Seismic activity along fault branches of the Dead Sea-Jordan Transform System: The Carmel-Tirtza fault system

TECTONOPHYSICS ELSEVIER

Tectonophysics267(1996)317-330

Seismic activity along fault branches of the Dead Sea-Jordan Transform System: The Carmel-Tirtza fault system A. Hofstetter a, T. van Eck b, A. Shapira a a Institute for Petroleum Research and Geophysics. P.O. Box 2286. Holon 58122, Israel b Institute of Earth Sciences, Unicersio'of Utrecht, P.O. Box 80021, 3508 TA Utrecht, The Netherlands

Received 23 August 1995; accepted 28 May 1996

Abstract About 550 earthquakes (1.0 _
1. Introduction The Dead Sea-Jordan rift which separates the Arabian plate from the African plate (Freund et al., 1970) started to form most probably during post-Early Miocene (Bartov et al., 1980). During the midCenozoic the Afro-Arabian continent was rifted and broken into several separate sub-plates. Freund et al. (1970) related the activity of the Dead Sea transform to the young normal faulting east of the Galilee and Lebanon. They noted the local intense faulting east

of the Galilee and Lebanon that is produced by motion along curved segments of the Dead Sea rift, and their absence in the southern parts of the rift. Recent studies, based on various geophysical evidence, suggest that the continental margin is divided into two major provinces, the boundary between which is the Mt. Carmel-Tirtza fault system and its postulated continuation offshore (Garfunkel and AImagor, 1985; Ben-Avraham and Ginzburg, 1990; Hofstetter et al., 1991). The southern province includes central Israel to Mt. Carmel in the north, and

0040- t 951/96/$15.00 Copyright © 1996 Elsevier Science B.V. All rights reserved. PII S0040-1951(96)00108-4

A. H@tetter et al. / Tectonophysics 267 (1996) 317-330

318

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Fig. 1. Earthquakes that are used in this study along the Mt. Carmet-Tirtza fault system (solid lines represent faults) observed during 1984 to 1994, with 13 composite focal mechanism solutions (Table ITable 2) and the seismic stations (solid triangle). Event cluster numbers are referred to in the text.

A. Hofstetter et al. / Tectonophysics 267 (1996) 317-330

319

the northern province includes northern Israel and Lebanon. The C a r m e l - T i r t z a fault zone, composed of numerous N W - S E - t r e n d i n g small faults, seemingly branching out from the Jordan Valley, has a prominent tectonic and morphological expression. The fault area is known to be of complex geological and tectonic structure (Achmon, 1986; Ron et al., 1991), with an inland structural uplift of more than 800 m above sea level. The combination of the Tirtza fault with its continuations along the Mt. Carmel fault and the Gilboa fault form a seismically active zone more than 130 km long. If the activity in off-coast Galilee is indeed related to the continuation of the Carmel fault in the Mediterranean Sea (Ron, 1984; Achmon, 1986; Ben-Gai, 1989), then the total length of this fault zone is at least over 170 km. Earthquakes in this region occur sometimes in swarms or clusters, with the largest widely felt event taking place near the Mt. Carmel fault ( M L = 5.3 on 2 4 / 8 / 1 9 8 4 ) . That event has arisen considerable interest in the region due to the previous lack of knowledge of the seismicity level. The effects of several historical earthquakes, probably occurred at the junction of Mt. C a r m e l - T i r t z a fault and the Dead Sea transform, are still debatable, i.e., the event of 1546 with the allegedly most destructive and severe damage (BenMenahem, 1979; Rotstein, 1987). In contrast, some moderate damage may have taken place (Ambraseys and Karcz, 1992).

given in IPRG Bulletins ( 1 9 8 4 - 1 9 9 4 ) and Van Eck and Hofstetter (1989). The second data set includes analog recordings of data for the period 1900 to 1984. The seismicity rate is obtained using both data sets during 1900 to 1994.

2. Data acquisition

3

Seismicity in Israel is monitored by the Israel Seismograph Network (ISN; IPRG Seismological Bulletins, 1984-1994), which includes a northern portable array comprising up to 13 seismic stations, and 2 three-component stations (Fig. 1) that were deployed over various periods o f time. During an eleven-year period, 1984 to 1994, about 550 microearthquakes in the region were digitally recorded by the seismic network. The determinable magnitudes range between 1.0 and 5.3, on the local magnitude scale, M L (Shapira, 1988). Using this data set we are able to obtain accurate epicenter location, source mechanism and spectral characteristics. A detailed description of a typical seismic station is

3. Distribution of seismicity in space and time 3.1. Epicentral location and clustering The crustal velocity model of Israel, determined from a refraction study using known quarry explosions, is relatively simple, and the location discrepancy [using a location algorithm based on search procedure (IPRG Seismological Bulletins, 1984-

Table 1 List of the clusters and single earthquakes that are used for the fault plane solution. Event cluster numbers are referred to in the text. Depth estimates refer to single event locations with errors of approximately + 2 km (see text for location errors) Cluster Date

