Recent and present-day stresses in the Granada Basin (Betic Cordilleras): Example of a late Miocene-present-day extensional basin in a convergent plate boundary

TECTONICS, VOL. 18, NO. 4, PAGES 686-702, AUGUST 1999

Recent and present-day stresses in the Granada Basin (Betic Cordilleras)' Example of a late Miocene-present-day extensional basin in a convergent plate boundary

J. Galindo-Zaldivar, • A. Jabaloy, • I. Serrano, • J. Morales? F. Gonzfilez-Lodeiro, • and F. Torcal •

Abstract. The diffuse convergent boundary between the Eurasian and African plates in the western Mediterranean is associated with a seismicity zone more than 300 km wide. Although the two plates are converging NW-SE, the Betic and Rif Cordilleras contain extensional structures that have been active since the Miocene. The extensional tectonics in the re-

gion, which occurred simultaneously with the uplift of the cordillera, have been analyzed in the southeastern sector of the late Miocene to recent Granada Basin, using earthquake focal mechanisms, the determination ofpaleostresses from the study of the orientation and kinematics of microfaults, and the study of the major structures. Both the geological surface data and the focal mechanisms indicate present-day regional conditions of NE-SW extension, with triaxial to prolate stress ellipsoids. However, the stress field is heterogeneous, with local varia- tions in stress over time, with different stresses sometimes even acting simultaneously in adjacent areas. The most fre- quent changes consist of pluridirectional or NE-SW extension, favored by the prolate character of the stress ellipsoids, and NW-SE subhorizontal compression, favored by the regional tectonic setting. Strike-slip faults are scarce even though they are the most likely structure to be expected in a region with SW-NE extension and NW-SE compression. Seismicity is concentrated in the upper crust and may correspond to the ac- tivity of low- to high-angle normal faults similar to the surface faults, although they can not be correlated with them. The lower cutoff of this seismicity probably coincides with the 300øC isotherm and suggests a low thermal gradient for the region. Present-day regional stresses have O l vertical at the surface but in depth plunge toward the SW.

1. Introduction

The Alborfin Sea Basin, filled with Neogene and Quater- nary rocks, lies between the Betic and Rif Cordilleras (Figure 1) over thinned continental crust [Working Group For Deep Seismic Sounding in the Albor•n Sea 1974-1975 (WGDSSAS

'Departamento de Geodinfimica. Universidad de Granada, Gra- nada. Spain.

:Instituto Andaluz de Geofisica, Universidad de Granada, Gra- nada. Spain.

•Also at Departamento de Fisica Te6rica y del Cosmos. Universi- dad de Granada. Granada, Spain.

Copyright 1999 by the American Geophysical Union.

Paper number 1999TC900016. 0278-7407/99/1999TC900016512.00

1974-19*5), 1978; Surihach and Vegas, 1993]. Seismic activ- ity in the Betic-Rif Cordilleras and in the Alborfin Sea is asso- ciated with the contact between the Eurasian and African

plates. In the study area the convergence between these two plates is NW-SE at a rate of 4 mm/yr [DeMets et al., 1990]. The spatial distribution of the earthquakes outlines a well- defined plate boundary between 30øW (Azores) and 12øW. In contrast, the seismicity between 12øW and 2øE (Algeria) is dispersed, and the plate boundary is diffuse. The zone delim- ited by the seismicity reaches a maximum width of 300 km [/lnderson and clackson, 1987; Bz.tforn et al., 1988], with earthquakes in this region attaining magnitudes of M, < 6.0 [Bz(forn eta/., 1988, 1995; Calvert et al., 1997].

Since the Miocene, both the A!borfin Sea and the Internal Zones of the Betic-Rif Cordilleras have undergone extensional tectonic activity that has produced thinning of the previously thickened continental crust [Platt and Vissers, 1989; Docherty and Banda, 1995; Gonzglez-Lodeiro et al., 1996]. However, the coeval development of compressional structures can also be observed: folds and thrusts in the external region of the cordilleras and folds and strike-slip faults in the Internal Zones [Galindo-Za/dh'ar et al., 1993]. In recent years, several geo- dynamic models have been proposed to explain continental extension in the Alborfin Sea within a general compressional context: rollback processes [Morley, 1993; Lonergan and White, 1997], mantle diapirism [Weijermars, 1985], exten- sional collapse [Platt and Vissers, 1989; Watts et al., 1993; l/issers eta/., 1995], and delamination [Docherty and Banda, 1995; Seber et al., 1996]. Watts et al. [1993], Docherty and Banda [1995], and Vissers et al. [1995] have discussed these models in detail. The analysis of the relationships between compressional and extensional structures has special interest in the study of continental extension in zones undergoing compression [Ruppe/, 1995].

The External Zones of the Betic Cordilleras consist of the

Prebetic and Subbetic Zones, comprising mainly Mesozoic and Cenozoic sedimentary rocks dominated by limestones and marls. The Internal Zones consist of several complexes: the Dorsal, Pre-Dorsal, and Alozaina, formed of Mesozoic and Cenozoic rocks, and another three complexes, the Nevado- Filfibride, Alpujfirride, and Malfiguide (Figure 1). The latter three contain, in addition, rocks dated or attributed to the Pa- leozoic and are eraplaced in the following order from bottom to top: Nevado-Filfibride, Alpujfirride, and Malfiguide. The Nevado-Filfibride and Alpujfirride have undergone Alpine metamorphism.

The term "Granada Basin" is given to an outcrop of Neo- gene to Quaternary sedimentary rocks lying over the NE-SW trending contact between the External and Internal Zones

686

GALINDO-ZALDiVAR ET AL.' STRESSES IN GRANADA BASIN, BETIC CORDILLERAS

,0 . 10,0.20,0 Km • INTERNAL ZONES J AFRICA

687

•7.0-

Alborfin Sea 20 km

41o Longitude (W)

INTERNAL ZONES

A!ozaina complexes

I 111111 complex /•••7• Nevedo-Filebride complex EXTERNAL ZONES

Subbetic zone

Neogene and Quaternary sedimentary rocks

Seismic station

o Earthquake epicenter

Paleostress determinations station

Figure 1. General geological sketch of the study area including epicenters, seismic network stations, and microfault measurement stations. A-B is the location of the cross section in Figure 2.

(Figure 1). The sedimentary sequence is over 2 km thick in some areas, containing lower Miocene to Recent rocks [Morales e! al., 1990]. The Tortonian marine rocks identified in many of the depressions of both the External and Internal Zones indicate that the Alborfin Sea covered a much greater area during the early Miocene, including the Granada Basin. From the upper Miocene to the present there has been rapid uplift of the cordilleras. The differential uplift of each area has determined basin boundaries since the Messinian and accounts

for the current uneven topography. The focal mechanisms in

the region, compiled by Galindo-Zaldivar et al. [1993], indi- cate the existence of normal faults, suggesting present-day extensional stress with a poorly defined direction of extension.

The main aim of this study is to analyze in detail the Recent and present-day stress state in the southeastern half of the Granada Basin and surrounding regions, on the basis of a combined study of the earthquake focal mechanisms and pa- leostresses determined from Recent structures. The study re- gion constitutes a good example of an extensional Neogene basin located at a convergent plate boundary. The existence of

688 GALINDO-ZALDiVAR ET AL.' STRESSES IN GRANADA BASIN, BETIC CORDILLERAS

a local seismic network has allowed us to analyze the present- day stresses and precisely locate the abundant microseismicity in the region with greater accuracy than that in previous works. We therefore compare and integrate field geology data with seismological data.

2. Seismicity in the Granada Basin

The Granada Basin and neighboring areas form the region with the highest rate of microseismic activity in the entire Ibe- rian Peninsula, with earthquakes of magnitude mh-< 5.5 [De Miguel eta/., 1989]. Although there is no instrumental record of earthquakes over 5.5 in magnitude, with the exception of the deep earthquake in 1954 m/,=7.1, M,,=7.8, and h=640 km [Chung and Kanamori, 1976], there are nevertheless historical data that clearly indicate the potential of some faults to cause destructive earthquakes [ Fidal, 1986].

