2004) 163–176. [6] I. Hunstad, P. England, An upper bound on the rate of strain in the central Apennines, Italy, from triangulation measurements between 1869 and 1963, Earth Planet. Sci. Lett. 169 (1999) 261–267.
G.P. Roberts / Earth and Planetary Science Letters 251 (2006) 44–51 [7] M. Anzidei, P. Baldi, G. Casula, A. Galvani, E. Mantovani, A. Pesci, F. Riguzzi, E. Serpelloni, Insights into present-day crustal motion in the central Mediterranean area from GPS surveys, Geophys. J. Int. 146 (2001) 98–110. [8] C. Doglioni, Some remarks on the origin of foredeeps, Tectonophysics 228 (1993) 1–20. [9] H. Anderson, J. Jackson, Active tectonics of the Adriatic region, Geophys. J. R. Astron. Soc. 91 (1987) 937–983. [10] L. Jolivet, C. Facenna, B. Goffe, M. Mattei, F. Rossetti, C. Brunet, F. Storti, R. Funiciello, J.-P. Cadet, N. D'Agostino, T. Parra, Midcrustal shear zones in postorogenic extension: example from the northern Tyrrhenian Sea, J. Geophys. Res. 103 (1998) 12,123–12,160. [11] E. Patacca, R. Sartori, P. Scandone, Tyrrhenian Basin and Apenninic Arcs: kinematic relations since late Tortonian times, Mem. Soc. Geol. Ital. 45 (1990) 425–451. [12] G.P. Roberts, A.M. Michetti, Spatial and temporal variations in growth rates along active normal fault Systems: an example from Lazio-Abruzzo, central Italy, J. Struct. Geol. 26 (2004) 339–376. [13] G.P. Cavinato, P.G. De Celles, Extensional basins in the tectonically bimodal central Apennines fold–thrust belt, Italy: response to corner flow above a subducting slab in retrograde motion, Geology 27 (1999) 955–958. [14] P. Montone, A. Amato, S. Pondrelli, Active stress map of Italy, J. Geophys. Res. 104 (1999) 25,595–25,610. [15] A.M. Michetti, L. Ferreli, E. Esposito, S. Porfido, A.M. Blumetti, E. Vittori, L. Serva, G.P. Roberts, Ground effects during the September 9, 1998 ML = 5.5 Lauria Earthquake in southern Italy and the seismic potential of the “aseismic” Pollino region: preliminary report, Seismol. Res. Lett. 71 (2000) 31–46. [16] D. Postpischl (Ed.), Atlas of Isoseismal Maps of the Italian Earthquakes, C.N.R., Quaderni de “La Ricerca Scientifica”, 114 (2A), 1985, 164pp. [17] E. Boschi, G. Ferrari, P. Gasperini, E. Guidoboni, G. Smriglio, G. Valensise, Catalogo dei forti terremoti in Italia dal 461 a.C. al 1980, Istituto Nazionale di Geofisica — SGA Storia Geofisica Ambiente, Grafica Ragno, Ozzano Emilia, BO, Italy, 1995, 973 pp. [18] C. Margottini, A. Screpanti, Temporal evolution of the seismic crisis related to the 13th January 1915, Avezzano earthquake, in: C. Margottini, L. Serva (Eds.), Historical Seismicity of Central-
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Multi-seismic cycle velocity and strain fields for an active normal fault system, central Italy Gerald P. Roberts ⁎ The Research School of Earth Sciences, Birkbeck and University College, University of London, Gower Street, London, WC1E 6BT, United Kingdom Received 26 October 2004; received in revised form 6 April 2005; accepted 28 November 2005 Available online 2 October 2006 Editor: B. Wood
Abstract Vertical offsets of Late Pleistocene–Holocene deposits and landforms are combined with slip-direction indicators on dipping active normal faults in central Italy to produce velocity and strain-rate fields averaged since 12–18 ka for a 5 by 5 km grid. Maximum strain-rates of 0.078 ± 0.0156 ppm/yr, or 0.052 ± 0.01 ppm/yr across a distance of 63,639 m imply maximum NE–SW extension rates of 5.0 ± 1.0 mm/yr or 3.3 ± 0.66 mm/yr depending on whether rates are averaged over 12 ka or 18 ka. The high spatial resolution reveals that strain-rates vary by up to 0.04 ± 0.008 or 0.06 ± 0.012 ppm/yr over a distance of only c. 7 km along the strike of the fault system, again depending on the timescale chosen. For the first time, these data allow a comparison of rates derived over time periods that would contain numerous seismic cycles with those measured with GPS that represent either interseismic strain, or displacements due to a single large earthquake with some interseismic strain. Overall, the rates compare reasonably well with those measured geodetically or derived from seismic moment summations (b 0.12 ppm/yr and generally in the range of 0.10–0.06 ppm/yr; 6.36–3.82 mm/yr). However, the present wide spacing of geodetic sites (30–40 km) means that the small-scale spatial variations in strain-rate since 12–18 ka have not yet been resolved using GPS. This makes it difficult to compare strain measured geodetically with that released during historical earthquakes to assess whether excess strain will be released in an impending earthquake. © 2006 Elsevier B.V. All rights reserved. Keywords: velocity field; Late Pleistocene–Holocene; active normal faults; Italy