ML Latitude Longitude Depth

yr mo dy hr mn 1 2

4 5 6 7 8 9 10 11 12 13

84 84 84 84 84 84 84 84 84 84 84 84 84 84 84 84 88 88 88 89 89 90 91 91 92 93 94

5 8 8 8 8 8 8 8 8 11 11 11 11 11 11 11 1 3 9 1 1 12 3 12 7 3 9

30 24 24 24 24 25 25 27 31 4 5 5 5 5 5 5 30 2 5 3 6 7 1 29 29 26 16

10 6 6 8 15 0 14 7 22 19 1 1 1 4 5 6 3 14 14 17 10 20 7 1 5 6 3

16 2 21 17 16 48 35 12 13 24 15 44 50 51 14 10 0 45 33 10 59 24 51 38 30 57 18

3.2 5.3 1.4 1.0 1.7 2.1 1.1 2.8 2.5 3.6 4.0 3.8 1.8 3.6 3.5 2.1 3.8 3.1 3.5 3.9 3.7 3.1 3.3 3.2 3.4 3.1 4.1

N

E

(km)

32.29 32.66 32.72 32.70 32.68 32.69 32.67 32.65 32.67 32.12 32.12 32.12 32.11 32.11 32.13 32.11 32.20 33.11 33.05 32.48 32.46 32.46 33.74 32.09 32.35 32.16 32.05

35.41 35.18 35.23 35.19 35.18 35.20 35.15 35.24 35.19 35.39 35.38 35.39 35.40 35.39 35.37 35.37 35.51 34.74 34.58 35.46 35.48 35.28 35.19 35.49 35.47 35.27 35.50

12 21 20 21 12 12 15 12 12 2 2 2 4 4 5 4 12 24 25 12 21 12 12 12 11 12 19

320

A. H of~tette r et al. / Tectonophvsics 267 (1996) 3 1 7 - 3 3 0

1994] between the reported and the calculated explosions is small (__1 km). Taking into consideration that not all station-event configurations are optimal, we may expect an epicentral location accuracy of + 2 km on the average. The location of the events used in this study are shown in Fig. 1, along with the distribution of recording stations, which provides a good azimuthal coverage for most of the events. The

seismic activity is mainly along the northwestsoutheast surface faults. This activity is not spatially uniform, in the sense that there is a relatively high concentration of events close to the junction of the Dead Sea Transform and the Carmel-Tirtza fault decreasing somewhat gradually in the northwest direction towards the Mediterranean Sea (also Salamon, 1993). Along the fault there are areas of little,

N

N

N

N

N

N

N

N

N

N

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6

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N

N

N

dilatation o

compression

Fig. 2. Composite fault plane solutions of 13 clusters along the Mt. Carmel Tirtza fault system, obtained using the algorithm of Reasenberg and Oppenheimer (1985). For cluster location see Fig. I and Table I.

A. Hofstetter et al. / Tectonophysics 267 (1996) 317 330

diffused or no activity, i.e., southeast of clusters 2 and 8 and along the Mediterranean segment in Fig. 1. It is obvious that the observation period is short and extrapolation of this inactivity during 11 years to longer periods of time can be doubtful. The seismic activity in the East Mediterranean Sea is off the main fault and cannot be correlated with the surface fault (Garfunkel and Almagor, 1985; Ben-Gai, 1989). The seismicity in the region northwest of the Gilboa fault (Fig. 1) is largely off the main surface geological fault. This fact is in good agreement with recent reflection study (Rotstein et al., 1993), suggesting that this region, up to the Mediterranean Sea, is characterized by a wide zone of deformation rather than a single fault line. Similar observations in their character of irregular seismic activity or cluster distribution along the San Andreas fault system, or diffused activity in the vicinity of the junction of the 102

branching faults, was shown by Hill et al. (1990). They found that along the San Andreas fault or its associated faults the seismicity follows a series of sub-parallel lineations. We define a cluster as a group of earthquakes that share almost the same focal origin and time as well and whose signals appear to be highly similar at the various stations. Many of the located earthquakes seem to be clustered. Part of this activity is associated with aftershocks such as that of the main earthquake of August 24, 1984, M L, or the November, 1984 felt swarm (clusters 2 and 3 in Fig. 1). Along the Tirtza fault, especially close to the junction with the Dead Sea rift, the activity is rather divided and concentrated in clusters, as it follows several mapped active faults in this region (Salamon, 1993). Towards the Mediterranean Sea, large portion of the activity is off the main obvious geological fault. i

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Fig. 3. The variation o f the mean annual frequency with respect to the date interval d T = 1 9 9 4 - t , for different groups of data, i.e. M L = 1.8, M L >_ 2.5, M L > 3.5 and M L > 4,0, for the years 1984, 1982, 1971 and 1954, respectively. The inflection points reflect the dates when data were incomplete (arrows).

A. Hofstetter et al. / Tectonophysics 267 (1996) 317-330

322

3.2. Relative location

event hypocentres are closely spaced, based on the less accurate absolute location; (b) the focal mechanisms are roughly identical. The second assumption can often be checked by rough first-motion focal mechanism determinations. Below we apply this technique for cluster 3 (Table 1) and we find that both assumptions are satisfied. For one cluster, consisting of 6 events (cluster 3), we obtained the relative hypocenter locations using cross-correlation methods similar to those used by Poupinet et al. (1984), Ito (1985), Deichmann and Garcia-Fernandez (1992) and others. Cluster 3 consists of a number of events in the upper crust with a depth range of 2 to 5 km that are qualitatively confirmed as they were felt locally. The events occurred in a small graben along the Carmel-Tirtza fault system and involve normal faulting (see Tables 1 and 2, Figs. 1-3).