Seismic series unassociated with a principal earthquake is one of the most notable characteristics of the seismicity in this region. The last series of 1979 [Fidal, 1986] covered the whole basin and surrounding areas (several tens of kilometers around the boundary of the depression) and lasted nearly a year. Earthquakes also occur as seismic swarms in which a great number of microearthquakes are concentrated in a very short period (days or weeks) and in a specific area (e.g., Loja, in 1985; Agr6n, in 1988).

The seismic activity is fundamentally located in the upper crust. Most of the earthquakes are generally deeper than 9 kin, and shallower than 16 km in the central sector, gradually deepening up to 25 kin in the southwestern sector [Morales et

a/., 1997] (Figure 2). There is also very deep activity (640 kin), with magnitudes between m/,:7.1 and m/,--4.5 in the years 1954, 1973, 1990 [BtqCorn et al., 1991] and 1993, although there is a lack of intermediate and deep earthquakes. The very deep earthquakes have been attributed to a detached lithos- pheric slab [Blanco and Spakman, 1993].

3. Earthquake Data Selection and Focal Mechanisms

To study present-day stresses from the solutions to the fo- cal mechanisms of microearthquakes and earthquakes, we have selected 43 events located on the southeastern edge of the Granada Basin (Figures 1 and 3a) with magnitudes of m•/>2.1 (Table 1). The earthquakes were registered by perlna- nent and portable seismic stations of the Andalusian Seismic Network belonging to the University of Granada (Figure 1). The seismic stations comprise Mark-L4C or field vertical seismometers with a natural period of To =1 s [Alguacil, 1986]. The signals recorded in the field are transmitted by telemetry to the recording center, which samples the signals at 100 sps (samples per second) via an analog to digital 12-bit converter. A short-term average / long-term average algorithm deter- mines the occurrence of an earthquake, which is then recorded digitally on the hard drive of a PC [Alguacil el al., 1990]. Al- though the coverage of the stations of the Andalusian SeislniC Network is certainly adequate for locating local earthquakes (Figures I and 3a), data from other stations throughout the zone belonging to other institutions (i.e. Instituto Geogrfifico Nacional, Ministerio de Folnento) have also been used in

• 20

SW

o o

A

10 Km

Horizontal scale: ! I

1

Figure 2. (a) Cross sections across the Granada Basin area, showing the depth distribution of the earthquake hypocenters (location in Figure 1' data froin Morales eta/. [1997]) and (b) the main geological features (location in Figure 3a). Dots show, earthquakes with a known focal mechanism.

GALINDO-ZALDiVAR ET AL.' STRESSES IN GRANADA BASIN, BETIC CORDILLERAS 689

A

12 25

1

22

I

3

,, 0

LEGEND

Subbetic zone

Malaguide complex Alpujarride complex

Nevado-Filabride complex Neogene and Quaternary I rocks Axial trace of synform Axial trace of antiform Normal fault Extensional detachment

Granada 39

42• •

33 0 19 3C

..

29

,

1

10 km !

4'00 3'80 3'60 3'40

Longitude (W)

N N N N

GS PINOS O O

B . ".:':':':':';"Z: t-- .i... x ,,,,, m k• -'"""•"_a_Granada ,?","f,•h•'-' •• G S

[-. IO'•;-, •-•---'•:•: .... .-..: :% ,.,%--

%',' ........ •,' ...... ?'&'• -'?' N

,' ,' ,' ,' ,' ,- ,' ,' ,' ̀, ,",-•, ................. ,- ,- ,-,-` ,,, g _ • '5',',','}}},',',',',',',',',',' ................. 10 km

/ 4'00 ,/• 6/ 3'Z• Lt•ngitude (W) % • N N 3' N

RI ZAFA NI 02 O

Figure 3. Geological sketches of the southern and central Granada Basin showing (a) earthquake focal mechanisms and (b) microfault measurement sites. I-I' is the location of the cross section in Figure 2. A to F point the location of the field examples illustrated in Figure 6. Diagrams are in stereographic projection lower hemisphere. Focal mechanisms show compression in white. Striae symbols in microfault diagrams indicate the following: normal faults (squares), reverse faults (triangles), and faults with unknown sense of motion (dots).

690 GALINDO-ZALDiVAR ET AL.' STRESSES IN GRANADA BASIN, BETIC CORDILLERAS

Table 1. Data of the Earthquakes Used in This Work Event Date Time, Latitude, Longitude, H Error, Depth, Z Error. Magnitude. rms.

UT øN øE km km km ma (s)

I June 24, 1984 1430 36.860 -3.730 0.49 11.3 1.1 4.9 0.12

2 May 2, 1988 1051 37.183 -3.608 0.84 15.1 1.6 3.2 0.10

3* July 18, 1988 2032 36.958 -3.965 0.46 28.1 0.5 3.7 0.06

4* July 25. 1988 0123 37.089 -3.821 0.46 11.1 0.7 2.9 0.08 5 Aug. 20, 1988 1642 37.179 -3.713 0.82 13.5 1.4 3.7 0.14 6 Sept. 15, 1988 0027 36.990 -3.737 0.76 10.6 1.3 2.7 0.11 7* Oct. 21, 1988 1205 37.022 -3.895 0.74 13.6 1.9 2.7 0.17

8* Oct. 22, 1988 1613 37.024 -3.846 0.38 14.5 1.1 2.2 0.08

9 Nov. 14, 1988 1120 37.042 -3.819 0.76 9.9 0.5 2.6 0.12

10' Dec. 5, 1988 2012 37.031 -3.835 0.57 14.0 1.4 4.0 0.12 11' Dec. 6, 1988 0609 37.024 -3.848 0.56 13.2 1.4 3.2 0.16

