18
3 Miocene basins in the Betic fold and thrust zone 35 Introduction The Betic Cordillera of southern Spain together with the Rif in Morocco and the Tell Mountains in northern Algeria form a morphologically distinct, arc- shaped orogen, which constitutes the western end of the Mediterranean Alpine chain (Fig. 3.1). The outer arc of this orogen consists of a fold and thrust belt (the Betic and Rif External Zones), whilst the inner arc is made up of an allochtonous pile of mostly metamor- phic rocks exposed in the Betic and Rif Internal Zones. Part of the Internal Zone is presently sub- merged in the Alboran Sea. The Betic-Rif arc geom- etry developed during the latest Oligocene and early to middle Miocene, when a combination of westward motion plus extensional deformation of the Internal Zone and slow but continuous African – European plate convergence resulted in outward thrusting and folding of the Mesozoic – Cenozoic cover in front of the migrating Internal Zone (Fig. 3.2). A number of hypotheses have been proposed in the literature to ex- plain the extensional collapse and westward drift of the Internal Zone, and the simultaneous folding and thrusting in the External Zones, in terms of orogenic processes at the African – European plate boundary (e.g., Platt and Vissers, 1989; García Dueñas et al., 1992; Royden, 1993; Seber et al., 1996; Vissers et al., 1995; Lonergan and White, 1997; Spakman and Wortel, 2004). During the Miocene, basins developed both on top of the growing fold and thrust belt of the External Zone and within the extending Internal Zone. This chapter focuses on Miocene basins in the fold and thrust belt, in particular on those in the eastern part of the Betic External Zone (Fig. 3.1). During the Neogene, the External Zone (Pre- and Subbetics) formed in essence a foreland domain, in which the Miocene is represented by a marine “Flysch” series (or “Tap” facies, or Moratalla Formation; e.g., Hermes, 1978; García-Hernández et al., 1980; Ott d’Estevou et al., 1988; Sanz de Galdeano and Vera, 1992). The flysch deposits pass upwards into late Miocene molasse type continental deposits (e.g., Dabrio, 1972; Calvo et al., 1978; Ott d’Estevou et al., 1988; Sanz de Galdeano and Vera, 1992). Both the Miocene and underlying Mesozoic – Cenozoic rocks are folded along approximately ENE to NE trending axes or are cut by generally northwest directed thrusts. Some of the Miocene basins, e.g., the Pontones basin (Figs. 3.2c and 3.3), have (partially) been overridden by thrust units and have been discon- nected from the surrounding basins. But within the Betic External Zone, and in the Prebetics in particular, there are also clearly extensional structures, which run parallel to the general ENE-WSW trend of the compressional structures as indicated on the geologi- cal maps of the Instituto Geologico y Minero de España (IGME). Some of the Miocene basins in the Prebetics are bounded by these extensional structures, such as for example the Santiago de la Espada basin (Figs. 3.2c and 3.3) which is bounded along its north- ern margin by a morphologically distinct normal fault. At first inspection, the geometry of this basin suggests an extension-related origin with significant displacements along the extensional faults, and a di- rect relationship between the compressional struc- tures (the folds and thrusts) and the extensional faults is not immediately obvious. Except of a study by Luján et al. (2000) and by Crespo-Blanc and Campos (2001), both in the Gibraltar fold and thrust belt, ear- lier structural studies of the External Betics (e.g., De Smet, 1984; Ott d’Estevou et al., 1988; Banks and Warburton, 1991; Platt et al.; 2003) seem to have ei- ther overlooked or ignored the existence and signifi- cance of these extensional structures. At least four different explanations may account for the development of extensional structures such as those in the Santiago de la Espada basin: (1) collapse of an overthickened fold and thrust belt or orogenic wedge (e.g., Davis et al., 1983; Platt, 1986; Dahlen, 1990), (2) displacement of the locus of extension from the Internal towards the External Zone (e.g., Crespo- Blanc and Campos, 2001), (3) stratal extension in the footwalls of thrust faults (Platt and Leggett, 1986), (4) hanging-wall collapse in response to a specific fault plane geometry (a shallowing-upward fault; Coward, 1983). The first two explanations imply that the extensional structures should accommodate regional,

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Page 1: Miocene basins in the Betic fold and thrust zone · PDF fileMiocene basins in the Betic fold and thrust zone 35 ... Main characteristics of the Betic External Zone ... complex rift

3Miocene basins in the Betic fold and thrust zone

35

Introduction

The Betic Cordillera of southern Spain togetherwith the Rif in Morocco and the Tell Mountains innorthern Algeria form a morphologically distinct, arc-shaped orogen, which constitutes the western end ofthe Mediterranean Alpine chain (Fig. 3.1). The outerarc of this orogen consists of a fold and thrust belt (theBetic and Rif External Zones), whilst the inner arc ismade up of an allochtonous pile of mostly metamor-phic rocks exposed in the Betic and Rif InternalZones. Part of the Internal Zone is presently sub-merged in the Alboran Sea. The Betic-Rif arc geom-etry developed during the latest Oligocene and earlyto middle Miocene, when a combination of westwardmotion plus extensional deformation of the InternalZone and slow but continuous African – Europeanplate convergence resulted in outward thrusting andfolding of the Mesozoic – Cenozoic cover in front ofthe migrating Internal Zone (Fig. 3.2). A number ofhypotheses have been proposed in the literature to ex-plain the extensional collapse and westward drift ofthe Internal Zone, and the simultaneous folding andthrusting in the External Zones, in terms of orogenicprocesses at the African – European plate boundary(e.g., Platt and Vissers, 1989; García Dueñas et al.,1992; Royden, 1993; Seber et al., 1996; Vissers et al.,1995; Lonergan and White, 1997; Spakman andWortel, 2004).