In relative event location (i.e. Douglas, 1967; Poupinet et al., 1984; Ito, 1985) one can use the relative arrival times of the different events recorded on the same station instead of the absolute arrival times, where a master event serves as a reference. Consequently only the small travel-time differences due to path differences near the source are considered, while the main part of the source-station raypaths are assumed to be the same. An obvious advantage of this method is that complicated crustal structure modeling can be avoided. Moreover, seismogram signals from different events recorded on the same station, that display a high degree of similarity, can be used to obtain accurate arrival time differences by applying cross-correlation techniques. In doing this one implicitly assumes that: (a) the

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Magn i rude Fig. 4. Incremental and cumulative f r e q u e n c y - m a g n i t u d e relation earthquakes during 1900-1994.

for

earthquakes along the C a r m e l - T i r t z a fault system for observed

A. Hofstetter et al. / Tectonophysics 267 (1996) 317-330 3.3.

Seismicity

rate

The frequency-magnitude relationship (Fig. 4) is determined by using earthquake data from the Carmel and Tirtza fault systems and their intersections with the Jordan rift valley. The data include earthquakes of known locations and magnitudes for the period 1900 to 1994. Hanks (1992) has discussed the physical significance of large and small earthquakes. Assuming that the observed seismicity is complete for the whole period, and can be characterized by a Poisson process, we may use the following log-linear relationship: lOgl0 n

= a -

bM e , Mma x >

ME>

(1)

Mmi n

where n is the number of events of magnitude M L in the time interval 1900-1994. The magnitudes Mai n and Mmax are the lowest detectable magnitude and the maximum probable magnitude, respectively. Based on the available earthquake catalogs (IPRG Seismological Bulletins, 1984-1994; Turcotte and Arieh, 1986) for the analyzed region (Fig. 1) and knowing the seismic monitoring history in Israel throughout the whole time period, i.e., the installation of new seismic stations (Station EIL as part of the WWSSN, in 1954) or national seismic network (ISN, in 1982) and later the upgrading of the latters (station EIL in 1971, and ISN in 1984), we infer that the 95 years are divisible into a discrete number of time intervals. For each time interval we can assign a different operative threshold magnitude, above which

323

the data become complete. Fig. 3 shows the variation of the mean annual frequency with respect to the date interval dT = 1994 - t, where t is a given year for different groups of data, i.e. M L > 1.8, M L > 2.5, M L >_ 3.5 and M L > 4.0, for the years 1984, 1982, 197l and 1954, until which date the data were incomplete (sloped line), respectively. Following Stepp (1973), Consentino et al. (1977), Weichert (1980) and Shapira (1983), we obtain the b-value for the Carmel-Tirtza fault system b = 0.89 _ 0.01. It is worth noticing that the b-value, typically to heterogeneous tectonic region, that includes compressional, extensional and transformal deformation, is practically the same as in previous studies of Ben-Menahem and Aboodi (1981) and Shapira and Feldman (1987) for different areas in and around Israel and for different time intervals.

4. Source mechanisms Assuming that the stronger events better represent the seismotectonics of the Carmel-Tirtza fault system, we use only earthquakes with local magnitude, M L >_ 3.0, for source mechanism determination (M L > 4 . 0 are from Salamon, 1993). In the case of one-component stations, the determination of source mechanism is commonly done by using P-wave first motions. The grouping of earthquakes into clusters, as described above, serves as a basis for combining many events into a composite solution (Brillinger et

Table 2 Fault parameters for c o m p o s i t e a n d single event solutions (Table 1). Event cluster n u m b e r s are referred to in the text No.

! 2 3 4 5 6 7 8 9 10 11 12 13

Obs

14 93 45 25 17 35 76 33 35 15 36 31 41

P

T

Fault plane

A u x i l i a r y plane

dip

azim

dip

azim

strike

dip

rake

strike

dip

rake

31 2 73 ll 42 21 62 68 3 41 10 10 10

319 87 34 132 79 76 84 314 264 76 304 102 328

4 48 14 52 33 33 10 12 10 10 3 17 14

51 180 182 237 206 181 193 189 354 176 214 195 236

1 325 100 15 141 311 255 301 130 120 - 11 - 30 13

71 60 60 65 85 82 41 35 85 70 80 85 72

- 26 40 - 80 50 - 60 40 - 130 - 59 10 - 40 5 20 3

100 212 260 258 240 215 125 85 39 226 85 238 - 78

65 56 31 46 30 50 60 60 80 52 80 70 87

- 160 142 - 106 144 - 170 170 -60 - 110 174 - 154 - 170 174 161

Misfit

Sta. dist.