12' Dec. 6, 1988 2015 37.029 -3.842 0.48 14.9 1.1 3.8 0.13

13' Dec. 7, 1988 1850 37.032 -3.842 0.49 17.1 0.9 3.6 0.12

14' Dec. 14, 1988 1700 37.051 -3.927 0.41 15.9 0.8 3.8 0.10

15 Dec. 25, 1988 1356 37.019 -3.674 0.78 13.8 1.0 3.6 0.15

16' Jan. 5, 1989 2018 37.019 -3.864 0.54 23.2 0.9 2.4 0.11

17' Jan. 5, 1989 2024 37.016 -3.866 0.67 23.3 1.3 2.9 0.15

18' Jan. 6. 1989 0633 37.016 -3.858 0.46 20.6 0.9 2.3 0.09

19' Jan. 12, 1989 1334 37.051 -3.807 0.92 9.5 0.9 2.2 0.20

20* Feb. 15, 1989 0004 36.979 -4.043 0.89 22.2 1.1 3.7 0.13

21 * Feb. 15, 1989 0154 36.996 -4.047 0.56 21.6 0.7 3.3 0.06

22* Feb. 15. 1989 0743 37.001 -4.037 0.79 15.5 1.1 2.1 0.19

23 Jul5 19. 1989 0430 37.020 -3.686 0.55 12.7 1.0 3.5 0.12 24 Oct. 7, 1989 0445 37.116 -3.632 0.61 12.9 0.9 3.5 0.12

25 July 1, 1990 2324 37.096 -3.823 0.48 13.3 0.9 3.2 0.17 26 Oct. 14. 1990 1602 37.012 -3.695 0.63 12.1 1.1 3.2 0.17

27 Nov. 7. 1990 0725 36.985 -3.713 0.90 14.0 1.4 4.2 0.18

28 Nov. 10, 1990 0121 37.004 -3.709 0.67 13.0 1.1 3.1 0.16

29 Nov. 23, 1990 1420 36.992 -3.860 0.46 13.9 1.0 2.8 0.08

30 June 21. 1991 2123 37.096 -3.581 0.62 11.0 1.3 3.1 0.17

31 July 6, 1991 1928 37.147 -3.460 0.64 10.1 1.3 2.4 0.17

32 Aug. 10, 1991 0934 36.988 -3.643 1.10 10.0 0.6 2.8 0.13 33 Nov. 12, 1991 0645 37.024 -3.853 0.40 13.5 0.7 3.3 0.15

34 Dec. 10. 1991 0351 36.921 -3.741 0.67 16.3 0.7 3.1 0.14

35 Jan. 12, 1992 0439 37.035 -3.584 1.18 17.0 1.4 3.0 0.18

36 Mar. 21, 1993 0817 37.112 -3.487 1.03 11.7 1.1 3.0 0.24

37 Mar. 21, 1993 1001 37.104 -3.495 1.13 10.1 1.4 3.6 0.24

38 June 19, 1993 2311 37.182 -3.727 0.49 14.8 0.9 3.0 0.12

39 Mar. 3, 1994 1309 37.186 -3.624 0.82 9.9 0.6 2.7 0.19

40 Aug. 8, 1994 2054 37.096 -3.431 0.95 9.3 0.4 2.6 0.19 41 Oct. 24, 1994 0301 37.000 -3.702 0.45 15.5 0.9 2.8 0.16

42 Dec. 6. 1994 0928 37.148 -3.589 0.77 11.6 0.8 3.2 0.15

43 dec. 23, 1994 1504 36.898 -3.857 1.14 7.1 1.2 2.9 0.20

H Error. horizontal error from location program; Z Error. vertical error from location program; nns. squared mean residual. * Earthquakes belonging to Agr6n area.

testing focal depths and determining the solutions of focal mechanisms.

The data utilized were gathered between 1988 and 1994, with rms and horizontal error of less than 0.25 s and 1.2 km, respectively (Table 1). Moreover, we have included one ear- lier earthquake which represents the largest-magnitude earth- quake recorded by instruments in the Granada Depression

(June 24, 1984). The order and distribution of earthquake oc- currence show seismicity in the area to be irregular (Figure 3a). However, 16 of the 43 earthquakes chosen for study are from the Agr6n area (Table ! and Figure 3a), and of these, 5 belong to a seismic swarm from December 5-!5, !988. This swarm began with an earthquake of m=4.0, releasing 290 mi- croearthquakes, of which only 59 were m>l.0. Earthquakes of

GALINDO-ZALDiVAR ET AL ß STRESSES IN GRANADA BASIN, BETIC CORDILLERAS 691

m>2.1 have been chosen to study this sector, owing to the palticularly good coverage of the seismic stations during this Z/ Z: Zs period. km km km

Taking into account the dependence of the focal depth and i 10.8 9.9 8.8 9.8 the velocity model used in our standard location program and 2 14.8 14.8 14.7 14.8 in order to obtain a more accurate focal depth, we have used 3 29.3 29.1 28.9 29.1 alternative methods [IVadati, 1933; Hales et al., 1981; Toth 4 12.1 11.6 1 i.0 11.6 and Kisslinger, 1984]. In these inethods the origin time (Ho)

5 13.5 13.5 13.4 13.5 of the earthquake and focal depth are estimated independently following the Wadati diagram. Following Hales e! al. [1981] 6 ii.2 11.0 10.9 il.0 and Toth and Kisslinger [ 1984], we calculated two estimations 7 15.3 i 5. i 14.8 i 5. i of the focal depth (Z• and Z2, Table 2). A third estimation of 9 9.8 9.6 9.4 9.6 the focal depth can be obtained using the method of Chapman 10 14.4 14.3 14.2 14.3 and Bollinger [1984] (Z•, Table 2). ii 13.9 13.3 12.5 13.2

Of the 43 earthquakes chosen, the depths of 38 have been 12 15.2 14.9 14.6 14.9 calculated with the afore-mentioned methods (Table 2) while 13 18.5 18.7 19.0 18.7 the remaining 5 (numbers 8, 16, 18, 19, and 22) had depths as- 14 17.2 17.1 i 7.0 17.1 signed by the location prograin (Table 1). In the 38 earth- 15 13.7 13.8 13.9 13.8 quakes to which the alternative methods were applied, the av- 17 23.1 23.2 23.4 23.2 erage of three estimates was taken as the true depth. Although

20 23.6 23.7 23.9 23.7 the focal depths given by the location program for earthquakes

21 22.4 22.5 22.7 22.5 in the Granada Depression are in accordance with the average estimates [Ibdhez, 1987; Morales et al.. 1997], there is never- 23 12.5 12.5 12.6 12.5 theless a decrease in the depth error (standard deviation) (Ta- 24 13.2 13.3 13.5 13.3 bles I and 2). 25 14.7 14.6 14.5 i 4.6

Once the focal depths had been constrained with the above 26 12.5 13.0 13.6 13.0 mentioned methods (Table 2), the focal mechanisms of the 27 16.7 16.2 15.7 16.2 chosen earthquakes were then determined using the first im- 28 i 5.4 15.1 14.6 15.0 pulse polarity of the P wave. The solution of the focal mecha- 29 15.2 15.5 15.9 15.5 nisln for low-magnitude earthquakes could be determined 30 11.5 10.6 9.5 10.5 owing to the location of the earthquakes, the number of P 31 1 i.2 i 1.3 1 i.5 1 i.3 wave polarity data, and the good azimuthal coverage of the

32 10.6 11.7 13.0 i 1.8

seismic stations on the focal sphere (see Table 3). 33 12.8 12.7 12.6 12.7

4. Determination of the Present-Day Stress State

Focal mechanisms have generally been used to determine present-day stresses [,4ngelier and Mechler, 1977; Angellet, 1984; Gephart and Forsyth, 1984; Gephart, 1990a, b]. The micropolar kinematic hypothesis of T•viss e! al. [1993] applied by T•,iss and Unruh [1998] points out that in a general setting, earthquake focal mechanisms only give an indication of the strain rates rather than of the stresses. These authors show that

the mathematics used to determine the stress tensor and strain

rate tensor are the same. However, stresses can be determined only in specific conditions that must include a relative vortic- ity equal to zero and isotropic rocks. They consider that fault system development may be interpreted as a result of cata- clastic flow. Blocks may have undergone rotations, and all the faults of the system had to move at the same time. In the Gra- nada Basin the timing of earthquakes shows that fault motions are heterochronous in each fault surface and do not agree well with a cataclastic flow model. However, they agree better with classical stress hypotheses in which each fault has a nearly in- dependent motion [,4ngelier and Mech/er, 1977; ,4ngelier, 1984; Gephart and Forsyth, 1984; Carey-Gai/hardis and Mercier, 1992].

The orientation of the P and T axes of the focal mecha-

nisms does not coincide with that of axes o• and o3 of the

Table 2. Estimated Focal Depth Using Wadati Diagrams Event Average Depth S.d.

Km km

34 i 8.8 18.2 17.5 18.2

35 17.6 17.7 17.8 17.7

36 ii.6 11.5 11.3 11.5

37 i 1.5 10.7 9.7 10.6

38 15.0 15.9 17.0 16.0

39 10.8 11.2 1 i.9 1 i.3

40 ii.7 10.9 10.1 10.9

41 16.9 16.3 15.7 16.3

42 12.9 13.9 15.2 14.0

43 8.8 9.3 10.3 9.5

0.8

0.0

0.2

0.4

0.0

0.1

0.2

0.2

0.1

0.6

0.2

0.2

0.1

0.1

0.1

0.1

0.1

0.0

0.1

0.1

0.4

0.4

0.3

0.3

0.8

0.1

1.0

0.1

0.5

0.1

0.1

0.7

0.8

0.4

0.7

0.5

0.9

0.6

S. d.. standard deviation.

stress ellipsoid [,4ngelier and Mechler, 1977; ,4ngelier, 1984; Gephart and Forsyth, 1984] owing to the presence of prior weakness planes and friction, nor does it coincide with the strain rate axes [T•viss and Unruh, 1998]. However, the o• axis must have an orientation inside the compressional dihe- drom, and the o3 axis must have an orientation inside the ex- tensional dihedrom (o•>cs•>cs_•, positive compression). The orientation of axes o• and o3 of the stress ellipsoid may be determined by right-dihedra diagrams [,4nge/ier and Mech/er, 1977; Ange/ier, 1984], even if slips have taken place along previous weakness surfaces. The right-dihedra diagrams are calculated by the combination of the focal mechanisms of a set of earthquakes and show the percentage of compression or

