deJong_Alpine Tectonics Betics & Pyrenees: Iberia's Rotation Pole Changes_Tectonophysics 1990

8/3/2019 deJong_Alpine Tectonics Betics & Pyrenees: Iberia's Rotation Pole Changes_Tectonophysics 1990

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Tectonophysics, 184 (1990) 219-296Elsevier Science Publishers B.V., Amsterdam

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Alpine tectonics and rotation pole evolution of IberiaKoen de Jong

Instrtute for Earth Sciences, Free Unruersity, P.O. Box 7161, 1007 MC Amsterdam (The Netherlands)(Received April 20,1989; revised version accepted January 26. 1990)

ABSTRACTDe Jong, K., 1990. Alpine tectonics and rotation pole evolution of Iberia. In: G. Boillot and J.M. Fontbote (Editors), Alpine

Evolution of Iberia and its Continental Margins. Tectonophysics, 184: 279-296.The geological evolution of the Betic Cordilleras and Pyrenees reflects the Cretaceous and Tertiary rotation pole and

kinematic evolution of the Iberian and African plates. New constraints on the Alpine tectonic evolution of the Iberian plateare provided by P-T-t data and regionally consistent stretching lineations from the metamorphic parts of the BeticCordilleras.

High-pressure low-temperature metamorphism in the Betic Cordilleras resulted from continent-continent collision whichcaused subduction to a maximum depth of 37 km. A preliminary 116 + 10 Ma radiometric age for this event corresponds tothe initiation of seafloor spreading to the west of Iberia which lasted until about 80 Ma. Intracontinental thrusting in theBetics between 99 Ma and 83 Ma took place after subduction ended. E-W to ESE-WNW trending stretching lineationsindicate the direction of thrusting, which resulted in extensional strains of 200-600%. The timing of thrusting in the Beticscoincides with a 95-80 Ma tectonic phase in northern Africa, during which E-W stretching lineations were formed. Thestretching Iineations are coincident with the 110-80 Ma motion vector of Africa-Iberia with respect to Eurasia. Thrusting inthe Betics and deformation in northern Africa was driven by convergence of Africa-Iberia and Eurasia. Cretaceousdeformation is further recorded by terrigeneous sedimentation in the Mauritanian Flysch and by the tectosedimentaryevolution of the Malaguide Complex. Crustal thinning, magmatism and metamorphism in the Pyrenees during the 110-85 Maperiod is governed by a left-lateral strike-slip of Africa-Iberia with respect to Eurasia around the same rotation pole asthrusting in the Betics.

During the 80-54 Ma period the rotation pole was situated west of Gibraltar, near the previous active collision zone. Thisinhibited large-scale overthrusting and related penetrative deformation in northern Africa and the Betic Cordilleras.Deformation was instead transferred to the northern boundary of Iberia, now acting as an African promontory. From theCampanian on wards, oblique convergence took place around the combined Gibraltar rotation pole. Deformation culminatedin the late Eocene, corresponding to spreading in the Norwegian-Greenland Sea at 55 Ma which induced an additionalcompression in western Eurasia. During the Pyrenean collision, high-pressure metamorphic rocks in the Betic Cordilleras wereexhumed and they cooled substantially. The cooling trend was disturbed by Oligocene extensional deformation andintroduction of a transient heat source, which correlates with the mantle being uplifted during extension. Heating culminatedat the Oligocene-Miocene boundary in the Betic Cordilleras and in northern Africa. This evolution agrees with thedevelopment of a plate boundary between Iberia and Africa at 30 Ma, after completion of the Pyrenean collision. The newplate boundary was connected to the western European rift system.

Renewal of compression and overthrusting in the Betic Zone took place after 20 Ma. Overthrusting is succeeded by twophases of wrenching, juxtaposing crustal segments with different Moho depths inherited from the late Oligocene to EarlyMiocene extension.

IntroductionThe Iberian peninsula (Fig. 1) is bordered by

two Alpine foldbelts, the Pyrenees to the northand the Betic Cordilleras to the south. These beltsseparate the Iberian plate from, respectively, theEurasian plate and the African plate. The Creta-ceous to Tertiary tectonic evolution of the Pyrenees

has been well documented (Mattauer and Henry,1974; Puigdefabregas and Souquet, 1986; Soula etal., 1986). Until now the Betic Cordilleras hasmost often been regarded as a Tertiary orogen,mainly on the basis of important Tertiary defor-mation in the non-metamorphic parts (Rondeeland Simon, 1974; De Smet, 1984). A Mesozoic agefor the early deformation has, however, been sug-

0040-1951/90/$03.50 0 1990 - Elsevier Science Publishers B.V.