0.07 0.09 0.00 0.12 0.00 0.03 0.04 0.06 0.17 0.13 0.08 0.00 0.00

0.57 0.73 0.70 0.63 0.57 0.74 0.69 0.66 0.55 0.66 0.64 0.76 0.61

324

A. H@tetter et al. / Tectonophysics 267 (1996) 317-330

al., 1980; Udias et al., 1982) as presented in Figs. 1 and 2 and in Tables 1 and 2. The focal depth calculations are often not well constrained as those of the epicenter locations and, consequently, a few fault plane mechanisms are examined in parallel using a range of possible depths. Only minor changes in the fault plane solutions are observed, which means that the solutions are stable with respect to the assumed model of wave propagation, in most of the cases there is a fairly good distribution of first motion observations, except in cases 5 and 6 for which only one nodal plane is well constrained (Fig. 2). In general, the focal mechanisms d o s e to the junction of the Dead Sea Transform and the Carmel-Tirtza fault are characterized mainly by normal faulting with some strike slip motion (Table 2 and Figs. 1 and 2). Towards the eastern Mediterranean Sea the dominant mechanism changes to strike slip motion with an element of compression in all the solutions, in agreement with the surface faulting (Ron, 1984; Achmon, 1986; Ben-Gai, 1989) and reflection studies (Rotstein et al., 1993). The trend of general seismic activity and the obvious correlation with surface fault suggest that the active planes are in the northwest-southeast direction. Some fault plane solutions indicate north-south motion (clusters 1, 4, 1 I and 13 in Tables 1 and 2 and Figs. 1 and 2) in the vicinity of the junction of the Carmel-Tirtza fault and Dead Sea rift system, that may be related to induced motion of the latter fault. The Gilboa fault is characterized by normal faulting (cluster 7 in Tables 1 and 2 and in Figs. 1 and 2). We use two statistical criteria to judge the reliability of the composite solutions, as proposed by Reasenberg and Oppenheimer (1985), and also applied by Van Eck and Hofstetter (1989, 1990) and Salamon (1993). The misfit parameter qualitatively measures the fitness of the fault plane solution to the available data. It varies from 0 which is a perfect fit to 1 which is a complete misfit. In general, we obtain an excellent fit ( < 0.17), in spite of the incomplete azimuthal coverage in some of the cases. The station distribution parameter describes the percentage of readings close to the nodal planes, where the value 0 means many stations close to the nodal plane, and 1 is for no stations in the vicinity of the nodal plane. Robustness is considered to be achieved if the value

is above 0.5. All the cases passed this test with values > 0.55. Consequently, our results (Table 2), with misfit parameter < 0.17 and a station distribution parameter >_ 0.55, show a very good fit to the available data. Therefore, we consider the presented composite solutions as being reliable. 4.1. Stress tensor inuersion

A first-order approximation of the present state of stress can be obtained by assuming one stress tensor for the whole region of investigation, i.e., the Carmel-Tirtza fault system. Therefore we performed a simple stress tensor inversion of all available focal mechanism solutions (Table 2) using the methodology of Gephart and Forsyth (1984). Details of the method, we used the so-called "approximate" method, are explained in Gephart (1990). Fig. 5 shows the final result (or I = 2 8 7 / 1 6 ; ~r~ = 115/74; % = 18/2), i.e., the best-fitting stress tensor, of a limited search. An earlier more extensive grid search indicated the area of interest for a more detailed inversion as presented in Fig. 5. A N N E - S S W near horizontal tensional axis crs, and a W N W - E S E near horizontal compressional axis % seem to model well the observed focal mechanisms.

5. Spectral analysis and spectral scaling We use spectra of S-waves recorded by the vertical seismometers, along with a few added horizontal components to estimate the seismic moment, M 0, the corner frequency, .fi> stress drop, Act, and the source radius, r 0, based on the dislocation model of Brune (1970, 1971). A detailed description of the selection of S-wave signals and estimation of the displacement spectra is given by Lee and Stewart (1981). Its application to the Israeli seismic network is described by Van Eck and Hofstetter (1989) and Shapira and Hofstetter (1992). Here we follow briefly similar arguments mainly given by Shapira and Hofstetter (1992) in a recent study of spectral parameters of earthquakes in Israel and its vicinity. 5.1. Seismic" m o m e n t determination

The seismic moment, M 0, can be estimated as: M o = 4rrpDfi sO,/(cF)

(2)

A. Hofstetter et al. / Tectonophysics 267 (1996) 317-330

where, p is the density, 2700 k g / m 3, [3 is the average S-wave velocity, 3600 m / s (Ginzburg et al., 1979); D is the distance correction, D = R ~ e ~R,

325

where R is the distance, the theoretical value of e~ is 5 / 6 (Ewing et al., 1957; Nuttli, 1978) and the empirically found value of ~ is 3.8 × 10 -3 km -~

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[] •

cyl b e s t fit (Y3 b e s t fit

misfits: 4 o _ 9° 9 ° _ 14 ° 14°-460

Fig. 5. Stress tensor inversion using a grid search (Gephart and Forsyth, 1984) and 13 focal mechanism solutions for the Carmel-Tirtza fault system (Table 3). The underlying assumption is that all focal mechanisms along the fault system can be explained by a crustal stress pattern characterized by one stress tensor. The robust solution shows a horizontal N N E - S S W extension and a W N W - E S E compression.