692 GALINDO-ZALDiVAR ET AL ß STRESSES IN GRANADA BASIN, BETIC CORDILLERAS

Table 3. Solution of Calculated Focal Mechanisms

Event A Nodal Plane B Nodal Plane 7'Axis P Axis

Phi Dip Rake Phi Dip Rake Trend Plunge Trend Plunge 1 320.0 45.0 -100.7 155.0 46.0 -79.5 237.6 0.5 143.8 82.4

2 67.0 90.0 90.0 247.0 0.0 90.0 337.0 45.0 157.0 45.0

3 248.0 70.0 90.0 68.0 20.0 90.0 158.0 65.0 338.0 25.0

4 12.0 45.0 -100.7 207.0 46.0 -79.5 289.0 0.5 195.8 82.4

5 174.0 90.0 90.0 354.0 0.0 90.0 84.0 45.0 264.0 45.0

6 66.0 90.0 90.0 246.0 0.0 90.0 336.0 45.0 156.0 45.0

7 163.0 65.0 - 103.7 13.0 28.3 -63.1 263.2 18.9 47.4 67.1

8 293.0 65.0 -90.0 113.0 25.0 -90.0 23.0 20.0 203.0 70.0

9 135.0 70.0 -55.0 251.0 39.7 -147.6 199.8 17.6 86.0 51.8

10 100.0 58.0 127.1 225.0 47.5 46.0 64.8 58.5 164.6 5.9

11 120.0 25.0 -94.5 305.0 65.1 -87.9 33.4 20.1 219.3 69.8

12 165.0 20.0 -75.9 330.0 70.6 -95.1 64.0 25.5 231.8 64.0

13 135.0 30.0 -57.8 279.0 65.0 -107.1 21.6 18.2 158.6 65.7

14 315.0 80.0 -90.0 135.0 10.0 -90.0 45.0 35.0 225.0 55.0

15 322.0 75.0 -90.0 142.0 15.0 -90.0 52.0 30.0 232.0 60.0

16 64.0 15.0 -39.0 192.0 80.6 -101.8 292.1 34.6 88.2 52.9

17 244.0 75.0 -90.0 64.0 15.0 -90.0 334.0 30.0 154.0 60.0

18 45.0 25.0 -90.0 225.0 65.0 -90.0 315.0 20.0 135.0 70.0

19 131.0 45.0 -64.5 277.0 50.3 - 113.3 23.3 2.8 121.9 72.0

20 123.0 65.0 0.0 33.0 90.0 155.0 345.2 17.4 80.8 17.4

21 45.0 45.0 -159.2 300.0 75.5 -46.9 359.3 18.8 251.0 42.6

22 30.0 45.0 - 151.5 279.0 70.3 -48.7 340.1 15.2 232.8 47.6

23 321.0 75.0 -90.0 141.0 15.0 -90.0 51.0 30.0 231.0 60.0

24 101.0 45.0 -100.7 296.0 46.0 -79.5 18.6 0.5 284.8 82.4

25 15.0 45.0 -113.8 227.0 49.7 -68.0 301.6 2.4 203.0 73.2

26 323.0 90.0 -90.0 143.0 0.0 -90.0 53.0 45.0 233.0 45.0

27 315.0 90.0 -90.0 135.0 0.0 -90.0 45.0 45.0 225.0 45.0

28 319.0 80.0 -90.0 139.0 10.0 -90.0 49.0 35.0 229.0 55.0

29 40.0 58.0 90.0 220.0 32.0 90.0 310.0 77.0 130.0 13.0

30 334.0 85.0 -90.0 154.0 5.0 -90.0 64.0 40.0 244.0 50.0

31 99.0 90.0 90.0 279.0 0.0 90.0 9.0 45.0 189.0 45.0

32 150.0 90.0 90.0 330.0 0.0 90.0 60.0 45.0 240.0 45.0

33 45.0 35.0 -129.3 270.0 63.7 -66.1 342.7 15.4 219.6 63.3

34 345.0 55.0 -134.8 225.0 54.5 -44.8 104.9 0.3 195.3 54.7

35 337.0 80.0 -90.0 157.0 10.0 -90.0 67.0 35.0 247.0 55.0

36 76.0 23.0 - 109.5 277.0 68.0 -82.0 0.8 23.0 200.9 65.7

37 76.0 23.0 -104.8 272.0 67.8 -83.8 357.0 22.6 193.1 66.6

38 128.0 35.0 90.0 308.0 55.0 90.0 218.0 80.0 38.0 10.0 39 171.0 57.0 90.0 351.0 33.0 90.0 81.0 78.0 261.0 12.0

40 90.0 28.0 81.2 280.0 62.4 94.7 200.9 72.2 6.5 17.2

41 296.0 80.0 -90.0 116.0 10.0 -90.0 26.0 35.0 206.0 55.0

42 75.0 45.0 -125.3 300.0 54.7 -60.0 9.0 5.3 267.7 65.3

43 90.0 20.0 -61.5 240.0 72.5 -99.8 337.8 26.9 135.4 61.3

Phi, strike in degrees for each nodal plane; dip, dip of the planes; rake, sliding angle.

extension dihedra of each orientation in stereographic projec- tion. When all the earthquakes originate in a region with the same stress ellipsoid, the orientation of o• is included in the COlnpression dihedra of all the focal mechanisms (100% P) and the orientation of o3 is included in all the extensional di- hedra (100% 7). However, when the focal mechanisms have formed under different stress ellipsoids, there are no orienta-

tions with 100% P and 100% L In this case, the orientations with the maxilnUln percentages of P and T correspond with the orientations of o• and 03 of the best represented stress el- lipsoid. In addition, Gephart and Forsyth's [1984] and Gephart's [1990a, b] method, based on verification of the suit- ability of stress states by search networks, allows the determi- nation of the shape of the stress ellipsoids. In this paper we

GAL1NDO-ZALDiVAR ET AL.: STRESSES IN GRANADA BASIN, BETIC CORDILLERAS 693

consider an axial ratio (ar=(o2-o•)/(o•-o•)) that is similar for present stresses and for palcostresses.

The focal mechanisms calculated for the southeastern Gra-

nada Depression correspond to earthquakes with depths be- tween 9.5 and 29.1 kin, although most are shallower than 20 kin, thus allowing the present-day stress state of this crustal region to be characterized. The stereographic projection of the /' axes of the focal mechanisms of earthquakes with m>2.5 (Figure 4) shows that these axes preferentially plunge SW. The 7' axes diagram (Figure 4) shows more dispersion, al- though generally the axes plunge north and NE.

The right-dihedra diagram of the focal mechanisms with ,7>2.5 (Figure 4) indicates that cy• plunges steeply SW and that cy•, more poorly defined, plunges moderately NE. Never- theless, it should be noted that no direction is compatible for all the compressional dihedra nor for the extensional ones. Therefore, although the regional stress state of the study area is extensional, some earthquakes indicate a local compres- sional setting. The right-dihedra diagram for normal fault mechanisms (Figure 4) indicates the same orientation of the stress axes as the general diagram does, and in addition, there are orientations showing compatibility for all the compression and extension dihedra of the focal mechanisms analyzed. The reverse-fault mechanisms are much scarcer, and their right- dihedra diagram shows a vertical o• axis and NW-SE or NE- SW trends (the latter being poorly defined) for o,.

The method of focal mechanism analysis proposed by Gepharl and Forsylh [1984] and Gephart [1990a, b] gives very similar results whether the rotation of the poles of striae with all the data (o•: plunge, 65/trend, 209; 03:25/28; at=0.2) or only with normal faults (o•:72/220; o3:19/40; at-0.1) are considered. These results are close to those obtained when

using the method allowing rotation of the poles of striae and planes with all the data (0,:89/220; o•:1/39; at=0.3) or only with normal faults (o•:89/220; o•:1/45; at=0.3). The stress el- lipsoids determined are prolate to triaxial, with a NE-SW ex- tensional trend and o; plunging shallowly NE. The stress state determined with plane and pole rotation in Gephart's method indicates that o• is vertical. Nonetheless, the fit using only pole rotation shows that o, is inclined SW (72 ø toward N220øE), a result compatible both with the right-dihedra dia- grams and with the preferential orientation of the P axes (Fig- ure 4).