During the Miocene, basins developed both on topof the growing fold and thrust belt of the ExternalZone and within the extending Internal Zone. Thischapter focuses on Miocene basins in the fold andthrust belt, in particular on those in the eastern part ofthe Betic External Zone (Fig. 3.1). During theNeogene, the External Zone (Pre- and Subbetics)formed in essence a foreland domain, in which theMiocene is represented by a marine “Flysch” series(or “Tap” facies, or Moratalla Formation; e.g.,Hermes, 1978; García-Hernández et al., 1980; Ottd’Estevou et al., 1988; Sanz de Galdeano and Vera,1992). The flysch deposits pass upwards into lateMiocene molasse type continental deposits (e.g.,Dabrio, 1972; Calvo et al., 1978; Ott d’Estevou et al.,

1988; Sanz de Galdeano and Vera, 1992). Both theMiocene and underlying Mesozoic – Cenozoic rocksare folded along approximately ENE to NE trendingaxes or are cut by generally northwest directedthrusts. Some of the Miocene basins, e.g., thePontones basin (Figs. 3.2c and 3.3), have (partially)been overridden by thrust units and have been discon-nected from the surrounding basins. But within theBetic External Zone, and in the Prebetics in particular,there are also clearly extensional structures, whichrun parallel to the general ENE-WSW trend of thecompressional structures as indicated on the geologi-cal maps of the Instituto Geologico y Minero deEspaña (IGME). Some of the Miocene basins in thePrebetics are bounded by these extensional structures,such as for example the Santiago de la Espada basin(Figs. 3.2c and 3.3) which is bounded along its north-ern margin by a morphologically distinct normal fault.At first inspection, the geometry of this basin suggestsan extension-related origin with significantdisplacements along the extensional faults, and a di-rect relationship between the compressional struc-tures (the folds and thrusts) and the extensional faultsis not immediately obvious. Except of a study byLuján et al. (2000) and by Crespo-Blanc and Campos(2001), both in the Gibraltar fold and thrust belt, ear-lier structural studies of the External Betics (e.g., DeSmet, 1984; Ott d’Estevou et al., 1988; Banks andWarburton, 1991; Platt et al.; 2003) seem to have ei-ther overlooked or ignored the existence and signifi-cance of these extensional structures.

At least four different explanations may accountfor the development of extensional structures such asthose in the Santiago de la Espada basin: (1) collapseof an overthickened fold and thrust belt or orogenicwedge (e.g., Davis et al., 1983; Platt, 1986; Dahlen,1990), (2) displacement of the locus of extension fromthe Internal towards the External Zone (e.g., Crespo-Blanc and Campos, 2001), (3) stratal extension in thefootwalls of thrust faults (Platt and Leggett, 1986), (4)hanging-wall collapse in response to a specific faultplane geometry (a shallowing-upward fault; Coward,1983). The first two explanations imply that theextensional structures should accommodate regional,

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Chapter 3

36

Figure 3.1. Geological map of part of the External Zone of the Betic Cordillera in southern Spain, modified from MapaGeologico de la Peninsula Iberica (IGME, 1:1.000.000, 1981). Structural data in the Prebetics after Platt et al. (2003).Rectangle marks area of interest in western Prebetic Miocene basins (Pontones and Santiago de la Espada basins). IEZB- Internal-External Zone Boundary. (1) External Betics profile from Banks and Warburton (1991); (2) Carzola and Nerpio-Fuensante sections from Platt et al., (2003); (3) Pontones-Santiago de la Espada section (this study); for sections, see Figure 3.2.

Murcia

External Zone

Internal Zone

PrebeticsMesozoic-Paleogene

SubbeticsMesozoic-Paleogene

NeogeneVolcanism

Miocene

Pliocene

Quaternary

Metamorphicunits

38º

3º 2º

Pal

om

ares

fault

Baza

Cieza

Jumilla

Lorca

Alh

ama

deM

urci

aFa

ult

Sierrade

Segur

a

Crevillente Fault

Hellin

Yetas

Calasparra

Iberian Meseta

MediterraneanSea

Triassic

Iberian Meseta:Variscan basementand Mesozoic cover

IEZB

Santiago dela Espada

(2)

(3)

(1)

P. de DonFadrique

(2)

Slipvector

Thrust

Syncline

Anticline

Moratalla

Villacarrillo

- - - ? - - - ? - - - ? - - - ? - - - ? - - -

-8° -6° -4° -2° 0°

34°

36°

38°

Iberia

Betic Cordillera

Rif

TellAlboran Sea

Atlantic

Ocean

Alboran DomainExternal Zone

40°

Africa

BalearicIslandsAlicante

Cadiz Malaga

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Miocene basins in the Betic fold and thrust belt

37

a)

Banks

and

Warb

urt

on

(1991)

b)

Pla

ttetal.

(2003)

c)

This

stu

dy

10

km

0

10

km

0

km 2 0

km 2 0

(Inte

rnalB

etic

Basem

ent)

SE

NW

10

km

0

SA

NT

IAG

Ode

laE

SP

AD

AS

EC

TIO

N

km 2 0

Decolle

ment

Iberian

Basem

ent

JT

r

K

T

Ponto

nes

B.

S.E

spada

B.

Figu

re 3

.2. B

alan

ced

cros

s-se

ctio

ns o

f the

Bet

ic E

xter

nal Z

one.

For

loca

tions

of s

ectio

ns se

e Fi

gure

1. a

) Cro

ss-s

ectio

n fro

m B

anks

and

War

burt

on (1

991)

. b) C

ross

-sec

tions

from

Pla

tt et

al.