A. Hofstetter et al. / Tectonophysics 267 (1996) 317-330

326

for Israel and its vicinity (Shapira and Hofstetter, 1992); 1] v is the zero frequency spectral level of the vertical ground displacement; c is the free surface correction, assumed to be equal to 2 and F is the

S-wave radiation pattern correction factor, where values of 0.18 and 0.61 are used for the SV-waves recorded on the vertical seismographs and for the S-waves recorded on the horizontal seismographs,

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Fig. 6. (A) Seismic signal with preceding background noise of an event that occurred along the Mt. Carmel fault system (event 6 in Table 3). The 10 s solid bar above the S-wave indicates the time window used for the site displacement spectra calculation. (B) The spectra of S-wave ground displacement and 10 seconds of the preceding background noise of the event in (A). Also shown the observed low frequency site spectral level, ~ v , and the observed corner frequency, f[~, as described in the text.

A. Hofstetter et al./ Tectonophysics 267 (1996) 317-330

327

Table 3 Dynamic source parameters of earthquakes that were used in the spectral analysis, where the date is composed of year, month, day, hour, and minute, M L is the local magnitude, M 0 is the seismic moment, Act is Brune's stress drop estimate, f0 is the corrected comer frequency and r 0 is the source radius according to Brune's formulation. The number in parenthesis is the associated uncertainty factor. The seismic moment of the second earthquake (8408240602) is taken from the ISC (1984) No. 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15

Date 8405301016 8408240602 8411041924 8411050144 8411050451 8411050514 8801300259 8803021445 8809051433 9012072024 9103010751 9112290138 9207290530 9303260657 9409160318

Latitude

Longitude

N

E

32.3 32.7 32.1 32.1 32.1 32.1 32.2 33.2 33.1 32.5 32.7 32.1 32.3 32.2 32.1

35.4 35.2 35.4 35.4 35.4 35.4 35.5 34.7 34.6 35.3 35.2 35.5 35.5 35.3 35.5

respectively ( B o o r e and Boatwright, 1984). illustrates a seismic signal and the associated d i s p l a c e m e n t s p e c t r u m for an e v e n t in cluster M o values range f r o m 0.7 × 1014 to 1017 N m 3), w h i c h in general are c o m p a t i b l e with the o f Shapira and Hofstetter (1992).

ML 3.2 5.3 3.6 3.8 3.6 3.5 3.8 3.1 3.5 3.1 3.3 3.2 3.4 3.1 4.l

Fig. 6 ground 3. The (Table values

fo

r0

No.

MPa

Hz

m

Obs.

4.13 (2.8)

5.1 (1.2)

262

19

2.49 (2.4) 3.32 (1.8) 1.31 (2.8) 1.94 (3.0) 9.57 (1.8) 16.13 (2.0) 16.42 (1.6) 6.21 (1.8) 4.60 (4.1) 3.36 (2.2) 0.87 (1.9) 1.04 (2.1) 9.36 (1.5)

3.0 (1.2) 2.7 (1.1) 2.6 (1.3) 3.0 (1.3) 3.1 (1.4) 3.8 (1.3) 2.9 (1.3) 4.6 (1.2) 3.2 (1.5) 3.5 (1.3) 3.0 (1.4) 4.7 (1.5) 3.4 (1.1)

437 490 522 446 428 353 458 287 415 382 424 269 386

11 3 14 13 3 13 7 19 15 23 13 22 3

1.7 (2.2) 1000 4.9 (2.7) 9.1 (2.5) 4.3 (2.7) 4.0 (3.3) 17.5 (1.5) 16.5 (2.0) 36.7 (1.7) 3.4 (1.9) 7.6 (2.2) 4.4 (1.9) 21.2 (1.5) 0.7 (1.6) 22.1 (1.6)

w h e r e 13s is the rupture velocity. A v e r a g e Ao-, for a g i v e n earthquake, are s h o w n in Table 3.

5.4. Source radius estimation

ro = a(j~Jfo)

The corner f r e q u e n c y using the distance correction m a y be f o r m u l a t e d as: (3)

where fo is the c o r r e c t e d c o r n e r frequency, f~ is the m e a s u r e d corner f r e q u e n c y at distance R (km), and 3' = 4.6 X 10 -4 k m - 1 (Shapira and Hofstetter, 1992). T h e corrected c o r n e r frequency, f0 varies b e t w e e n 2.6 to 5.1 Hz (Table 3), and with a g o o d a g r e e m e n t with Van E c k and Hofstetter (1989) and Shapira and Hofstetter (1992).

5.3. Stress drop estimation F o l l o w i n g the m o d e l o f B r u n e (1970, 1971) the static stress drop can be c o m p u t e d f r o m the equation:

A o - = 8.47Mo( f o / ~ s ) 3

A~r

1014 Nm

The source radius can be estimated from:

5.2. Corner f r e q u e n c y correction

)Co = f ; e z'l¢

M0

(4)

(5)

where the p a r a m e t e r a gets the value 0.37 ( B m n e , 1970), (Brune, 1971). Typical source radius values are b e t w e e n 260 m to 520 m (Table 3). The source radius values are similar to those obtained in California (Thatcher and Hanks, 1973) or those of Israel (Van E c k and Hofstetter, 1989; Shapira and Hofstetter, 1992).