1. Introduction Geodetic determinations of velocity and strain fields for areas undergoing active deformation are now relatively common, and have been used to comment on seismic hazards [1,2]. Specifically, it may be that if strain measured geodetically exceeds that implied by seismic moment summations for a given time period, ⁎ Tel.: +44 207 679 7713. E-mail address: [email protected]. 0012-821X/$ - see front matter © 2006 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2005.11.066
then the excess strain may be released during impending earthquakes. Such geodetic data are commonly collected over time periods of a few years using GPS, or exceptionally via re-occupation of old triangulation networks with GPS, allowing deformation to be measured over tens to hundreds of years. However, in areas of relatively slow intra-plate extensional deformation where rates of only a few millimetres per year are experienced, recurrence intervals for large magnitude earthquakes (N Ms 5.5) are hundreds to thousands of years. Geodetic measures of deformation thus reveal
G.P. Roberts / Earth and Planetary Science Letters 251 (2006) 44–51
either interseismic elastic strains (which here we take to include small magnitude earthquakes b c. Ms 5.5 that do not break the ground surface), or contain permanent displacements related to a single large earthquakes (Nc. Ms 5.5) plus some interseismic elastic strain. It would be interesting to know whether strain-rates over such time periods are representative of longer time periods, and whether strain has been sampled at a high enough spatial resolution to recognise discrepancies between geodetic and seismic strains. Unfortunately, such multi-seismic cycle deformation data with sufficient spatial coverage are extremely rare (e.g. [3,4]). This paper uses a deformation rate field derived from faulted Late Pleistocene–Holocene deposits and landforms to produce velocity and strain-rate fields for an active normal fault system in Lazio-Abruzzo, central Italy (Fig. 1). The deformation rates cover a time period from 12–18 ka to the present, a time period that contained numerous seismic cycles [5]. These rates are then compared with rates derived from a GPS re-occupation of an 1875 A.D. triangulation network, a 1994–1999 GPS campaign, and seismic moment summations using historical records and the CMT catalogue [2,6,7]. Results show that strain-rates from the different time periods are similar, but with higher spatial resolution gained through the use of offset geology. 2. Background Active extension in the Apennines occurs in previously shortened continental crust positioned within the zone of convergence between the Eurasian and African Plates [8–10]. Northward motion of the African plate through the Late Mesozoic–Recent has led to subduction of Tethyan ocean crust and collision of fragments of continental crust which now form the northern margins of the Mediterranean Sea, and the NW–SE striking fold– thrust belt in the Apennines. Thrusting continues to the present day on the Adriatic side of the Apennines, but in general, NE-directed thrusting in central Italy ceased in the Pliocene [11]. Extension in the Apennines began at about 2.5–3.0 Ma [12,13] with earlier extension to the west that produced thinned continental crust and a volcanic province on the western coast of Italy, and ocean crust in the Tyrrhenian Sea. Extension may be due to rollback of the Calabrian subduction zone [10,11]. Focal mechanisms and borehole breakout data indicate a general NE–SW extension in central and southern Italy [14], confirmed by studies of striated active normal faults at outcrop [12,15]. The record of large magnitude earthquakes in the region spans over 1000 years with many written reports of damaging earthquakes that were
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most-likely normal faulting events located in the Apennines [16,17]. The 1915 Fucino Earthquake (Ms 6.9–7.0) produced surface ruptures that exhibited maximum coseismic throws of about 1 m, and 33,000 fatalities [18]. GPS studies have documented an extension rate of 6 ± 2 mm/yr in a thin-corridor crossing the central Apennines [19] with GPS re-occupations of older triangulation networks across larger portions of the Apennines and seismic moment summations agreeing that extension rates are of the order of a few millimetres per year [2,20]. Similar deformation rates have been measured in the southern Apennines [2,7]. For the central Apennines, Roberts and Michetti [12] provide a database that allows extension rates averaged over 12–18 ka to be compared to those from geodesy and seismic moment summations (Fig. 1). They have measured the directions and amounts of slip that have occurred across active fault scarps in central Italy since the Late Pleistocene to Early Holocene (12–18 ka) for 71 sites. They measured the orientations of striations and corrugations on exposed active normal fault planes along the scarps to reveal the plunge and plunge directions of the motions. They measured the faultedoffsets of landforms and deposits that formed due to the demise of the last glacial maximum to derive the rates of fault motion. Ice cover at the last glacial maximum was restricted to small mountain–valley glaciers so the measured offsets are tectonic and are not produced by glacial unloading. Indeed, 36Cl cosmogenic exposure dating has revealed that the fault planes associated with these scarps have formed since about 12 ka, and have grown incrementally through 1.5–3.0 m slip events, probably associated with repeated Mw 6.7–7.0 earthquakes ([5] and references therein). Measured vertical offsets since the last glacial maximum are 4–20 m, indicating typical rates of vertical motion (throw-rates) of 0.2–1.7 mm/yr across individual faults [12]. These vertical motions are produced on fault planes that mostly dip towards the SW. However, the faulting is not pure dip-slip, with fault-slip directions varying from N–S to E–W due to the converging nature of normal fault hangingwall displacements (e.g. [12,15,21–23]). Additional rates and slip-directions are also available from the vicinity of palaeoseismological trench sites reported in the literature (e.g. [12,24–26]). With both the magnitude of vertical offset, and the plunge and plunge direction of the slip-vector known for each of the localities in the database of Roberts and Michetti [12] (Fig. 1), it is possible to use trigonometry to calculate the horizontal projection of the magnitude and orientation of the fault-slip vectors, which is done for the first time herein. These vectors can be applied to a grid of
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G.P. Roberts / Earth and Planetary Science Letters 251 (2006) 44–51
hypothetical geodetic stations to produce a velocity field for the deformation: strains and strain-rates can be calculated by comparing the relative velocities of the stations and then dividing by appropriate time intervals. 3. Method The database of Roberts and Michetti [12] was updated with new trench site data [27,28] and data from an exposure age dating site [5]. Trigonometry was then used to convert vertical offsets into horizontal vectors using both the plunge and plunge directions of the striations and corrugations on the fault planes (Fig. 2). The rates of motion were calculated by dividing the displacements by both 18 ka and 12 ka because the demise of the last glacial maximum may have been as early as 18 ka and the scarps may have begun to form as late as 12 ka [5,29]. This brackets the rates between their likely maximum and minimum values. The locations of these vectors were recorded in the field via UTM coordinates using a hand-held GPS in 1998, so the locations are within c. 100 m of their actual positions. This poor spatial resolution makes no difference to the calculations herein because the sites where offsets are recorded are larger than 100 m across; in other words, all points within the geographic area of uncertainty from the hand-held GPS measurements have experienced the same deformation. A 5 by 5 km grid of hypothetical geodetic stations was constructed in a UTM map projection to capture the spatial variations in fault-slip which occur over only a few kilometres along the strikes of faults (Fig. 1b). Vectors representing the motions across the faults were applied to the hypothetical geodetic stations in a stepwise manner from NE to SW, that is, the regional extension direction [12,14]. This produced a second grid showing new station locations. Comparison of the two grids produces velocity vectors showing the motion of each station. A restoration of the motions was chosen to display the results. In other words, the present positions of the stations were restored to their pre-deformation positions (i.e. their positions at 18 ka or 12 ka). This provides an 18 kyrs or 12 kyrs velocity field for sites in the hangingwalls of the faults located furthest to the NE in Fig. 2; the footwalls of the faults in the NE remain fixed in space. As it is generally agreed that no other active