The spatial distribution of the focal mechanisms (Figure 3a) indicates that reverse-fault mechanisms are distributed among normal-fault ones. In addition, they are also dispersed over time. Their coexistence is characteristic of the stress state in

the •study region, leading us to examine in detail the focal mechanisms determined for earthquakes with m>2 in the Agr6n series of 1988-1989 (Tables 1, 2, and 3). The reverse- fault mechanisms are distributed in space and time with nor- real-fault mechanisms, as in the entire SE of Granada Depres- sion. While the 7' axes diagram (Figure 5) shows notable dis- persion, most of the P axes plunge steeply except for three earthquakes for which P plunges only shallowly NNW, SSE, and ENE. The right-dihedra diagram Ibr all the focal mecha- nisms (Figure 4) shows that the general stresses have an ENE- WSW trend of extension. However, not all the data are in ac- cordance with this general setting. If the focal mechanisms are classified by magnitude, the diagram of mechanisms with

m>2.5 (Figure 4) has the same characteristics as the general diagram does. Notwithstanding, the diagram for earthquakes with m<2.5 (Figure 4) indicates a NW-SE subhorizontal trend of extension, compatible with all the data. If the m>2.5 in- compatible mechanisms are separated into two groups, normal and reverse faults, then compatible solutions are obtained for all the mechanisms (Figure 4). The normal-fault mechanisms indicate NE-SW or NW-SE trends of extension. and the re-

verse-fault mechanisms indicate NW-SE compression. Thus the subhorizontal NW-SE trend corresponds with an approxi- mate orientation for o: of regional stresses, with the orienta- tion of cy• deduced from the reverse mechanisms and the ori- entation of cy• deduced from the m<2.5 mechanisms.

5. Surface Geological Data and Palcostress Determination

In order to study the recent evolution of deformations and palcostresses in the Granada Basin, we present the main geo- logical features of the area. In the region SE of the Granada Basin the sedimentary rocks of the depression lie over the metamorphic rocks of the Betic Cordilleras Internal Zones. These metamorphic rocks are ascribed to two superimposed complexes. The upper one, the Alpuj/trride, is formed of vari- ous tectonic units whose lithological column comprises, from bottom to top, dark schists with some metasandstones attrib- uted to the Paleozoic; fine-grained schists, bluish, gray or violet phyllites, and slates with levels of conglomerates, quartzites and calcschists attributed to the Permian or Wer- fencan [Ga//egos, 1975], capped by Middle to Late Triassic limestones and dolostones [•..larlin, 1980]. These units have undergone high-pressure/low-temperature metamorphism [G(?.//• el a/., 1989; Azahdn el aI., 1992] that ended 25 Myr ago [.l/onid e[ a/., 1991], followed by nearly isothermal de- compression and cooling to below 350øC at 19 Ma [Moni• el a/., 1991].

The lower complex, the Nevado-Fil/tbride, has a lithologi- cal sequence similar to that of the Alpuj/trride. It begins with schist-bearing graphite, metasandstones, and orthogneisses, overlain by gray schists and metasandstones. They are fol- lowed by calcschists and marbles. In the upper part of the se- quence there are bodies of serpentinites, amphibolites, eclogi- res, and orthogneisses. The rocks have undergone Alpine metan•orphism under high-pressure/low-temperature condi- tions, evolving to a stage of intermediate pressure and tem- perature [Ptt,,,,a. 1976] at 23 Ma [,•1oni• el aI., 1991]. These rocks cooled to below 350øC at 17-16 Ma [.lloni• el aI., 1991]. These complexes are separated by a brittle extensional detachment with a S W movement of the hanging wall and as- sociated ductile, brittle-ductile, and brittle extensional struc- tures in the tbotwall [Ga/i,do-Zahtivar el a[., 1989; daba/o), el a/. 1093]. This detachment was active until the early Torto- nian [(h,•zd/ez-Lodeiro el a/.. 1996; Johnson el aI.. 1997] and was deformed b3 a NE-SW upright fold during the Pliocene [.Jo/,•xo,, 1997]. During the Quaternary, and coetaneous with the lblding, the detachment underwent partial reactivation as the basal t•ault of a high-angle fault system with NW-SE strikes dipping mainly toward the SW, with some conjugate f•aults [(hz/izzdo-Zahth'az' el a/., 1996].

694 GALINDO-ZALDiVAR ET AL.' STRESSES IN GRANADA BASIN, BETIC CORDILLERAS

N N

c N N

D

N

..... ß .... • 1.00% P

60%T

100% T

Figure 4. P and T axes projection and right-dihedra diagrams for m>2.5 earthquakes in the Granada Basin: '(a) stereographic projection of ? axes (38 data), (b) stereographic projection of T axes (38 data), (c) right-dihedra diagram of all earthquakes (38 earthquakes), (d) right-dihedra diagram including focal mechanisms with normal faults (31 •earthquakes), and (e)right- dihedra diagram including focal mechanisms.with reverse faults (7 earthquakes)

GALINDO-ZALDiVAR ET AL.' STRESSES IN GRANADA BASIN, BETIC CORDILLERAS 695

The sedimentary infill of the SE Granada Depression in- cludes middle Miocene rocks that crop out in the SE areas [Rodriguez-Fe•'ndndez, 1982], as well as late Miocene to Quaternary rocks. The Tortonian rocks are marine, and their outcrops reach up to 1500 m above sea level [Rodriguez- Ferndndez e[ a!., 1989], while the Turolian and subsequent rocks are continental. The sedimentary infill, like the meta- morphic basement, is affected by upright open folds trending NE-SW and by the above mentioned normal-fault set with NW-SE strikes. Other high-angle normal faults, affecting up to the Quaternary rocks, have E-W and NE-SW strikes (Figure 3) and are coetaneous with the activity of the NW-SE set. These data suggest a non plane strain for the recent evolution of the region and are compatible with conditions that varied fi'om pluridirectional (radial) to clear NE-SW extension.

In order to establish the characteristics of the stress ellip- soid that acted in a region, the orientation data of the brittle structures are usually analyzed. Microfault orientation pro- vides the most complete information about the deviatoric stress ellipsoid. Various methods have been proposed to study them, all assuming that the stria on the fault plane is parallel to the trend of maximum shear stress on that surface [Bott, 1959]. In this paper we have used the Search Grid method [Galindo-Zaldh'ar and Gonzdlez-Lodeiro, 1988], which cal- culates the axial ratio and the orientation of the axes of the

stress ellipsoids that have acted during superimposed fractur- ing stages. This method may be applied taking into account several restrictions in the analyzed outcrops [Galindo- Zaldi•'ar and Gonzdlez-Lodeiro, 1988] to obtain a consistent solution: homogeneous stress state, isotropic material, and no internal block rotation. These conditions are similar to those

considered by other authors [Carey and Brunier, 1974; geller and .•techler, 1977; Angelier, 1984; Etchecopar eta!., 1981; A!eksandro•'ski, 1985]. If there is overprinting of dif- ferent deformational stages, each one should be represented by a single stress state. In an outcrop the necessary field observa- tions to try to fulfill these conditions are the following: meas- urements taken in a small area, flat fault surfaces with straight striae, small slips, and uniform !ithology. We have also meas- ured the orientation of conjugate faults to obtain complemen- tary data on the orientation of the stress ellipsoid axes.

We have examined 11 measurement sites in the study area (Figure 3b and Table 4), 8 of which are located in metamor- phic basement rocks and 3 of which are located in the sedi- mentary rocks [see also Galindo-Zaldivar et al., 1989, 1993; •labaloy eta!., 1992]. In the metamorphic basement they are located in the Middle-Late Triassic A!pujfirride limestones and dolomites and in the Nevado-Fi!fibride serpentinites. Keeping in mind that the Alpujfirride and Nevado-Fi!fibride rocks cooled to below 350øC at 19 and 16 Ma, respectively, and that the brittle-ductile transition in calcite-bearing rocks takes place between 350 ø and 250øC [Suppe, 1985], the age of the brittle deformation should be younger than this cooling age [:•1oni• eta!., 1991]. In the sedimentary rocks, microfaults are located in marls and sandstones with gypsum froin the middle Miocene and late Tortonian. Nevertheless, many of the larger structures clearly deform Quaternary deposits (Figure 3).