6. D i s c u s s i o n a n d c o n c l u s i o n s

W e analyzed, quantitatively and qualitatively, m o r e than 550 small earthquakes along the C a r m e l Tirtza fault r e c o r d e d during an e l e v e n year period ( 1 9 8 4 - 1 9 9 4 ) . S e i s m i c m o m e n t s b e t w e e n 0.7 X 1014 and 1017 N m h a v e been calculated for earthquakes o f 3.1 < M L < 5 . 3 (Table 3), which are in g o o d a g r e e m e n t with observations f r o m the D e a d Sea basin or Israel (Van Eck and Hofstetter, 1989; Shapira

328

A. Hqfi~'tetter et al. / Tectonophysics 267 (1996) 317-330

1020 E

10'9

'

I

'

'

o / . /

M0 versus ML

Zo 10'8 .--

@~f/~ x~~@~

10 ~7 -

10'6 E

•

0

10 ~

g: .m_

10

E .o

" ' ~;.

,•

,,% ::l,..,/i

" ~

~ @.q,×

TM

U) 10'3

o•

":"Y/4" "

,o<

. :: •

i.

!

"

•

I

2

3

4

i

5

magnitude(M,)

6

7

Fig. 7. Local magnitude, MI, versus seismic moment, M o, for events in Israel and its surroundings. The observations include those in Van Eck and Hofstetter (,1989) and Shapira and Hofstetter (1992). The thick line indicates the magnitude seismic moment relation of Hanks and Kanamori (1979). The ligther line indicates the Mt~ M0 relation as obtained by Van Eck and Hofstetter (1989) with a small additional term of 0.2 on the right hand side.

and Hofstetter, 1992), or southern California (Thatcher and Hanks, 1973; Bakun, 1984). In Fig. 7 we show the relation between M 0 and M L for all events within Israel and its surroundings with M L and M 0 estimates. The observations are from our study (Table 3) and those of Van Eck and Hofstetter (1989) and Shapira and Hofstetter (1992). For larger events the relation of Hanks and Kanamori (1979) fits well. For smaller events the relation in Van Eck and Hofstetter (1989) slightly modified seems to fit the observations. Bakun (1984) and Hanks and Boore (1984), and also apparent in Fig. 7, showed that there is an increase of the magnitude coefficient (in the m o m e n t - m a g n i t u d e relation) with the increase of the earthquake magnitude from 1.2 to 1.5. The area, which exhibits relatively high seismicity, had a fairly normal rate of activity during the observation time (Fig. 1). The overall seismicity follows the main trend of the Mr. C a r m e l - T i r t z a fault. In the southeastern part, we distinguished a rather diffused activity along the series of active Tirtza faults. In the northwestern part of the fault, the seismicity is located in a rather wide deformation zone, as sug-

gested also by results of a recent reflection study (Rotstein et al., 1993). Although earthquakes along the Dead Sea transform accommodate most of the relative plate motion, we can interpret the seismic activity along the Mt. C a r m e l - T i r t z a fault system as part of this interaction, involving deformation of the brittle crust. The lateral extent of this deformation is delineated by the seismic activity, as suggested also by Salamon (1993). Mr. Carmel fault appears to be a boundary between two micro plates. South of the fault the micro plate is stable and deformation is close to the Dead Sea transform, while north of the fault the deformation is distributed in a wide range. In general, the fault should be viewed as a mechanism to absorb part of the deformation of the Dead Sea transform, due to the relative motion between the Sinai sub-plate and the Arabian plate. Regarding the San Andreas fault system and its associated faults, similar interpretation of diffused activity close to the junction of branching fault and deformation along a broad region encompassing the fault, is given by Dickinson (1981) and Hill et al. (1990). The dominant source mechanism changes along the Mt. C a r m e l - T i r t z a fault system, as was suggested also by Salamon (1993). Normal faulting is apparent close to the junction of the C a r m e l - T i r t z a fault and the J o r d a n - D e a d Sea transform. The earthquakes with strong normal faulting components occurred generally in or along local grabens. Along the northwestern part of the C a r m e l - T i r t z a fault the northwest-southeast left lateral strike slip motion is dominant with an element of compression, in good agreement also with results of a reflection study (Rotstein et al., 1993). Geological and tectonical studies suggested a complicated break-up pattern around this fault branch and that north of the Mt. C a r m e l - T i r t z a fault the stress field is of complex nature (Ron, 1984; Achmon, 1986; Ben-Gai, 1989). Recently Salamon (1993) calculated the stress field based on geological data since mid-Miocene until recent time (Ron, 1984) and also the seismicity data ( M > 4), and found it to be of normal faulting (o- u = 3 3 3 / 5 7 ; o ' 2 = 140/31; ( Y 3 = 2 3 4 / 5 ) . This stress field is compatible with the calculated focal mechanisms in our study (~Yl = 2 8 7 / 1 6 ; (r~= 115/74; ~r3 = 18/2). Historical recorded seismicity does not suggest that large earthquakes occurred on or nearby the