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normal faults exist between these faults and the Adriatic coast [12,13,19], this can be thought of as a fixed Adriatic coast reference frame. However, note that stations move in a direction generally towards the Adriatic because a restoration has been chosen. To measure the strain, distances in a direction of 225° between equivalent stations located on two NW–SE oriented lines (A–B and C–D on Fig. 2) were calculated before and after the deformation. As these lines are located on opposite sides of, and generally parallel to the active fault system, the strain across the fault system can be derived by comparing the original and deformed distances between these lines. Strain-rates can then be derived by dividing by either 12 kyrs or 18 kyrs. These rates can then be compared with those from geodesy and seismology. Errors have been propagated through the calculations for directions, velocities and strain-rate. Errors for faultslip directions are shown as ± 5 times the average 99% confidence cone radius about the mean value (5°; N = 71 locations supported by 1706 measurements of the slipdirection), calculated using Fisher statistics (i.e. ± 25°) [12]; 5 times the value was used because up to 5 faults may be crossed on a NE–SW transect. Thus the error will be smaller than that shown for the majority of stations where less faults are crossed. Errors for velocities and strain-rate are shown for ± 20% measurement error of the vertical offsets of Late-Pleistocene and Holocene features, as described in the original paper containing the data [12]. 4. Results The orientations of lines representing the velocities of hypothetical geodetic stations reflect the converging patterns of slip across the individual faults and their summation between faults located across strike from one another (Fig. 2). In places this has produced herringbone patterns where velocity vectors from the NW tip of one fault (c. N–S slip) and the SE tip of another fault (c. E–W slip) combine (e.g. where the L'Aquila and Fucino Faults overlap). The lengths of the lines represent the magnitudes of slip across individual faults and their summation across neighbouring faults in a fixed Adriatic coast reference frame. The restorations imply an extension generally towards the SW in accordance with determinations of the active stress
Fig. 1. (a) Location map showing active normal faults in Italy and geodetic stations used to monitor them. (b) Fault map showing the plunge directions of striae/corrugations for specific localities along the faults. (c) Scarp height variation converted into throw-rate profiles assuming an 18 ka age for the scarps (see Discussion), augmented by throw-rates from palaeoseismological trench sites. Data from (b) and (c) are used to make Fig. 2. (d) Cumulative values for throw across the scarps summed along transects shown in Fig. 1b and c.
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Fig. 2. (a) Velocity field for Lazio-Abruzzo (located in b) overlain on an elevation map with a UTM grid, in an Adriatic Coast fixed reference frame showing the motions of points on a 5 by 5 km grid back in time from the present to 12–18 ka. Note this is a restoration back to 12–18 ka, so points move generally towards the Adriatic in a fixed Adriatic reference frame (e.g. see the white and black dots for an example). Data input are the throwrates from Roberts and Michetti [12] (Fig. 1), which have been converted into 2 D horizontal vectors using the plunge, plunge-direction and displacement along measured fault-slip directions. Errors for vectors are shown for ±20% error on the displacement and ±5 times the average 99% confidence cone radii on the slip direction from [12]. Errors are shown only for a strip parallel to the line A–B; errors further SW are the same whilst errors further NE are smaller. Line A–B is the transect along which the velocity field has been sampled (see Fig. 3). Strain is calculated between lines A–B and C–D (see Fig. 4). The total length of both A–B and C–D is only partially shown on this view. The velocity calculations were performed by adding the displacement vectors from the faults to the UTM coordinates of points on the grid in a stepwise fashion from NE to SW. The strain calculations (Fig. 4) were performed by finding the change in distance between equivalent points on the deformed and undeformed grids. (b) Location map showing known active normal faults in Italy [12]. Lines 1 and 2 show transects which cross the Apennines, but no known active normal faults, explaining why strain decreases to zero at X UTM coordinates 300,000 and 420,000 on Figs. 3 and 4.