......... The markers used to establish the microfault regime were marker displacement, slicken-sides, growth fibers, crushed tails associated with resistant clasts, Riedel fractures, and pla-

nar and crescent-shaped extensional fractures. However, al- though in some faults the striae have been observed, the sense of motion could not have been established, owing to the poor development of kinematic indicators. The Search Grid method can use faults with striae showing known or unknown senses of motion together [Galindo-Zaldh'ar and Gonzdlez-Lodeiro, 1988]. ,dle/•'xamh'o•'ski [1985] shows graphically that for a stress ellipsoid. determined by the orientation of its axes and its axial ratio. there is only one specific pattern of slip trends on variable-oriented fault planes. If we consider only the pat- tern of striae orientation, without taking into account the sense of motion. there are only two ellipsoids that fit this pattern, differentiated by the opposite senses of motion of the striae. These ellipsoids are coaxial, with an equal orientation of switched c• and o;, and complementary axial ratios (at-- x and ar- 1-x respectively). The Search Grid method looks for the best fit of the observed striae orientations with the theoretical

patterns of all the possible stress ellipsoids in a grid with a step in orientation of axes of 2 ø and 0.01 in axial ratios. The best stress ellipsoid should have theoretical striae parallel to the measured striae and the same sense of motion when it has

been determined. Then the striae with unknown senses of mo-

tion help to determine the stress ellipsoid, although it is neces- sary to know the regime of some faults to choose the stress el- lipsoids between the two possibilities. In addition, this method can distinguish between certain (S) and probable (P) stress el- lipsoids, according to the quality of the criteria used for the determinations of the fault regimes. The distributions of the angles between the orientation of the calculated and measured striae allow the ellipsoids to be classified into three groups depending on whether the fit is (1) poor, with high angles for most of the striae population, (2) medium, or (3) good, with low angles for most of the striae (Table 4). When the angle between the calculated and real stria is greater than 15 ø, the stria is considered to belong to another faulting stage and may be used to calculate a new overprinting stress ellipsoid. This value coincides with that proposed for the inverse inethods by •trmO'o eta/. [1982] and with the lnaxilnum measurelnent er- ror allowed by Angellet [1984].

The general results (Table 4) are similar to those described by Ga/indo-Zaldivar eta/. [1989, 1993] and dabaloy et al. [1992], where they distinguish the two most frequent groups of ellipsoids, although there are also intermediate ellipsoids. The first group they describe has oblate to triaxial ellipsoids with o; NE-SW subhorizontal while the second group has prolate to triaxiai ellipsoids with O l subvertical, indicating nearly radial extension. The first group of ellipsoids is dated as prior to the Tortonian as these ellipsoids have been detected in the metamorphic basement and in Serravallian sandstones (stations 9 and 10, Table 4) [•laba/oy et al., 1992]. The second group is Tortonian to the present, determined in whole rocks, including station 7 (Table 4) with upper Tortonian sandstones and station 11 (Table 4) located near a large fault (Figure 6d) that affects up to Holocene deposits. The study of new out- crops has also allowed the identification of prolate to triaxial ellipsoids with low plunging NW-SE Ol (stations 5, 9, and 10, Figure 3b and Table 4) and oblate to triaxiai ellipsoids with subvertical (station 5, Figure 3b and Table 4).

The Neogene sedimentary rocks contain outcrops with normal high-angle conjugate faults trending ESE-WNW (Fig-

696 GALINDO-ZALDiVAR ET AL.' STRESSES IN GRANADA BASIN, BETIC CORDILLERAS

A N N ß

ß ß

N N

c D

E • F •

G N

-lOO% P

•_:: ..... 60%P

60% T

100% T

GALINDO-ZALDiVAR ET AL.' STRESSES IN GRANADA BASIN, BETIC CORDILLERAS 697

Table 4. Paleostress Ellipsoids Determined From Microfaults Analysis Station Latitude, Longitude, Height Estriae Ellipsoids Axial o• o,• Fitting Rocks (Age)

øN øE Meters Number No ratio Trend Plunge Trend Plunge '1 (GS1) 37.166 -3.443 1160 37 I 0.23 I l 80 262 3 P2 Dol-ALP (Triassic) 2 (GS3) 37.193 -3.434 1100 11 I 0.44 257 64 68 26 S3 Dol-ALP (Triassic) 3 (D23) 37.084 -3.505 1440 33 I 0.03 288 76 18 0 P3 Dol-ALP (rl'riassic) 4 (O1) 36.915 -3.428 640 30 I 0.30 27 82 331 2 S2 DoI-ALP (r!'riassic) 5 (O21) 36.880 -3.421 390 21 I 0.26 304 33 99 54 S2 DoI-Ai, P (rl'riassic)

2 0.55 28 12 227 77 S3

6 (RI17) 37.013 -3.481 2150 25 I 0.97 309 62 212 4 SI Serpent-NF 2 0.34 141 68 46 2 S2

7(PINOS) 37.166 -3.509 792 12 I 0.24 90 76 274 14 S3 Sands (Up Tort) 8 (ZAFA) 36.919 -4.101 635 14 I 0.25 8 44 98 0 S2 DoI-ALP (Triassic)

2 0.83 279 6 186 23 P2

9 (QUI) 37.120 -3.468 982 26 I 0.36 161 26 272 37 S2 Sands (Serf) 2 0.59 116 18 214 24 S2

10 (QU2) 37.213 -3.481 1106 17 I 0.03 145 40 291 45 Sl Sands (Serf) 2 0.68 69 56 242 34 S3

11 (NIGU) 36.985 -3.539 942 19 I 0.09 32 78 274 11 S3 DoI-ALP (Triassic)

Dol-ALP. dolomites of Alpujarride Complex: Serpent-NF. serpentines of Nevado-Filabride Complex: Sands (Up Tort). upper Tortonian sandstones: Sands (SerO. Serravallian sandstones.

ure 6), forming a systeln of domino faults over the contact between the Alpujfirride and the Nevado-Filfibride [Ga/ino- Za/dh'ar eta/.. 1996]. The outcrops of both the major and mi- nor faults in the system have been studied by San de Ga/- deano [1983] and Ga/indo-Za/dh'ar eta/. [1996], who both report that part of the NNE-SSW extension is post-Tortonian. Although solne of the faults are capped by Pliocene rocks (Figure 6), others are currently active. The geometry of the fault system is in agreement with the paleostress ellipsoids determined.

6. Discussion

6.1. Distribution of Seismic Activity

The earthquakes studied were distributed throughout the southeastern half of the Granada Depression (Figure 3), as the faults with recent activity were. The distribution of deep earthquakes is related to the lithology, stress states in the crust, and temperature, all properties that determine the me- chanical characteristics of the crust. The seismicity indicates that the deformation is occurring in conditions of brittle fault- ing. Most of the earthquakes analyzed took place at depths of h<16 km, although they vary up to depths of h<23-25 km to- ward the S W sector of the Granada Depression [Morales et a/., 1997]. The existence of a maximum depth at which earth- quakes occur in the crust may be related to a boundary be- tween an upper brittle crust, allowing the existence of fric- tional instability that produces seismic activity, and a ductile

lowel' crust causing aseismic deformation by stable frictional sliding.

Smith and Bruhn [1984] observed in the Basin and Range- Colorado Plateau that the brittle-ductile transition in the crust

occurs at the depth above which 80% of earthquakes occur. Sibson [1982], studying the seismic-aseismic transition in various regions of the United States, reached the conclusion that the lowel' boundary of microseismic activity depends on the thermal state of the regions. Tse and Rice [1986] demon- strated that the base of the seismogenic layer could only be understood as the transition from nonstable to stable frictional

slipping due to an increase in temperature. In quartz-feldspar continental crustal environments this seismogenic layer is re- lated to the 300øC isotherm [Scho/z, 1990]. Therefore, lacking heat flow data for this zone and considering a quartz-feldspar upper crust, the depth distribution for the seismic activity sug- gests that the 300øC isotherm must be between 14 and 16 km. These data point to a low geothermal gradient, of around 19 ø- 20øC/km, for the region.