A. Hofstetter et al./ Tectonophysics 267 (1996) 317-330

Carmel-Tirtza fault (Turcotte and Arieh, 1986; Ambraseys and Karcz, 1992). Based on Shapira and Shamir (1994) we may estimate a maximum magnitude of Mmax = 6.5. Using the relations log M 0 = 16 + 1.5M (Hanks and Kanamori, 1979), as verified for regional events in Israel (Shapira and Hofstetter, 1992), M 0 = ~uLW= ~cxL2W (Kanamori and Anderson, 1975), where the modulus of rigidity is = 3 × 1 0 1 ° N / m 2, c~= 1.4× 10 -5 , and W = 1 5 km is the width of the fault, we get that the displacement is u = 0 . 4 2 m and the fault length is L = 3 0 kin, and the latter is in agreement with the on-surface geological faults (Fig. 1). In conclusion we may say that the Carmel-Tirtza fault system is a seismically active branch of the Jordan-Dead Sea transform fault. This study shows that from a relatively short period of observation (11 yr) one obtains a fairly good indication which are the seismic active fault segments and what is the corresponding stress pattern: dominant left lateral strike slip along the northwestern part and a complex tectonics on the southeastern part. The observations of this study, that include the epicenter data, seismic moment calculations and focal mechanism determinations, are needed for any research dealing with the estimation of seismic energy release in connection with geological analysis of tectonic deformations and the estimation of potentially dangerous zone. Though the observation period (1984-1994) is fairly short, it is enough to obtain more than a just general idea of the type of seismic activity in this region, and to improve our understanding of potentially intermediate-size damaging earthquakes.

Acknowledgements C. Ben-Sasson, L. Feldman and B. Reich did the initial data processing. D. Kadosh, D. Levi, U. Peled and Y. Schwartz kept the seismic network operating. The constructive comments of both reviewers are appreciated. This is a joint study supported by both the Earth Science Research Administration, Ministry of Energy, Israel and the Geodynamics Research Institute, University of Utrecht, The Netherlands (publication 95.046).

329

References Achmon, M., 1986. The Carmel Border Fault. M.Sc. Thesis, The Hebrew Univ., Jerusalem, Israel (in Hebrew with English abstract). Amhraseys, N. and Karcz, I., 1992. The earthquake of 1546 in the Holy Land. Terra Nova, 4: 253-262. Bakun, W.H., 1984. Seismic moments, local magnitudes, and coda-duration magnitudes for earthquakes in central California. Bull. Seismol. Soc. Am., 74: 439-458. Bartov, Y., Steinitz, G., Eyal, M. and Eyal, Y., 1980. Sinistral movement along the Gulf of Aqaba--Its age and relation to the opening of the Read Sea. Nature, 285: 220-221. Ben-Avraham, Z. and Ginzburg, A., 1990. Displaced terranes and crustal evolution of the Levant and the eastern Mediterranean. Tectonics, 9: 613-622. Ben-Gai, Y., 1989. The Border Faults of the Carmel Structure on the Continental Shelf off Northern Israel. M.Sc. Thesis, Tel Aviv Univ., Tel Aviv, Israel (in Hebrew with English abstract). Ben-Menahem, A., 1979. Earthquake catalogue for the Middle East. Boll. Geofis. Teor. Appl., 21: 245-313. Ben-Menahem, A. and Aboodi E., 1981. Micro- and macroseismicity of the Dead Sea rift and off-coast eastern Mediterranean. Tectonophysics, 80: t99-233. Boore, D. and Boatwright, J., 1984. Average body-wave radiation coefficients. Bull. Seismol. Soc. Am., 74: 1615-1621. Brillinger, D.R., Udias, A. and Bolt, B., 1980. A probability model for regional focal mechanism solutions. Bull. Seisol. Soc. Am., 70: 149-170. Brune, J.N., 1970. Tectonic stress and the spectra of seismic shear waves from earthquakes. J. Geophys. Res., 75: 4997-5009. Brune, J.N., 1971. Correction. J. Geophys. Res., 76: 5002. Consentino, P., Fierrarra, V. and Luzio, D., 1977. Truncated exponential frequency-magnitude relationship in earthquake statistics. Bull. Seismol. Soc. Am., 67: 1615-1624. Deichmann, N. and Garcia-Fernandez, M., 1992. Rupture geometry from high-precision relative hypocentre locations of microearthquake clusters. Geophys. J. Int., 110:501-517 Dickinson, W.R., 1981. Plate tectonics and the continental margin of California. In: G.W. Ernst (Editor), The Geotectonic Development of California. Prentice Hall, Englewood Cliffs, pp. 2-28. Douglas, A., I967. Joint epicentre determination. Nature, 215: 47-48. Ewing, W.M., Jardetzky, W.S. and Press, F., 1957. Elastic Waves in Layered Media. McGraw Hill, New York. Freund, R., Garfunkel, Z., Zak, I., Goldberg, M., Weissbrod, T. and Derin, B., 1970. The shear along the Dead Sea Rift. Philos. Trans. R. Soc. London, Ser. A, 267: 107-130. Garfunkel, Z. and Almagor, G., 1985. Geology and structure of the continental margin off northern Israel and the adjacent part of the Levantine basin. Mar. Geol., 62: 105-131. Gephart, J.W. and Forsyth, D.W., 1984. An improved method for determining the regional stress tensor using focal mechanism data: application to the San Fernando earthquake sequence. J. Geophys. Res., 89: 9305-9320.