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gested by Kampschuur and Rondeel(l975) owingto the Mesozoic age of the flysch deposits in thewestern Betics. New data discussed in this paperalso suggest important Cretaceous tectonics in themetamorphic Internal Zone of the BeticCordilleras. Ceochronological studies in the Al-pine collision belt of northern Africa (Monie etal., 1984a; 1988) show a tectonic evolution whichis comparable to that of the Betics-Cretaceousmetamo~~c ages and an important Tertiary re-setting. This paper aims at tying the new tectonicmodel for the Betic Zone and the thermotectonicevolution of the northern African belt to the well-constrained tectonic evolution of the Pyrenees.The tectonic evolution of the erogenic beltsbordering the Iberian plate will be shown to beconsistent with the Cretaceous and Tertiary rota-tion pole and kinematic evolution of the Iberianand African plates discussed by Savostin et al.

K. DE JONG

(1986) Srivastava and Tapscott (1986) and Klit-gord and Schouten (1986).

Regionally consistent stretching lineationswhich were formed during early Alpine thrustingat lower crustal levels are a salient feature of thetectonic evolution of the Internal Zone. Theycoincide with the mid-Cretaceous motion vector ofSavostin et al. (1986) of the African plate (includ-ing Iberia at that time) with respect to Eurasia.Because stretching lineations appro~mate themovement direction in shear zones (Esscher andWatterson, 1974) they probably also trace platemotion directions. A relationship between the di-rection of thrusting and plate motion has beensuggested for an number of orogens (Shackletonand Ries, 1984), including the Alps (Baird andDewey, 1986; Choukroune et al., 1986). Duringthe later stages of the tectonic evolution of thearcuate western Alps, radial thrusting occurred

IBERIAN MESETA

GULF DE LION

Infernal Zones of the Betlc Cordilleras and Rlf;:

DIAbne metamorphic rocks)a

Fig. 1. Sketch map of the westernmost Mediterranean area (modified after Ricou et al., 1986) showing the major Alpine structuralprovinces. The eastern Betic Cordilleras of southern Spain are delineated.


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ALPINE TECTONICS AND ROTATION POLE EVOLUTION OF IBERIA 281

(Choukroune et al., 1986) at higher crustal levels,this thrusting clearly bearing no relationship toplate motion vectors. Relatively large finite dis-placements and rotations during such a stage willdisturb the original pattern of older stretchinglineations formed at deeper levels. The easternpart of the Betic Cordilleras does not demonstratean arcuate form, and therefore no pervasive re-orientation of older structures is to be expected.

Palaeostress analyses in stable forelands do notusually suffer the disadvantage of reorientation, asfinite strain is in general small. A clear relation-ship between (successive) palaeostress directionsand plate motion vectors is therefore recorded inthe Alpine foreland (Letouzey, 1986; Bergerat,1987). However, in Iberia the Mesozoic palaeos-tress directions do not mimic the plate vector veryaccurately (Malod, 1989). This is partly the resultof reorientation and heterogeneities induced bypre-existing faults. Therefore, the regionally con-sistent stretching lineations in the Internal Zoneare considered as an important constraint in theearly kinematic evolution of the Iberian plate.Motions around different rotation pole positionsduring orogeny will, due to overprinting and re-orientation, not be recorded by successive genera-tions of stretching lineations. Shifting of rotationpoles has, however, a marked effect on tectonics inmetamorphic belts, as will be discussed later.Evolution of the tectonic zones bordering Iberia

The Alpine collision belts bordering Iberia arecharacterized by Jurassic to Early Cretaceous ex-tensional deformation and related strike-slip de-formation. A Middle to Late Jurassic strike-slipfault between Iberia-Africa is indicated by platereconstructions (Savostin et al., 1986; Klitgordand Schouten, 1986). The occurrence of a frag-ment of an ophiolite sequence of Late Jurassic agein northern Africa (Bouillin et al., 1977) accordswith these reconstructions. Continuing motion intothe Cretaceous is indicated by flysch depositsculminating in Aptian-Albian times in the FlyschDomain (Bouillin et al., 1986). The non-metamor-phic External Zones of Iberia and northern Africaare palaeogeographically unrelated (Bouillin et al.,1986) this also indicating their initial separation.