(200

3). c

) Pon

tone

s-Sa

ntia

go d

e la

Esp

ada

sect

ion

cons

truc

ted

in th

is st

udy;

for r

esto

red

sect

ion

see

Figu

re 3

.9. A

ll cr

oss-

sect

ions

show

n at

sam

e sc

ale;

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J, K

, and

T d

enot

e Tr

iass

ic, J

uras

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and

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ary,

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ectiv

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at s

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phic

thic

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s of

the

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ozoi

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its in

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ses

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ards

the

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h, a

nd th

atsh

orte

ning

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ccom

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ated

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a ge

ntly

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ppin

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t with

in T

rias

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evap

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eria

n Pa

leoz

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men

t. To

tal a

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rten

ing

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eth

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rly

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imat

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t 200

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6 km

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ks &

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tt et

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3), o

f whi

ch 4

0 or

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raz /

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anks

&W

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latt

et a

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003)

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ntia

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ada

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e la

te M

ioce

ne.

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Chapter 3

38

presumably crustal-scale extension of the Pre- andSubbetic domain. The latter two explanations implythat the extensional structures have developed due tolocal extension within a compressive setting in theBetic fold and thrust belt, and that they are inherent tothe development of the thrust structure. The geometryof the extensional structures in the Santiago de laEspada basin indeed suggest that they formed in re-sponse to late Miocene thrusting, hence that theMiocene basins in essence developed in acompressive tectonic setting.

Main characteristics of the Betic External Zone

Primarily on the basis of its Mesozoic stratigraphyand facies, the Betic External Zone is divided into twogeological domains, namely the Prebetic zone and theSubbetic zone (Hermes, 1978; García-Hernández etal., 1980). The Prebetic zone, exposed in the easternpart of the Betic External Zone, mainly consists ofMesozoic up to Miocene continental to shallow ma-rine deposits originating from the former southernIberian margin. The rocks of the Subbetic zone, on theother hand, constitute a thicker and more completeMesozoic sedimentary record with facies commonlyindicating deeper marine conditions, most likely moredistal with respect to the Iberian Meseta and south ofthe Prebetics. Relic Mesozoic extensional structuresand lateral thickness variations of the Mesozoic coversuggest that during the Mesozoic both the Prebeticsand Subbetics have intermittently been subjected toextension in response to the opening of the AtlanticOcean in the west and opening of the Piemonte-Ligurian (or Alp-Tethys) Ocean in the east, eventuallyresulting in the separation of Iberia from Africa(Banks and Warburton, 1991, Reicherter et al., 1994;Hanne et al., 2003). These events led to the develop-ment of the various palaeogeographic realms andcomplex rift and wrench basin configurations identi-fied on the thinned Iberian continental crust (García-Hernández et al., 1980; De Ruig et al., 1987;Reicherter et al., 1994; Hanne et al., 2003). In the sub-sequent post-rift stage, i.e., in the late Cretaceous andPaleogene to earliest Miocene, sediments were depos-ited in continental and shallow marine environmentsin the Prebetics and in a deep marine setting in theSubbetics (Hermes, 1978; Banks and Warburton,1991). All of these deposits are overlain by a trans-gressive marine series of early to middle Mioceneage, referred to as the Flysch deposits, the “Tap”facies, or the Moratella Formation, and lie

unconformably on earlier Miocene or pre-Miocenerocks (e.g., Hermes, 1978; García-Hernández et al.,1980; Sanz de Galdeano and Vera, 1992; Montenat etal., 1996). The Miocene marine series pass upwardsinto an upper Miocene continental series (e.g., Calvoet al., 1978; Montenat et al., 1996).

The Betic fold and thrust belt is made up of NE toENE trending folds, NW directed thrusts and SE di-rected back-thrusts (e.g., De Ruig, 1987; Ottd’Estevou et al., 1988; Banks and Warburton, 1991;Frizon de Lamotte et al., 1991; Van der Straaten,1993; Lonergan et al., 1994; Platt et al., 2003; see Fig.3.1). Shortening was accommodated along a gentlysouthward dipping detachment in the Triassic horizonat the base of the External Zone stratigraphy, whichprobably continues underneath the Internal Zone(Banks and Warburton, 1991; Fig. 3.2). From thecross-sections (Fig. 3.2) it is evident that the thrustsform part of a complex structure, which is a result ofpiggy-back, overstep, out-of-sequence and break-back thrusting (e.g., Sabat et al., 1988; Banks andWarburton, 1991; structural terms cf. Butler, 1982,1987). On the other hand, De Smet (1984) has pro-posed that the Subbetic zone is part of a major flowerstructure associated with the motion on a crustal-scalestrike-slip fault (the Crevillente fault) in the Betic(and Iberian?) basement, but structural evidence tosupport this interpretation seems to be lacking (Banksand Warburton, 1991; Platt et al., 2003). Along theentire length of the Betic External Zone including theBalearic Islands, thrusting and folding initiated in thelatest Oligocene – early Miocene, and has continuedtill the present (e.g., Sabat et al., 1988; Geel et al.,1992; Lonergan et al., 1994; Geel, 1996; Crespo-Blanc and Campos, 2001). Beets and De Ruig (1992)have recognized unconformities in the Miocenestratigraphy, which they interpret in terms of thenorthward migration of a peripheral or fore-bulge andforeland basin. Several of these unconformities andassociated stages of folding and thrusting have beenidentified: the stages have been dated as latestOligocene – Aquitanian, Aquitanian – Burdigalian,Burdigalian - Langhian and Serravallian – earlyTortonian (e.g., Calvo et al., 1978; De Ruig et al.,1987; Beets and De Ruig, 1992; Montenat et al.,1996). These data are in agreement with data from thegeological maps (IGME) and recent studies (e.g.,Calvo et al., 1978; De Ruig et al., 1987; Ott d’Estevouet al., 1988; Lonergan et al., 1994; Geel, 1996; Platt etal., 2003), showing that compressional deformationinitiated in the early Miocene in the most internalparts of the External zone (the Internal-External Zone

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Miocene basins in the Betic fold and thrust belt

39

Boundary, Lonergan et al., 1994; and Fig. 3.1), whilstit became progressively younger (up to Tortonian –Messinian) near the outer parts of the External Zone.Thrusting and folding in the most external parts of thePrebetics eventually led to uplift and closure, in thelate Miocene, of the northern Atlantic – Mediterra-nean corridor (also known as the “North Betic Strait”;e.g., Calvo et al., 1978; Sanz de Galdeano and Vera,1992; Soria et al., 1999; Martín et al., 2001; Braga etal., 2003; Sanz de Galdeano and Alfaro, 2004). Itshould be emphasized that most of the Miocene basinsin the Prebetics at present occur at an altitude of sev-eral hundreds up to 1500 meters above sea level (seealso Sanz de Galdeano and Alfaro, 2004).