330

A. H@tetter et al. / Tectonophysics 267 (1996) 317-330

Gephart, J.W., 1990. FMSI: A FORTRAN program for inverting fault/slickenside and earthquake focal mechanism data to obtain the regional stress tensor. Comput. Geosci., 16: 953989. Ginzburg, A , Makris, J., Fuchs, K., Prodehl, C., Kaminski, W. and Amitai, U., 1979. A seismic study of the crust and upper mantle of the Jordan-Dead Sea rift and their transition toward the Mediterranean sea. J. Geophys. Res., 84: 1569-1582. Hanks, T.C., 1992. Small earthquakes, tectonic forces. Science, 256: 1430-1432. Hanks, T.C. and Boore, D.M., 1984. Moment-magnitude relations in theory and practice. J. Geophys. Res., 89: 6229-6235. Hanks, T.C. and Kanamori, H., 1979. A moment magnitude scale. J. Geophys. Res., 84: 2348-2350. Hill, D., Eaton, J. and Jones, L., 1990. Seismicity, 1980-86. In: R. Wallace (Editor), The San Andreas Fault System, California. USGS Prof. Pap., 1515: 115-151. Hofstetter, A., Fetdman, L. and Rotstein, Y., 1991. Crustal structure of Israel: constraints from teleseismic and gravity data. Geophys, J. Int., 104: 371-380. IPRG Seismological Bulletins, 1984-1994. Earthquakes in and around Israel, v. 3-12, Seismological Division, Institute for Petroleum Research and Geophysics, Holon, Israel. ISC, 1984. Regional Catalogue of Earthquakes. International Seismological Centre Bulletin, Berkshire, United Kingdom. Ito, A., 1985. High resolution relative hypocentres of similar earthquakes by cross-spectral analysis method. J. Phys. Earth, 33: 279-294. Kanamori, H. and Anderson, D.L., 1975. Theoretical basis of some empirical relations in seismology. Bull. Seismol. Soc. Am., 65: 1073-1095. Lee, W.H.K. and Stewart, S.W., 1981. Principles and Applications of Microearthquake Network. Academic Press, New York. Nuttli, O., 1978. A time domain study of the attenuation of 10-Hz waves in the New Madrid seismic zone. Bull. Seismol. Soc. Am., 68: 343-355. Poupinet, G., Ellsworth, W.L. and Frechet J., 1984. Monitoring velocity variations in the crust using earthquake doublets: An application to the Calavaras fault, California. J. Geophys. Res., 89: 5719-5731. Reasenberg, P. and Oppenheimer, D., 1985. FPFIT, FPPLOT and FPPAGE: Fortran Computer Programs for Calculating and Displaying Earthquake Fault-plane Solutions. USGS, Open File Rep. 85-739. Ron, H., 1984. Paleomagnetic Investigation of the Fault Structure of Galilee - - Northern Israel. Ph.D, Thesis, The Hebrew Univ., Jerusalem, Israel (in Hebrew with English abstract).

Ron, H., Nur, A. and Hofstetter, A., 1991. Late Cenozoic and recent strike slip tectonics in Mr. Carmel, northern Israel. Ann. Tecton., 4: 70-80. Rotstein, Y.~ 1987. Gaussian probability estimates for a large earthquake occurrence in the Jordan Valley, Dead Sea rift. Tectonophysics, 141 : 95-105. Rotstein, Y., Bruner, I. and Kafri, U., 1993. High-resolution seismic imaging of the Carmel fault and its implications for the structure of Mr. Carmel. Isr. J. Earth Sci., 42: 55-69. Salamon, A., 1993. Seismotectonic Analysis of Earthquakes in Israel and Adjacent Areas. Ph.D. Thesis, The Hebrew Univ., Jerusalem, Israel (in Hebrew with English abstract). Sbapira, A., 1983. A probabilistic approach for evaluating earthquake risks, with application to the Afro-Eurasian junction. Tectonophysics, 91: 321-334. Shapira, A., 1988. Magnitude scales for regional earthquakes monitored in Israel. Isr. J. Earth Sci., 37: 17-22. Shapira, A. and Feldman, L., 1987. Microseismicity of three locations along the Jordan rift. Tectonophysics, 141: 89-94. Shapira, A. and Hofstetter, A., 1992. Source parameters and scaling relationships of earthquakes in Israel. Tectonophysics, 217: 217-226. Shapira, A. and Shamir. G., 1994. Seismicity Parameters of Seismogenic Zones in and around Israel. Inst. Petrol. Res. Geophys., Rep. Z 1/567/79(109). Stepp, J.C., 1973. Analysis of completeness of the earthquake sample in the Puget Sound area. In: Contributions of Seismic Zoning. Natl. Oceanic Atmos. Adm. Tech. Rep. ERL 267ESL30, 94 pp. Thatcher, W. and Hanks, T., 1973. Source parameters of southern California earthquakes. J. Geophys. Res., 78: 8547-8576. Turcotte, T. and Arieh, E., 1986. Catalog of Earthquakes in and around Israel. 1PRG Report for the Israel Electric Corporation Ltd., June 1986. Udias, A., Buforn, E., Brillinger, D. and Bolt, B., 1982. Joint statistical determination of fault-plane parameters. Phys. Earth Planet. Int., 30: 178-184. Van Eck, T. and Hofstetter, A., 1989. Microearthquake activity in Dead Sea region, Geophys. J. Int., 99: 605-620. Van Eck, T. and Hofstetter, A., 1990. Fault geometry and spatial clustering of microearthquakes along the Dead Sea-Jordan rift fault zone. Tectonophysics, 180: 15-27. Weichert, D.H., 1980. Estimation of the earthquake recurrence parameters for unequal observation periods for different magnitudes. Bull. Seismol. Soc. Am., 70: 1337-1346.