In the External Zone of the Betic Cordilleras analgal platform broke up at the Middle to LateJurassic boundary (Geel, 1979) resulting in strongpalaeogeographical differentiation (Hermes, 1978).An extensional tectonic regime is indicated bypillow basalt intrusions in the Sub-Betic (Hermes,1978; De Smet, 1984). Important hiatuses, turbi-dite deposits and the occurrence of Middle Jurassiclithoclasts in Albian-Aptian marls (Hermes, 1978)indicate important vertical motions continuing intothe Cretaceous. Basaltic intrusion in the InternalZone of the Betics is of Jurassic age (146 + 3 Ma,Rb/Sr age, Hebeda et al., 1980; 200 f 5 Ma,K/Ar biotite age, Besems and Simon, 1982).

At the northern boundary of Iberia, in thefuture Pyrenees, carbonate platform breakup oc-curred in the Early to Middle Jurassic(Puigdefabregas and Souquet, 1986). At the north-western margin of Iberia, Late Jurassic riftingpossibly occurred; important rifting started in theBerriasian to earliest Valanginian (144-140 Ma,Boillot et al., 1989). The end of emplacement ofultramafic rocks by ductile normal faulting hasbeen dated at 122 f 0.6 Ma (Feraud et al., 1988).Final emplacement by brittle deformation oc-curred before the late Aptian breakup unconfor-mity (around 115 Ma), which marks the onset ofseafloor spreading (Boillot and Malod, 1988; Boil-lot et al., 1989; Malod, 1989). Opening of the Bayof Biscay occurred between the Aptian andCampanian and induced several hundred kilo-metres of strike-slip on the North Pyrenean Fault(Le Pichon et al., 1971; Choukroune and Mat-tauer, 1978; Savostin et al., 1986; Srivastava andTapscott, 1986; Klitgord and Schouten, 1986;Boillot and Malod, 1988; Malod, 1989). Deforma-tion coincided with a general change in the sedi-mentation pattern in Aptian times (Souquet et al.,1985) and Pyrenean magmatism and metamor-phism between 110 and 85 Ma (Albarede andMichard-Vitrac, 1978; Montigny et al., 1986).During this period the North Pyrenean Fault zonewas characterized by high heat flow in response tocrustal thinning related to strike-slip tectonics(Choukroune and Mattauer, 1978; Vielzeuf andKornprobst, 1984; Golberg et al., 1986). EarlyCretaceous metamorphic ages are also well re-corded by the 4oAr-39Ar stepwise heating method


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of samples from mylonite zones in the easternPyrenees (110-100 Ma and 90 Ma, Costa andMaluski 1988) and in strike-slip basins at thenorthwestern termination of the Iberic Cordillera(100 Ma, Golberg et al., 1988).

The extensional regime in the Pyrenees changedto compression in the latest Cretaceous (Vielzeufand Kornprobst, 1984); oblique convergencestarted in the Campanian (PuigdefBbregas andSouquet, 1986). Strike-slip ceased dr~aticallyduring the middle Eocene, when major thrustswere developed parallel to the mylonite zones(Soula et al., 1986). The northern Spanish passivemargin was converted into an active margin in thePaleocene-Eocene interval as a result of plateconvergence (Boillot and Malod, 1988). Duringconvergence, Variscan and Early Cretaceous faultswere reactivated (Soula et al., 1986; McCaig andMiller, 1986; Majoor, 1988). Radiometric ages inmylonite zones indicate a latest Cretaceous tomiddle Eocene age for reactivation (McCaig andMiller, 1986; Costa and Maluski, 1988;Majoor,1988). Compressional deformation in the non-metamo~hi~ zones cul~nated in the Eocene(Mattauer and Henry, 1974; Puigdefabregas andSouquet, 1986). Piggy-back thrusting in the centralsouthern Pyrenees migrated southward with time(Williams, 1985), concomitant with the progressivesouthward development of molasse basins (Mat-tauer and Henry, 1974) and their incorporation insubsequently formed thrust units (Puigdef~bregasand Souquet, 1986). Flexure modelling ildb shownthat the Ebro foreland basin formed as a result ofthe Pyrenees, Catalan Coastal Range and IbericCordillera loads (Zoetemeijer et al., 1990). It con-tains Paleocene and thick Eocene-Oligocene de-posits, recording erosion of rising mountain chains(Mattauer and Henry, 1974; Nagtegaal and DeWeerd, 1985). The southward propagating defor-mation front reached the Ebro Basin after deposi-tion of the Oligocene molasse (Williams, 1985).