The basins chosen in the context of this study, i.e.the Pontones and Espada basins (Figs. 3.1 and 3.3), liein the western part of the Prebetics. The Pontones ba-sin, the Santiago de la Espada basin, and the basin inbetween, here referred to as the Almorchón basin,contain an almost complete record from the early tothe late Miocene (Dabrio, 1970; Dabrio et al., 1971;Dabrio, 1972). These Miocene sediments are consid-ered part of the Santiago de la Espada Formation andlie in general unconformably on marine Cretaceous orshallow marine Paleocene-Eocene deposits (Dabrio,op. cit.). The total stratigraphic thickness of theMiocene sediments appears to increase from at least140 meters in the Pontones basin up to about 600 me-ters in the Almorchón and Santiago de la Espada ba-sins (Fig. 3.3). Currently, these basins are narrow,elongate structures separated from each other. TheMiocene basin fill is folded and disrupted by bothextensional and thrust faults. In case of the Pontonesbasin, it is disrupted and overridden from the south-east by a number of discontinuous thin-skinnedthrusts, which has resulted in an imbricate structure ofrepeatedly stacked uppermost Cretaceous andMiocene rocks (Fig. 3.2). Both the Santiago de laEspada basin and the Almorchón basin are overriddenfrom the southeast by thrust units and are cut by amajor extensional fault along their north-western mar-gins. The basins are underlain by a gradually south-ward thickening sequence of over a kilometre thick,ranging from the Triassic up to the Oligocene.

Methods

The structural and stratigraphical data for thisstudy were collected in the Pontones basin, in out-crops along the road from Pontones to Santiago de la

Espada, and in the Santiago de la Espada basin. Geo-logical maps and stratigraphic data from earlier stud-ies by Dabrio (1970, 1972) and Dabrio et al. (1971)were used in this study and modified where needed.

Rock or clay samples were collected about every 5to 10 meters of section for biostratigraphic studieswhich were performed by G.J. van der Zwaan andW.J. Zachariasse (Stratigraphy and Paleontologygroup of the Department of Earth Sciences at theUtrecht University). Paleobathymetry analyses of in-dividual marine clay samples using the ratio of plank-tonic and benthic foraminifera (P-B ratio) were per-formed by D.J.J. van Hinsbergen. An explanation ofthe methods of this paleobathymetry measurement isdescribed in Van Hinsbergen et al. (2005).

Kinematics and slip directions of faults were deter-mined on the basis of both structures on fault planes(such as tensile fractures, Riedel fractures, striations)and shear structures in fault gouges (Riedel, P, Y, R2and X shears and striations on these shear planes) asdescribed by e.g., Logan et al. (1979), Rutter et al.(1986), Gamond (1987), Hancock and Barka (1987),Petit (1987), Means (1987), Sylvester (1988), andWoodcock and Schubert (1994). In the absence oflineations on a fault plane, the slip direction along thefault was inferred on the basis of the geometrical rela-tionship between the main fault and secondary shearfractures, e.g., Riedel fractures.

A cross-section along the Pontones and Santiagode la Espada basins, parallel to the general direction ofshortening, was constructed on the basis of both out-crop and published map data (Dabrio, 1972). Con-struction and restoration of this cross-section was per-formed by conventional principles, methods and tech-niques following Dahlstrom (1969), Hossack (1979)and Groshong (2002). In general, line length and, ifpermitted, area balance techniques were used in thereconstruction and restoration of the cross-section.The restoration and balancing of the deformed sectionwas performed using the software program 2D-Move,kindly provided by Midland Valley Exploration Ltd*1. This program provides a number of restorationtools for balancing and restoration of a constructed(scanned and digitized) profile as well as tools for for-ward and backward modelling of restored and de-formed cross-sections, respectively.

*1 Midland Valley Exploration Ltd: main office in Glasgow, United King-dom. Website: www.mve.com.

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Chapter 3

40

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Miocene basins in the Betic fold and thrust belt

41

Leg

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Chapter 3

42

Tb

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timescaleLourens et al. 2004

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Miocene basins in the Betic fold and thrust belt

43

Miocene stratigraphy of the Pontones andSantiago de la Espada basins

The oldest Neogene sediments consist of an atleast 60 meter thick sequence of alternating grey-redlime and yellow sandstone and conglomerate beds offluvial-continental origin. This unit is mainly exposedin the Almorchón basin, and (unconformably?)overlies Paleocene-Eocene shallow marine deposits(column 3 in Fig. 3.4).