6.2. Present-Day Stress and Permutation of Stress Axes

The analysis of mechanisms of m>2.5 from the SE part of the Granada Depression, as well as the m>2.1 mechanisms from the Agr6n series, indicates the existence at present of a heterogeneous stress state that is not compatible with all the mechanisms. The regional stress state has a NE-SW exten- sional trend, with o• plunging slightly NE and o• plunging steeply S W. The stress state is only compatible for all the

Figure 5 (opposite). P and T axes projection and right-dihedra diagrams for the Agr6n earthquake series: (a) stereographic projection of P axes (16 data), (b) stereographic projection of T axes (16 data), (c) right-dihedra diagram of all earthquakes (16 earthquakes), (d) right-dihedra diagram including all the focal mechanisms of m> 2.5 earthquakes (11 earthquakes), (e) right-dihedra diagram including focal mechanisms with normal faults and m> 2.5 earthquakes (8 earthquakes), (f) right-dihedra diagram including focal mechanisms with reverse faults and rn> 2.5 earthquakes (3 earthquakes), and (g) right-dihedra diagram including all the focal mechanisms of earthquakes with m< 2.5 ( 5 earthquakes).

698 GALINDO-ZALDiVAR ET AL.: STRESSES IN GRANADA BASIN, BETIC CORDILLERAS

Figure 6. Field examples of Recent deformations in the region. Location is in Figure 3b. (a) Overview of P!iocene sediments in the hanging wall of the Rio Dilar Fault. Note wedge-shaped layers and fan-like array. (b) Cross section of normal conjugate faults deforming Tortonian sediments and fossilized by Piiocene sediments, Barranco de !as Angustias. Split pebbles in the Tortonian rocks, Barranco de las Angustias. (d) Overview of Nig•ielas Fault. Note triangular facets and deformed Quaternary alluvial fans. (e) Detailed view of the surface of Nig[ielas Fault. (f) Conjugate faults deforming turbiditic sandstones and marls of late Tortonian-Messinian age.

GALINDO-ZALDiVAR ET AL.' STRESSES IN GRANADA BASIN, BETIC CORDILLERAS 699

mechai•isms if we consider groups classified by magnitude or type of faults, which indicates the existence of variations in the stresses with respect to the regional stress state. These data suggest the presence of local permutations in the stress axes, varying in space and time, with the main stress ellipsoid axes remaining approximately subvertical and subhorizontal NW- SE and NE-SW. The most common permutations detected in the Agr6n series are of two types. The first type consists of a reversal between o: and 03. Thus, the general stress setting, with NE-SW extensional conditions, changes to NW-SE di- rected extension. represented by earthquake mechanisms with m<2.5. 'In the second type the three axes flip when the re- gional stress changes to a NW-SE compressional setting. Al- though other permutations are also possible, the first type is particularly viable owing to the prolate nature of the general stress ellipsoid. which allows the exchange between axes o: and o• because their values are nearly identical. In addition, the main tectonic deformations in the region, with NE-SW normal faults and compressional structures (mainly folds), in- dicating NW-SE shortening, favor the second type of permu- tation.

The permutation of present-day stresses has also been de- tected in the Rif (Figure 1) by Medina [1995], on basis of the study of focal mechanisms. lYesnouky and Jones [1994] and Bellier and Zobac]c [1995], in the Basin and Range, and Cc,'e)'-Gailhc,'dis and ,tlercie•' [1992], in the Rhine graben, have also determined a rapid variation in the stress state over short time spans. Whereas Wesnouky and Jones try to model a fault system with incompatible stresses but a compatible dis- placement field, Belllet and Zoback study the same area and show that some faults have two striae indicating reactivations. Therefore some outcrops have been deformed under different stress states in very short periods of time. Carey-Gailhardis and Mercier used earthquake focal mechanisms to determine that a single stress state cannot model all the data of the Rhine graben. They found a main stress state and other stress states that are coaxial with the main one and propose that this is the result of the incompatibility of block motions in an assem- blage of rigid or elastic blocks and discard the permutations of the principal stress axes of the regional stress field. Although this mechanism may be active in our study region, we propose another mechanism for the Granada Basin that may produce main and secondary stress states in the area. The study of the development of orthogonal extension joint systems [Hancock, 1985] suggests that the permutation of axes c•2 and c•3 is a con- sequence of the evolution of brittle deformations. Joint devel- opment allows the relaxation of stresses perpendicular to the joint, which produces the change in the axial ratio of the stress ellipsoid, even switching the stress axes, but without varying the axial orientation. Likewise, fault activity can relax the stress acting on that fault [Meissner and Slrehlau, 1982] and can cause the axial ratio of the stress ellipsoid to vary. There- fore we propose that in spite of the fact that there are regional stresses, activity in the fault systems can cause relaxation of the stresses. This relaxation is mainly evident as a local per- mutation of the main stresses, although the preferential orien- tations of the axes are maintained. This fall in the stress value

tends to be greatest at depth, where the maximum number of earthquakes are located [Meissner and Slrehlau, 1982]. This setting may be similar to that determined for the earthquakes

of the Granada Basin, which clearly show a change in the po- sition of the main stress axes.

6.3. Relationship of Surface Structures, Paleostresses, and Seismicity

Surface data indicate that the southeastern part of the Gra- nada Basin has been undergoing intensive extensional tecton- ics since the Miocene, resulting in the rapid uplifting of high- pressure / low-temperature metamorphic rocks [Moni• el a!., 1991; .Johnson e! a!., 1997]. This extension is accommodated by low-angle normal faults and shear zones, mainly active during the early Miocene, that indicate SW-NE extension [Gali•do-Z•ddi•'ar et al.. 1993]. Among the ductile structures there are folds showing coetaneous NW-SE shortening [Gal- indo-Zaldh¾,' el al.. 1989]. All these structures were fossilized in the early Tortonian [Gonzc•lez-Lodeiro et al., 1996; Johnson ez al.. 1997] and are associated with oblate stress ellipsoids with a subhorizontal SW-NE trending c•3 [Galindo-Zaldivar el al.. 1993].

The extensional deformations continue at the present with the activity of normal faults showing SW-NE extension [Gal- indo-Zcddh'ar el a!., 1996; Johnson et al., 1997]. The NW-SE trend has acted as a compressional direction, originating E-W to NE-SW folds in the upper Tortonian [Eslevez eta!., 1982] and Pliocene [Johnson, 1997]. Remarkably, it also acted sometimes as an extensional trend, originating normal faults and resulting in stress states characterized by radial extension [Galindo-Zaldi•'w' el a!., 1993], corresponding to the prolate to triaxial palcostress ellipsoids with subvertical C•l. In addition to these ellipsoids, there are others indicating NW-SE com- pression that may be related to the development of the folds.

The focal mechanisms determined are not associated with

specific faults cropping out on surface. However, the present- day stresses that can be deduced from them have the same general features as the palcostresses since the late Miocene, the period in which the cordilleras began to rise with respect to the Alborfin Sea and in which the basins started to form.

These deformations are characterized by the permutation of stress axes favored by the regional tectonic context. The coe- val existence of different stresses in nearby areas and the switch in the same region from NE-SW extension with folds indicating perpendicular compression could be explained by a horizontal deformation ellipse whose .V (extensional) axis trended NE-SW and whose }: (shortening) axis was NW-SE. This ellipse is similar to that proposed for other regions of the eastern Betic Cordilleras [i.e., Montenat el a/., 1987]. None- theless, it should be noted that while strike-slip faults are common in the eastern Betic Cordilleras, they are quite scarce in the Granada Basin. The relaxation of the iaaain stresses

might give rise to radial extension situations that would ex- plain the variably trending or even the NW-SE trending nor- mal faults. Analogously to the process of stress and fracturing relaxation during the formation of joint systems [Hancock, 1985], this process can be repeated several times, such that NW-SE compression and SW-NE extension can alternate with periods of radial extension, approximately maintaining the orientation of the iaaain axes of the stress ellipsoid. Further- more. if the conditions of stress relaxation are appropriate, even NW-SE extensional events can be found. The depend- ence of this relaxation process on fault activity would explain

700 GALINDO-ZALDiVAR ET AL.' STRESSES IN GRANADA BASIN, BET[C CORDILLERAS

the alternating stress states at different times in distinct parts of the same region.