The interiors of the Iberian plate also experi-enced extensional phases during the Late Jurassicto Early Cretaceous period, which accord wellwith phases of spreading in the North Atlantic(Malod, 1989). Late Eocene to late Oligocenenorthward compression reactivated Variscan andMesozoic faults (GuimerB, 1984; Viallard, 1985).

Raising of geotherms during extension caused ef-fective weakening of the lithosphere, which is alsoa si~ificant factor in the localization of in-tracontinental Tertiary compression (Zoetemeijeret al., 1990).

The External Zone of the Betic Cordillerasrecords only relatively small increments of thePyrenean collision (De Ruig, this issue). This colli-sion resulted in differential block movements dur-ing the Paleocene and Eocene (Kenter et al., 1990).The main compression occurred at the Middle toLate Miocene boundary (Simon, 1987) duringnorthward thin-skinned tectonics with intermit-tant strike-slip deformation (De Ruig et al., 1987).Compression also began in the Middle Miocene inthe Tell and External Rif of northern Africa.

In the following I present indications of animportant phase of tectonism in the Internal Zoneof the Betic Cordilleras between the Barremianand Campanian.Tectonic evolution of the Internal Zone of theBetic Cordilleras

Regionaf scale structureThe Internal Zone occurs to the south of the

External Zone, which constitutes the margin ofIberia characterized by Mesozoic rifting (Hermes,1978). These two zones are presently separated bya strike-slip fault of Miocene age (Hermes, 1978;Le Blanc and Olivier, 1984; De Smet. 1984). TheTriassic stratigraphy of the overthrust units in theInternal Zone bears no resemblance to the Triassicstratigraphy of the External Zone (Simon, 1987).The Internal Zone can thus be considered as alloc-hthonous to Iberia. (Very) low grade Triassicmetarno~~c rocks of the Almagride Complexoccur in windows below the Alpujarride Complexof the Internal Zone (Fig. 2); they show strikingsimilarities with Triassic rocks of the eastern Sub-Betic (Besems and Simon, 1982; Simon, 1987).This indicates an original overthrust contact be-tween the Internal Zone and rocks that were prob-ably separated from the External Zone duringMesozoic rifting. The Miocene strike-slip faultbetween the Betic Zone and the External Zone cannot therefore be interpreted as a plate boundary,


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but as a later structure of relatively minor impor.tance.

The Internal Zone consists of four stackedcrustal segments. The lowest segment, the VeletaComplex, is characterized by low-pressure, low-temperature (LP/LT) metamorphism (Puga andDiaz de Federico; 1978), for which recently apre-Alpine age has been suggested (Gomez-Pugnaire and Franz, 1989). This complex has beenoverthrust by the Mulhacen Complex, which bearsevidence of early Alpine high-pressure, low-tem-perature (HP/LT) metamorphism partially over-printed by medium-grade metamorphism (Gomez-Pugnaire and Femandez-Soler, 1987; Bakker etal., 1989). The Alpujarride Complex occurs on topof the Mulhacen Complex and is also char-acterized by early H P/LT met~o~~srn (Bakkeret al., 1989; GoffC et al., 1989). In the westernBetics the Alpujarride Complex experienced high-temperature metamo~~sm at high to low pres-sure related to emplacement of ultramafic rocks(Westerhof, 1977) of Early Miocene age (Priem etal.. 1979; Zindler et al., 1983). The uppermosttectonic unit, the Malaguide (Ghoma~de in north-ern Africa) Complex, is almost entirely non-meta-morphic (Egeler and Simon, 1969). Condensed,but continuous, Mesozoic and Paleogene stratigra-phy (Roep, 1980) indicates that this crustal do-main has always retained near the crustal surface.The Dot- sale calcaire and Pre-dorsalian Zone rep-resent the margin of this Malaguide/Ghomaridedomain with dominant Triassic and Jurassic shelfand slope sedimentation (Bouillin et al., 1986).