The continental sediments pass upwards into amarine sequence that forms part of the “Formacion deSantiago de la Espada”. Biostratigraphic analysis sug-gests that these marine sediments in the Pontones,Almorchón and Santiago de la Espada basins are ofmiddle Miocene age (Fig. 3.4). This result is at vari-ance with earlier studies by Dabrio (1970, 1972) andDabrio et al. (1971), who claim that the marinesediments at the base of the “Formacion de Santiagode la Espada” are early Miocene. We therefore assumethat the underlying fluvial-continental series is mostlikely of Oligocene to early Miocene age, which is inagreement with observations in other parts of thewestern Prebetics (e.g., Jerez Mir and Abril Hurtado,1979). The middle Miocene marine series (unit Ta,Fig. 3.4) consist of a 65 up to a possibly 230 meterthick unit of bioclastic and algal limestone beds occa-sionally intercalated with a marl or silt bed. The mas-sive limestone unit at the base is referred to as the“Calizas bioclasticas de Pontones” (Dabrio, 1970 and1972; Dabrio et al., 1971; and Fig. 3.5) and could betime equivalent to the “Roble Limestone” of Hermes(1978). The 230 meter stratigraphic thickness is mostlikely overestimated due to folding and thrusting. Theoccurrence of benthic foraminifera, such as repre-sentatives of Amphistegina, Elphidium, Miogypsina,Bolivina, and Borelis, and few planktonicforaminifera, such as Orbulina spp., G. trilobus and G.scitula, point to a middle Miocene age (pers. comm.G.J. van der Zwaan, W.J. Zachariasse and W.Renema). The lower limestone beds contain high per-centages of quartz detritus and fragments of algae,brachiopods, echinoderms, gastropods and undeter-minable larger foraminifera. According to Dabrio etal. (1971) the siliciclastic material was most likelyderived from metamorphic and igneous sources on theIberian Meseta. The litho-facies suggests a gradualtransition from shallow marine near the base, to shelfconditions at the top of this unit.

The middle Miocene unit (Ta) changes upwardsinto a series of white marls intercalated withmudstone, turbiditic calcarenite beds and mass flow

deposits (unit Tb), which reach a thickness of up to300 meters (Fig. 3.4). This value, again, is most likelyoverestimated due to folding and thrusting. The oc-currence of Orbulina spp., G. trilobus, G. menardii 3,P. mayeri and G. partimlabiata suggest a Serravallianage (pers. comm. G.J. van der Zwaan and W.J.Zachariasse), which in terms of geochronology corre-sponds with the time interval 12.77 to 12.07 Ma(Lourens et al., 2004). Paleodepth estimates using P-B ratios in individual samples indicate a water depthof 750 – 1000 m. Independent depth markers, e.g., S.reticulata and P. araminensis, suggest a depth of 500to 600 meters, however (pers. comm. D.J.J. vanHinsbergen). The lithofacies includes mass flow de-posits, and point to a relatively deep marine environ-ment. These mass flow deposits of the Santiago de laEspada Formation are time equivalent with the 200 to275 meter thick flysch-type deposits of the MoratallaMarlstone member near Moratalla to the east(Hermes, 1978; for location see Fig. 3.1).

The marls with intervening calcarenites pass up-ward into a thick and relatively massive unit ofdolomites or crystalline limestones, wackestones andgrainstones with large coral and bryozoa fragments.At Santiago de la Espada, however, these limestonesalternate with marl beds. The presence of menardii-type globorotalids in thin sections suggests a lateSerravallian – Tortonian age for this part of the se-quence (pers. com. G.J. van der Zwaan). This lime-stone unit passes upwards into a series of silty claysand calcareous turbidite beds rich in siliciclastic detri-tus (unit Tc/1). The dominance of algae and benthicforaminifera, such as representatives of Elphidium,Amphistegina, Ammonia, Borelis, Cibicides andSpiroplectammina may suggest that relatively shal-low marine conditions prevailed during this time.

In the Pontones and Almorchón basins these ma-rine series are tectonically sealed by thrusts. In theSantiago de la Espada basin (Fig. 3.6), however, theshallow marine series changes abruptly into a thickhomogeneous unit of red loam and fluvial silts, sands,and poorly sorted, angular conglomerates andbreccias of the Don Domingo Formation (unit Tc).According to Dabrio (1972), an angular unconformityseparates the marine sediments from these molasse-type continental deposits above, however, in the re-gion of Santiago de la Espada we have found no evi-dence of this unconformity. We assume, therefore,that the continental deposits are of most likelyTortonian age, and that they may be time equivalentsof the lacustrine deposits seen to the northeast nearYetas (Jerez Mir and Abril Hurtado, 1979) and of the

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Chapter 3

44

continental-lacustrine deposits near Hellin (for loca-tions see Fig. 3.1). The latter deposits contain micro-mammal fossils which point to a late Turolian(Messinian) age (Calvo et al., 1978). Note that in themost external parts of the Prebetics, towards thenorthwest near Villacarrillo, fluvial-continental de-posits intercalate with marine limestones and marlscontaining micro-faunas indicative of Tortonian andMessinian age (Martínez del Olmo and NuñezGaliana, 1973).

Structures in the Pontones and Santiago de laEspada basins

The structures observed in outcrops of the Creta-ceous, Paleogene and Miocene rocks comprise bothcompression and extension-related structures, i.e.,thrusts and folds, as well as low-angle to steeply dip-ping normal faults. In addition, there are few NW-SEoriented strike-slip faults (Figs. 3.7 and 3.8). In theMiocene deposits in particular, these structures areneither soft-sediment nor syn-sedimentary structures,i.e., the folds and most of the faults clearly developedafter sediment deposition. In the uppermost Miocenedeposits of the Santiago de la Espada formation (unitsTb) at Santiago de la Espada, however, small-scaleextensional faults show evidence of syn-sedimentarydisplacements.