The seismic activity may be related with strongly dipping active faults, similar to the surface fault system, and with slightly dipping faults, such as the detachment between the Nevado-Fil•_bride and Alpuj•_rride. Although this detachment is the deepest one currently cropping out in the Betic Cordille- ras and has recently been reactivated at the western margin of Sierra Nevada [Ga/indo-Za/dh,ar e! a/., 1996], it has no asso- ciated large seismic activity, probably owing to its shallow position. The presence of faults of this type may be favored by variations in the mechanical behavior of the crust on both

sides of the 300øC isotherm. The existence of crustal detach-

men! levels at depths fi'om 10 to 20 km is suggested by the •leep reflection seismic profiles that cut across the Internal Zones, 30 km east of the study region [Ga/indo-Za/d/var e! a/., 1997]. the seismic refraction profiles that intersect the seismic reflection profiles and pass through the study area [Banda el a/., 1993], and the magnetoteiluric data [Potts e! a/., 1995].

The current crustal thickness in the Internal Zones of the

Betic Cordilleras is 35-38 kin, thinning gradually to 33-35 kin toward the Iberian Massif and abruptly to 15 kin toward the Alborfin Sea [Galindo-Zaldh•ar el al., 1997]. Marine regres- sion and uplifting of the cordilleras have taken place since the Tortonian, probably as a response to compressional deforma- tions (large folds in the Internal Zones and thrusting of thein over the Iberian Massif) [Walls et ah, 1993; Galindo-Zaldivar el al., 1997; Johnson el al., 1997]. However, the deformations observed for that period in the upper part of the crust are

mainly extensional, suggesting that the structures responsible for the crustal thickening must be found mainly in the lower part of the crust, as suggested in the aforementioned works. The increased ductility at these depths inhibits seismicity, thus making it impossible to establish what types of active struc- tures exist in the lower part of the crust. Alternative models for the Betic Cordilleras combining uplift and extension in the upper crust may invoke thermal crustal doming or an isostati- cally uncompensated orogen. However, the moderate to low thermal gradients deduced from the distribution of the seis- micity in the region do not support the existence of a thermal dome in the area. In addition, field research [Galindo-Zaldiva/' el a/., 1989; P/at! and Fissers, 1989; Gonzd/ez-Lodeiro el a/., 1996] shows that the trends of extension are parallel to the cordillera, in contrast with those expected in an isostatic uplift of!he orogen, as in this case, the trends of extension may be directed outward.

All the above data suggest a crustal model for the Granada Basin (Figure 7)characterized by uplift of!he region since the Tortonian as a consequence of crustal thickening. The defor- mations related to crustal thickening may be folds in the upper part of the crust and other probable NW-SE compressional structures in the lower crust. In the upper crust the most com- mon deformations are high-angle normal faults and detach- ments subject mainly to NE-SW extensional stress, which lo- cally switch to NW-SE compressional stress, favored by the regional tectonic setting. Other permutations can also be de- tected, with NW-SE extension and radial extension. Present- day regional stresses seem to be characterized by a O l axis that plunges fi'om vertical to steeply SW at depth.

•W Asymmetric basin development

-Okm .... • ••-'""---"'• • / • I'

:,o / / .... ,,,,,,.,,,/:,,.,,,,.,,,,,,,::

b;5:•"•:•:•::•::•::*:.•:.•::::::•:•:;:•::':':'•' :;" •:•.5.•:•:;•::•:.*•:::• ::•.:•::•::•::•::•::•::•::::•:: • bond of eorthquokes • ..•:::::...:.:::.:.:.. :...:::•:•.:.:.:.. '....:.:.:...:: :":.:U•I•T:.:::::::::::::::::::::::::: ::::•:: of Iorgest mognitude 0 5km • // .... , , , , Detochment • •Trojectories

Figure 7. Simplified and schematic model of the upper part of the crust of Granada Basin. Local stress field includes permutation of stress axes.

GALINDO-ZALD[VAR ET AL ß STRESSES IN GRANADA BASIN, BETIC CORDILLERAS 701

7. Conclusions

The N\•'-SE convergent boundary between the Eurasian and Afi'ican plates is diffuse in the western Mediterranean and is associated v, ith a seismicity zone more than 300 km wide embracing the Betic-Rif Cordilleras and the Alborfin Sea. Anal\ sis of the focal mechanisms in the southeastern half of

the Granada Basin. together with the outcropping structures, has provided inforlnation on the lnain features of the defor- mation and the stress state during the development of this ex- tensional basin linked to a convergent plate boundary.

The main structures active in the Granada Basin and sur-

rounding areas fi'om the late Miocene to the present. which is the period comprising the uplift of the cordillera, are the fol- 1oxxino' lame NNE-SW to E-W folds and normal-fault sys- tems. Although there are normal faults with highly variable strikes. NX•'-SE high-angle faults are more frequent. In addi- tion. on the v, estern boundary of Sierra Nevada the contact bet,aeen the Nevado-Filfibride and the Alpuj•rride is a low- angle normal fault that has recently been reactivated with a top-to-the-S\V movement.

Seislnicity in the region is concentrated in the upper crust' up to a depth of 14-16 kin in the southeastern Granada Basin and reaching up to 25 kin toward the S W. The base of the seismogenetic zone possibly corresponds to the 300øC iso- therm. suggesting a low geothermal gradient for the region of about 20øC/kin. The seismicity has the same distribution as the active faults do. although there is no clear association be- tween hypocenters and surface structures. Most of the activity is likely produced by high-angle and detachment faults, with characteristics similar to those of the surface fault system.

Present-day stress deduced from focal mechanisms is very similar to the palcostresses determined from the late Miocene to present-day microfaults. The stress ellipsoids in the upper part of the crust indicate a heterogeneous stress field. The

most common ellipsoids, indicating the regional stress field, are triaxial to prolate, with NE-SW subhorizontal extension. Axis o• is vertical on the surface, plunging SW in depth, while axis o; seems to plunge shallowly NE. However, although the orientations of the lnain axes of the stress ellipsoid tend to be constant. there are secondary stress fields, corresponding to permutations of the main axes that vary in position over time. The most common change is the switch of ellipsoids with NW-SE subhorizontal compression, likely favored by the re- gional geological setting. NW-SE extensional permutations are also fi'equent, as well as radial extension due to the low axial ratio of the regional stress ellipsoids.

The Granada Basin crust therefore seems to have been un-

dergoing uplift since the Tortonian, with marine sediments of this age up to 1500 m above sea level, causing extensional structures in the upper part of the crust. This may be a conse- quence of a process of crustal thickening, with COlnpressional structures located preferentially in the lowest nonseismic part of the crust. This setting, active up to the present-day, ac- counts for the abundance of NE-SW extensional structures in

relation to NW-SE compressional structures and allows the permutation between the two types of stresses. Nevertheless, there are no large strike-slip faults, as occurs in the eastern re- gion of the cordillera. The study of focal mechanisms and de- formations in the southeastern part of the Granada Basin may contribute to the characterization of the stresses and structures

that develop in extensional areas on convergent plate bounda- ries.

Acknov• ledgments. We are gratef•l to Peter Coney, recently de- ceased. George Zandt and Susan Beck fbr their constructive criticism. We also xxish to thank John S. Oldo•v and Dogan Seber fbr their de- tailed i'eviexxs that helped to improve the paper. C. Laurin has re- x iexxed the English st3 le of the paper. This work has been supported b5 the research group 4057 of the Junta de Andalucia and by the CICY•I projects AMB97-1113-C02-01 and PB96-1452-C03-01.

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