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00 4cxJ 500 600T(C)Fig. 3. P-T-r paths of the Mulhacen Complex (light stipple)and the Almanzora Unit (dark stipple). Tourmahne K/Ar.agesfor D, and D,_t are indicated. Ages (Ma) of deformationphases used to constrain exhumation are indicated by circles,inferred ages by diamonds. Boxes indicate P-T conditions(after Bakker et al., 1989). (I ) albite-jadeite + quartz (Newtonand Kennedy, 1968); (2) glaucophane stability (Maresch,1977); (3) staurolite-in (Hoschek, 1969); (4) anorthite+HzO= kyanite f zoisite + quartz (Newton and Kennedy, 1963); (5)

Al-silicate triple point (Holdaway, 1971).

Tectonometamorphic evolution

The pressure temperature (P-T) and tectonicevolution of the Mulhacen Complex and the Al-manzora Unit of the Alpujarride Complex hasbeen reconstructed by Bakker et al. (1989) byrelating deformation and metamorphism and P-Tdeterminations by microprobe analyses.

Both complexes are characterized by initialHP/LT metamorphism, indicating the dis-turbance of the pre-existing pattern of isothermsby subduction. The maximum metamorphic pres-

sures in the Mulhacen Complex (1 GPa) (Fig. 3)occur in the western Sierra Nevada to the easternSierra de 10s Filabres (Velilla and Fen011 Hach-Ali,1986; Gomez-Pugnaire and Fernandez-Soler, 1987;Bakker et al., 1989). The Almanzora Unit experi-enced a pressure of 0.7 GPa (Bakker et al., 1989),which agrees with the maximum pressure in theother Alpujarride units (GoffC et al., 1989). Subse-quent isobaric heating in both units indicates thestarting relaxation of the disturbed pattern of iso-therms and cessation of further underthrusting ofcooler crustal segments (England and Thompson,1984). At the end of the isobaric heating trajectorythe first phase of penetrative deformation (D,_,)took place. The most penetrative deformation (D,)occurred at peak temperature conditions of about


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ALPINE TECTONICS AND ROTATI ON POLE EVOLUTION OF IBERIA

570 C in the Mulhacen Complex and 450 o C inthe Almanzora Unit (Fig. 3) (Bakker et al., 1989).

D, effectively transposed all earlier fabrics. Inthe Almanzora Unit D,_, deformation fabricswere only left as internal fabrics in glaucophaneand crossite crystals. In the Mulhacen Complexpenetrative D,._, fabrics were left unaffected onlylocally in glaucophane schist facies amphibolites,local conglomerate bodies and in the core of a 5.5km3 gneiss body. D, _ 1 is characterized by E-W toESE-WNW trending stretching lineations (Fig. 2)on flat-lying foliations. Strain analyses on con-glomerate pebbles indicates extensional strains oflOO-300% (Fig. 4A). Borradaile (1976) inferredESE-WNW extensional strains in excess of 150%in the gneisses. Top-to-the-west shear is indicatedby asymmetric tails at the extremities of K-felds-par porphyroclasts. Ductile D,_ 1deformation wasthe result of thrusting within the Mulhacen Com-plex (Bakker et al., 1989). D, is also characterizedby important regionally consistent E-W to ESE-WNW stretching (Fig. 2) paralleling axes of lo-cally pronounced sheath-like folds. Plagioclaseporphyroclasts in amphibolites record extensionalstrains of 200-600% (Fig. 4B) which, however,might contain an unknown D,_, component. Aminimum D, stretching amount is provided by therotation of syn-D, garnets. Rotation angles of90-110 indicate a shear strain of 3.5 using theRosenfeld (1970) equation. If this rotation is

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285

formed during plane strain, the shear strain indt-cates 250% extension. Vissers (1989) arrived at asimilar strain estimate. Because the contact be-tween the internal garnet fabric and the externalfabric is lost most often during post-blastesis D,deformation and garnets have been boundinaged,this amount is a minimum estimate. Strain analyseson pebbles in quartzites with a penetrative D,fabric in the Alpujarride Complex indicate 300-600% extension (Fig. 4C). Similarity of main phasedeformation fabrics in both complexes combinedwith comparable amounts and directions ofstretching directions indicate a similar tectonichistory. Strain in the Alpujarride Complex ismuch more heterogeneously developed, the core ofthe conglomerate body indicating maximum ex-tension of 125% (Fig. 4C).