Folds, in outcrops as well as on the scale of thegeological map, commonly have NE-SW trendingfold-axes, they are upright to inclined and their trueprofiles vary from close to open. The limbs of folds, inparticular in the turbidite deposits of the Miocene Tbunit, are often stretched and boudinaged. Few foldsare evidently thrust-related, such as in the case of dragfolds in the footwall, and fault-bend or fault propaga-tion folds in the hanging wall. Thrust planes in generaldip at low to gentle angles, and kinematic indicatorssuggest hanging wall transport to the northwest. In afew examples, however, hanging wall transport is tothe southeast. The strike-slip faults, both dextral andsinistral, have steeply inclined to vertical, NW-SEtrending fault planes (Fig. 3.7b), i.e., they are parallelto the general transport direction of the fold and thrustbelt, and clearly accommodate differentialdisplacements between adjacent structural units. Theslip-vectors on these fault planes vary in plunge fromhorizontal to oblique (66°), which makes these faultseither tear / transcurrent faults (Sylvester, 1988) orhanging wall drop faults (Butler, 1982), respectively.The compressional structures are associated with a

NW-SE direction of shortening (Figs. 3.3 and 3.7b).The extensional faults are mainly located along the

north-western margins of the Almorchón and San-tiago de la Espada basins (Fig. 3.3). They have devel-oped in Cretaceous, Paleogene and Miocene rocksand apparently cut earlier (outcrop-scale) compres-sional structures (Fig. 3.7a). The extensional faultsrun parallel to the main trend of the compressionalstructures in the region. The larger-scale extensionalfaults in particular dip to the southeast at moderate tohigh angles (Figs. 3.3 and 3.6). Slip vectors, such asgrooves and lineations indicate hanging-wall move-ment to the southeast, which is opposite to the maintransport direction of the fold and thrust belt. In theAlmorchón basin, extensional faults form a networkreminiscent of a relay pattern in the sense of Biddleand Christie-Blick (1985). Paleogene and Miocenemarker beds on both sides of these faults (observed inthe field and on the geological map; Dabrio, 1972) in-dicate that several tens up to 80 meters of throw (ver-tical displacement) has occurred along each of thesefaults, leading to a total of at least 500 meters ofthrow. In the Santiago de la Espada basin, on the con-trary, displacement is accommodated along a singleextensional fault, that may have up to several hun-dreds (>600) of meters of throw (Fig. 3.6). Theextensional structures in these basins are associatedwith NW-SE directed extension, which clearly con-flicts with the NW-SE direction of shortening de-duced from the compressional structures.

The cross-cutting relationships of extensional andcompressional structures, as shown in Fig. 3.7, sug-gest an extensional overprint on earlier compressionalevents. Fig. 3.8, however, shows a clear example of ageometric and genetic relationship between thrust tec-tonics and the local development of normal faults inthe hanging wall. This kind of structural relationship,in which normal faults develop in the hanging wall inresponse to a change in the fault plane geometry of thethrust or reverse fault, has been described previouslyby Coward (1983) and will be discussed below.

Restoration of the section constructed across thispart of the Prebetic zone leads to a viable solution(i.e., a viable cross section in the sense of Elliott,1983) as shown in Figure 3.9. In the restoration, thebase of the Miocene, which is in essence a pre-Miocene erosion surface, has been used as a referencelevel. The stratigraphic thickness of the differentMesozoic and Cenozoic units, as well as the depth to,and dip of the decollement zone in the Triassicevaporite, were estimated from surface data on thegeological map, as well as from published seismic

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Miocene basins in the Betic fold and thrust belt

45

Upper Cretaceous

Miocene (Ta)Miocene (Tb)

Miocene

thrust

Miocene

Upper Cretaceous

Upper Cretaceous

SW NWSE

Pontones

Fuente Segura

thrust A

Calizas de bioclasticas de Pontones

Figure 3.5. Panoramic view of the Pontones basin as seen from the NE along strike of the major thrust faults. On the lefta NW-vergent thrust (labelled A on Figure 3.3) emplacing upper Cretaceous rocks on top of Miocene sediments.Foreground shows upper Cretaceous and Miocene deposits in the hanging wall emplaced on Miocene sediments in thefootwall. The thrust forms part of an imbricate structure (Figure 3.3).

profiles and well data (Banks and Warburton, 1991).These estimates are to a large extent in agreementwith Banks and Warburton (1991) and Platt et al.(2003), as shown in Figures 3.2a and b. Thestratigraphic thickness of the Mesozoic units north ofPontones is approximately 1000 to 1400 meters (Fig.3.9) and increases towards the south to values of up to2500 to 3000 meters. This increase in thickness oc-curs stepwise across thrust faults, such as for exampleat fault A on the map and cross-section (Figs. 3.3 and3.9). These faults apparently are Mesozoicextensional faults, as shown in the restored cross-sec-

tion (Fig 3.9b), which were reactivated during theNeogene.

From the restored section (Fig. 3.9b) it is evidentthat, prior to the folding and thrusting in the Prebetics,the Pontones, Almorchón and Santiago de la Espadabasins formed part of a single large basin. Estimatesof the amount of shortening in the Tranco zone northof Pontones are up to 13 km, while compressionalstructures in the Pontones – Santiago de la Espadasection (P-P’) indicate that approximately 6.3 km ofshortening has occurred in this part of the Prebetics.This latter estimate is of the same order as the about 4