field from borehole breakouts, focal mechanisms and striated faults (e.g. [14]). Fig. 3 shows a sample of the velocity field along line A–B located in Fig. 2. The displacements achieve a maximum horizontal magnitude of 60.2 ± 12.04 m in a direction of 039.3° ± 25° in a fixed Adriatic coast reference frame, implying a velocity of 3.35 ± 0.67 mm/yr or 5.02 ± 1.0 mm/yr depending on whether the displace-
ment is averaged over 18 kyrs or 12 kyrs respectively. These values decrease along strike to zero at the lateral terminations of the fault system noted by Roberts and Michetti [12]. In these locations, they drew serial crosssections across the geology and showed that it is possible to cross the Apennines in a NE–SW direction either without crossing a known active normal fault or by crossing very close to lateral fault tips (Fig. 1); hence
G.P. Roberts / Earth and Planetary Science Letters 251 (2006) 44–51
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Fig. 3. Sample of the 18–12 ka velocity field along a NW–SE transect located SW of the active faults (line A–B), in an Adriatic coast fixed reference frame.
the rates decrease to close to zero. Other active faults exist further NE and SW (Fig. 1), but these must form part of separate fault systems in the sense of Cowie and Roberts [30] as they lie outboard of the displacement minima shown in Fig. 1d. The strain in a direction of 225° reaches a maximum where there is 59.92 ± 11.86 m horizontal difference between the positions of stations on the original and deformed grids whose original separation was 63,639.61 m. This implies strain-rates of 0.052 ±
0.01 ppm/yr or 0.078 ± 0.0156 ppm/yr depending on whether the displacement is averaged over 18 ka or 12 ka respectively (Fig. 4). Again, the strain and strainrates decrease along strike to zero at the lateral terminations of the fault system noted by Roberts and Michetti [12]. By comparison (see Fig. 4), the upper bound to the strain-rate across the Apennines was given as 0.12 ppm/yr by Hunstad and England [6] and revised to between 0.06–0.10 ppm/yr by Hunstad et al. [2] who re-occupied an 1875 A.D. triangulation
Fig. 4. Strain over 12–18 ka in the direction 225° between lines A–B and C–D compared to published values from geodesy and seismology. For the 12–18 ka strains, 63,639.61 m is the undeformed distance between hypothetical geodetic stations, so 0.1 ppm/yr = 6.36 mm/yr across this distance.