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46

Fault scarp

~~~~ ? ~~~~~~~~~

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Miocene (unit Tb)

Miocene (unit Tb)

Faultscarp in fig. A

Santiago de la Espada

Ba

r

arr ne B

j

c d

o

ol e me

Paleogene - Miocene

Miocene (unit Tc)

Riedel

ba

ba

Figure 3.6. Main features of the Santiago de la Espada basin. a) Fault scarp of the extensional fault along the north-western margin of the basin. Scarps reveal clear kinematic features such as Riedel fractures (136/76) and down-dipgrooves (141/56 on 141/57) and slickenside lineations (142/41 on 148/41) on the fault surface, indicating dip-slip motion.b) Panoramic view of the basin seen from the south. To the northwest tilted middle-late Miocene marine deposits(unconformably?) overlain by late Miocene - Pliocene continental sediments of unit Tc. Note exposed fault plane of Figure3.6a in far distance.

km of shortening estimated by Banks and Warburton(1991) and Platt et al. (2003) in the correspondingparts of their sections.

Discussion

Four different explanations were proposed in theintroduction that may account for the development ofextensional structures in the Betic fold and thrust belt.The first two would imply crustal-scale extension ofthe Betic fold and thrust belt for which there is clearlyno evidence. The other two explanations (i.e., stratal

extension in thrust-fault footwalls, or hanging-wallcollapse above shallowing-upwards reverse faults)both imply that the extensional structures are in someway inherent to the process of thrusting in the devel-oping fold and thrust belt. The available data lend sup-port to such a thrust-related origin of the extensionalfaults as follows.

Platt and Leggett (1986) explain the developmentof extensional faults in the footwall of a thrust plane interms of small-scale stratal extension, due to varia-tions in sliding resistance along the thrust plane,whilst the hanging wall block remains undeformed.These extensional faults are in essence low-angle nor-

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47

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Chapter 3

48

mal faults dipping in the transport direction of thethrust, i.e., they have the same orientations as Riedelfaults or shears. However, the extensional structuresseen in the Prebetic fold and thrust belt have devel-oped in the hanging walls of the thrust faults. Stratalextension of the kind described by Platt and Legget(1986) seems therefore inadequate to explain thestructures observed.

According to Coward (1983), normal faults maydevelop in the hanging wall of a thrust in response to achange in fault plane geometry of the thrust or reversefault, in a way that the dip-angle of the fault decreaseswith decreasing depth. Movement on such a fault, re-ferred to as a shallowing-upwards fault, can result instretching of the hanging wall block directly abovethe ramp-flat structure (Fig. 3.8b) in particular if thereis no layer-parallel slip along the bedding planes inthe hanging wall. Stretching in the hanging wall inevi-tably leads to the development of normal faults, whichinclude synthetic but mostly antithetic faults, i.e., withrespect to the sense of displacement of the thrust fault.The stretched hanging wall block as well as the geo-metrical relationships of the normal and thrust faultsin outcrops, such as the one shown in Figure 3.7, areclearly consistent with the structures expected for ashallowing upwards fault. The structural and geo-metrical relationships of the large-scale thrust faults Aand B, indicated on the map in Figure 3.3 and in theprofiles of Figure 3.9, and the main extensional faultsalong the north-western margins of the Almorchónand Santiago de la Espada basins are quite similar tothe smaller-scale structures seen in outcrop as shownin Figure 3.7. This strongly suggests the presence of ashallowing upwards fault at the scale of the thrustwedge, as shown in the profiles in Figure 3.9. Thepresent interpretation is possibly strengthened by thefact that fault A initially developed as a Mesozoicextensional fault, which was reactivated as a reversefault in the late Miocene, and became progressivelyshallower upwards during continued thrusting.

Both structural and stratigraphic data, such as thelithological correlations in the Miocene stratigraphy,indicate that the Pontones, Almorchón and Santiagode la Espada basins, prior to folding and thrusting inthis part of the Betic External Zone, have been part ofa single and large marine basin during the (early-)middle Miocene. This large marine basin most likelyformed a connection between the Atlantic Ocean andthe Mediterranean Sea, often referred to as the “NorthBetic Strait” (e.g., Calvo et al., 1978; Sanz deGaldeano and Vera, 1992; Soria et al., 1999; Martín etal., 2001; Braga et al., 2003; Sanz de Galdeano and

Alfaro, 2004).The “abrupt” facies change from a continental –

shallow marine facies in the late Oligocene – earliestMiocene to a deep marine setting in the middleMiocene can not be explained with a global (eustatic)sea level change. As a matter of fact, the middleMiocene period is characterised by a high (up to 100-150 m) but decreasing eustatic sea level (Haq et al.,1988, and Fig. 3.4). However, paleobathymetry esti-mates of the middle Miocene sediments show a rapidrelative sea level rise (up to 500-600 meters of waterdepth) in an evidently underfilled basin. Both themagnitude of the relative sea level rise inferred frompaleobathymetry and the sedimentary facies domi-nated by mass flow and turbidite deposits imply a tec-tonic cause for the rapid basin subsidence. These ob-servations are in agreement with the Miocene subsid-ence histories inferred for both the Jumilla – Cieza re-gion by Kenter et al. (1990) and for the westernPrebetics, in particular the Santiago de la Espada re-gion, by Hanne et al. (2003). Two obvious candidatesfor such a tectonic cause are: (1) the migratingdepocenter of the foredeep domain ahead of the mov-ing thrust mass of the fold and thrust belt (Kenter etal., 1990; Beets and De Ruig, 1992; Hanne et al.,2003) and (2) a rapid (“instantaneous”) loading of theIberian plate due to the emplacement of both theSubbetic and the Betic Internal Zones. According toBeets and De Ruig (1992) and Hanne et al. (2003) thefirst foreland basins already developed in the lateOligocene in the south due to loading of the Iberianplate by the Betic Internal Zone. This foreland basinsystem progressively migrated towards the northwestduring the early and middle Miocene when it reachedthe realm of Santiago de la Espada and Pontones. Therapid subsidence of the Pontones – Santiago de laEspada basins, therefore, most likely reflects the ap-proaching foredeep basin in front of the moving thrustmass. It is noted, however, that according to Van derBeek and Cloetingh (1992) the load of the Betic Ex-ternal and Internal Zones is insufficient to explain theflexure of the Iberian lithosphere as observed today.

The subsequent change in the basins from marineto continental facies and allied basin uplift in theTortonian is almost coeval with the initiation of fold-ing and thrusting in the western part of the PrebeticZone, and evidently marks the moment that the for-ward propagating deformation front of the fold andthrust belt has approached the realm of Pontones andSantiago de la Espada. Since the Tortonian (~10 Ma),approximately 50 to 60 km of shortening has been ac-commodated in the western Prebetics, and the region

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Miocene basins in the Betic fold and thrust belt

49

Road C321 / A-317 Pontones

N E S

ds

~65 m

10 m

normal fault 136/74thrust 142/47

U. Cretaceous rocks

a

a

U. Cretaceous -Miocene rocks

CD E

O

A’

BB’

A

B�

b

Figure 3.8. a) Deformational structures in Upper Creta-ceous rocks north of Pontones. Thrust to the left emplacesUpper Cretaceous rocks in the hanging wall on UpperCretaceous and Miocene deposits in the footwall. In thehanging wall, a normal fault transports material over adistance (ds) of ~2 m in a direction opposite to themovement direction of the thrust. The normal and thrustfaults appear to be genetically and geometrically related.b) Diagram taken from Coward (1983), showing howextensional strains are developed above a shallowing-upwards fault. If bed A would be displaced to point D andthere is no change in length of the beds along the hangingwall of the fault, then for a supposed kink band geometry,bed length CD will have to expand to length A’C.Stretching of the hanging wall over the top of such ashallowing-upwards fault increases with increase of α anddecrease in β. This extension may occur over a broad zonein the hanging wall, leading either to ductile extension ofthe upper layers or to the development of brittleextensional faults.

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Chapter 3

50

1000

500

0

-500

-1000

-1500

-2000

-2500

-3000

-3500

~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~

2000

1500

1000

500

0

-500

-1000

-1500

-2000

-2500

-3000

P Pontones basin

Jurassic

Triassic

Iberian basement

lower to middle Cretaceous

Miocene (Ta-Tb)

A

x

x

y

y

Iberian basement

Pontones basin

NW

NWShortening in the Tranco zone ~13 km

A

B

Jurassic (~600 m)

Triassic

lower to middle Cretaceous (~700 m)

Cretaceous (> 300 m)

middle-late Miocene (> 300 m)

Top Tortonian at 500 m

Jurassic (600 - 700 m)

Miocene (~150 m)

Triassic?

10000 2000 3000 m

meters

meters

Tranco zone

unconformity

imbricate fan

Figure 3.9. a) Balanced section across the region of Santiago de la Espada. For location of section see Figures 3.1(profile 3) and 3.3 (profile P-P’). b) The Santiago de la Espada cross-section restored. Marker pins X, Y and Z referto common points in the deformed and undeformed sections. White Roman letters refer to large-scale thrusts shownon the map in Figure 3.3. Stratigraphic thickness of the Mesozoic and Cenozoic units increases towards the SE asevident from field observations, the IGME geological map (Dabrio, 1972), and documented by, e.g., Hernandez et al.(1980), Hermes (1978), and Banks and Warburton (1991). Folding and thrusting initiated after deposition of theMiocene sediments, i.e., in latest Miocene and Pliocene period. Mesozoic extensional structures were reactivated asreverse faults (e.g., fault A), which shallow upwards. Tectonic transport direction is in general towards the NW.Normal faults, however, move with opposite movement sense. Total throw along sets of normal faults is up to 500-600meters. Main decollement is located within the Triassic horizon, which dips at ~2.5° to the SE. The minimum amountof shortening in the Santiago de la Espada region is estimated at 6.25 km.

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Miocene basins in the Betic fold and thrust belt

51

~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~ ~? ? ?

P’Santiago de la Espada basinAlmorchón basin

Cretaceous

Triassic

Pg

Jurassic

Cretaceous

Tb

Ta-b

B

C

D E

AB

C D

E

z

z

Santiago de la Espada basin

Almorchón basin

SE

SE

Shortening in the Santiago de la Espada section at least 6.25 km

Triassic

Jurassic (1500 to 2000 m thick)

Cretaceous (up to 1000 m thick)

Paleogene - earliest Miocene

Triassic

Jurassic

Cretaceous

Mesozoic extensional fault

Triassic?

Triassic?

Miocene (Tb)

Paleogene -earliest Miocene

Miocene (Ta)

Miocene (200-275 m)

Miocene-Pliocene (~200 m)

?

>690 mTc

~530 m

1000

500

0

-500

-1000

-1500

-2000

-2500

-3000

-3500

meters

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52

Chapter 3

has experienced an uplift of up to 2000 meters. Assuggested in chapter 2, African-Eurasian plate con-vergence is most likely responsible for this lateMiocene to recent compressional deformation in theBetic External Zone. Regional uplift of the BeticCordillera continues still today (e.g., Giménez et al.,2000; Braga et al., 2003; Sanz de Galdeano andAlfaro, 2004). A study of the causes of this uplift liesbeyond the scope of this study, however, part of thelate Miocene to recent uplift may well be related toflexural isostatic processes (e.g., Watts, 1992) of theIberian lithosphere beneath.

Conclusions

Careful analysis of the structure and stratigraphyin the Miocene basins in the western part of thePrebetic Zone near Santiago de la Espada leads to thefollowing conclusions:

The development of outcrop and map-scaleextensional structures is related to the process ofthrusting and reverse faulting. Some of these reversefaults were probably initiated as Mesozoic exten-

sional faults, and were reactivated as reverse faults inthe late Miocene in response to continued thrustingand folding in the Betic External Zone. During faultpropagation upwards, the reverse faults became pro-gressively shallower (i.e., less inclined), which inevi-tably led to extension in the hanging wall block andthe subsequent development of normal faults. Thisimplies that irrespective of the extensional nature ofsome of the faults, the Miocene basins in the Prebeticsin essence developed in a compressive setting.

Prior to the late Miocene folding and thrusting inthe Prebetics, the Pontones, Almorchón and Santiagode la Espada basins formed part of a large marine ba-sin in the early-middle Miocene. The abrupt subsid-ence of this basin during the middle Miocene waslikely associated with the migrating foredeep part ofthe foreland basin, in front of the growing thrust massof the External Zone. Shallowing of the basin in thelate Miocene was immediately followed by the onsetof folding and thrusting in the western Prebetics,which led to segmentation of the large basin intosmaller basins and closure of the northern Atlantic-Mediterranean connection.