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network with GPS. Anzidei et al. [7], studying an area further south, but still across the axis of the Apennines, reported strain-rates of 0.021 ppm/yr from a GPS campaign spanning 1994–1999, and strain-rates from seismic moment summations of 0.005–0.06 depending on whether the Harvard CMT catalogue or 100 years of historical seismicity are used. 5. Discussion Overall, there is a reasonable agreement between strain-rates measured over (1) a few years from GPS and seismology, (2) about 100 years from re-occupation of triangulation networks and historical seismicity, and (3) the Late Pleistocene and Holocene; all provide strainrates that are less than 0.12 ppm/yr, implying rates of extension of less than about 6 mm/yr over a distance of 63,639.61 m (the distance between lines A–B and C–D in a direction of 225°). However, the fact that strains over different time intervals are measured at different spatial resolutions (Fig. 1), makes it difficult to extract information pertaining to the seismic cycle on specific faults. Specifically, one would like to compare the strain measured geodetically with that released in historical earthquakes to see if any impending earthquakes are implied. However, the problem is that because the strain-rates derived from geodetic and seismological data are averages across areas of tens of kilometres, due to the large distances between geodetic stations (30–40 km in Hunstad et al. [2]), they do not resolve the small-scale spatial variations revealed by the 12–18 ka strain-rate field. These variations are significant as their magnitude is as great as 0.04 ± 0.08, or 0.06 ± 0.012 ppm/yr over a distance of only 7 km along the strike of the Apennines (calculated using the difference between sites located close to X UTM coordinate 380,000 on Fig. 4). The geodetic and seismological data have also not resolved the fact that active extension on large faults dies to zero at the NW and SE ends of the fault system described by Roberts and Michetti [12] (Fig. 1d). It is thus difficult to be sure that differences between strains determined geodetically and those determined through seismic moment summations can be viewed simply as missing seismic energy release, that is, impending earthquakes. This is because the 12–18 ka data suggest that one cannot be sure that one is comparing the geodetic strain across an individual fault with the seismic energy released on the same fault due to the poor spatial resolution of present geodetic networks. This problem was noted by Hunstad et al. [2], and its resolution must await geodetic strain measurements at higher spatial resolution.
6. Conclusions Velocity and strain fields can be constructed from measurements of striated faults offsetting Late Pleistocene and Holocene features. These allow a view of the multi-seismic cycle deformation for comparison with geodetic and seismological data that by necessity cover only periods of interseismic elastic strain or a period containing a large magnitude earthquake and some interseismic elastic strain. Overall the velocities and strain-rates compare well, but the geodetic data appear to lack sufficient spatial resolution to resolve small-scale spatial variations in deformation rates. This lack of spatial resolution makes it difficult to decide whether strain measured geodetically in excess of that implied by earthquake moment release suggests an impending earthquake. There is a need for more studies of multiseismic cycle deformation fields to establish whether the patterns described herein are the rule or the exception. Acknowledgements The author thanks Ioannis Papanikolaou, Alessandro Michetti, Patience Cowie, Greg Tucker, Alex Whittaker, Nigel Morewood and Kerry Sieh for discussions that have led to this study. The work was funded by NERC GR9/02995, NERC NE/B504165/1, and Birkbeck, University of London. The Benfield Greig Hazard Research Centre at UCL is thanked for support. Iain Stewart, Andy Nicol and Bernie Wood are thanked for their referees' and Editor's comments. References [1] P. Clarke, R. Davies, P. England, B. Parsons, H. Billiris, D. Paradissis, G. Veis, P. Denys, P. Cross, V. Ashkenazi, R. Bingley, Geodetic estimate of seismic hazard in the Gulf of Korinthos, Geophys. Res. Lett. 24 (1997) 1303–1306. [2] I. Hunstad, G. Selvaggi, N. D'Agostino, P. England, P. Clarke, M. Pierozzi, Geodetic strain in peninsular Italy between 1875 and 2001, Geophys. Res. Lett. 30 (4) (2003) 1181, doi:10.1029/2002GL016447. [3] P. England, P. Molnar, The field of crustal velocity in Asia calculated from Quaternary rates of slip on faults, Geophys. J. Int. 130 (1997) 551–582. [4] S. Lamb, Vertical axis rotations in the Bolivian orocline, South America 2. Kinematic and dynamical implications, J. Geophys. Res. 106 (2001) 26,633–26,653. [5] L. Palumbo, L. Benedetti, D. Bourles, A. Cinque, R. Finkel, Slip history of the Magnola fault (Apennines, Central Italy) from 36Cl surface exposure dating: evidence for strong earthquakes over the Holocene, Earth Planet. Sci. Lett. 225 (2004) 163–176. [6] I. Hunstad, P. England, An upper bound on the rate of strain in the central Apennines, Italy, from triangulation measurements between 1869 and 1963, Earth Planet. Sci. Lett. 169 (1999) 261–267.
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