On the basis of the models of Davy and Gillet(1986) the P-T-t path of the Alpujarride Com-plex is explicable by a screening effect of theunderlying Mulhacen Complex and heating to-gether with loading by overlying crustal segmentsduring D,. In the last stages of D, the MulhacenComplex overthrust the Veleta Complex (Fig. 5B),the latter never having experienced very deeptectonic burial. The contact zone is characterizedby mylonites with E-W trending stretching linea-tions (Fig. 2). Preferred orientations of quartz-c-axes indicate top-to-the-west shear (De Jong, inprep.). Large-scale imbrication occurred in a crust

Fig. 4. Log-strain diagram (Wood, 1974) presenting strain analyses from the Internal Zone; averages indicated by triangles. (A) D,_ 1quartz pebbles, Mulhacen Complex, location (c) in Fig. 2, central Sierra de 10s Filabres. (B) D, elongated plagioclase porphyroclastsin amphibolites, Mulhacen Complex, location (d) in Fig. 2, eastern Sierra de 10s Filabres. (C) D, quartz pebbles, high strain areaindicated by crosses, low strain area by dots, Alpujarride Complex, location (i) in Fig. 2, central Sierra de las Estancias; the singlesquare displays one determination from the Almanzora Unit near location (8).


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A LD, CONFIGURATION 1 AU MCi- T--r 80-85 Ma

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Fig. 5. (A) Crustal configuration during the 85-80 Ma phase of crustal scale imbrication a peak thermal conditions during D, (insetsarc P-T-f paths). (B) Oligocene extension of the crustal wedge. juxtaposing the Almanzora Unit (AU) and the Mulhacen Complex(MC) during D, + ,

that had already been largely thermally equi-librated after initial thermal disturbance by sub-duction. The Almagride Complex is also char-acterized by ESE-WNW trending stretching lin-eations (Fig. 2) in carbonate mylonites, whichrecord top-to-the-west shearing by asymmetricpull-apart structures.

After D,, decompression and cooling dominatesthe thermal evolution, in the absence of anypenetrative deformation. The form of the P-T-ttrajectory precludes rapid exhumation of the In-ternal Zone. Continued thermal equilibration to-wards a new steady-state isotherm generates anincreasing apparent geothermal gradient, whichreaches about 35O/km. Erosional unroofing pre-sumably played an important role during thistrajectory.

The Almanzora Unit was placed on top of theMulhacen Complex along a low-angle ductile nor-mal shear zone during D,, ,, which represents aphase of heterogeneous crustal thinning and ex-tension (Bakker et al., 1989). The latter authorsestimated the excision of a 6 km thick crustalsection along a single shear zone during eastwardslip of the hangingwall. Thermal consequences ofextension were retarded; temperature increasestarted during Dx+* and culminated during theD X+3 second thermal peak at 525 o C (Bakker etal., 1989). Retardation of heating at a particularcrustal level with respect to the timing of exten-sion is also shown by thermomechanical models ofcrustal extension (Crough and Thompson, 1976;Thompson, 1981; Moretti and Froideveaux, 1986).Crustal thickening by large-scale S-verging D, +Z


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ALPINE TECTONICS AND ROTATI ON POLE EVOLUTION OF IBERIA 287

folding and associated thrusting enhanced this.The second thermal peak may indicate the effectof magma addition (Thompson, 1981) or introduc-tion of magmatic fluids (England and Thompson,1984) at higher crustal levels. The large amount ofheating, about 100 o C (Fig. 3) indicates the intro-duction of a transient heat source in the InternalZone; this was the consequence of crustal andlithospheric mantle extension and resulted in em-placement of ultramafic rocks in the western Bet-its (Platt, 1987; Bakker et al., 1989; Doblas andOyarzun, 1989) in the Early Miocene (22 + 4 Ma,Priem et al., 1979; 21.5 + 1.8 Ma, Zindler et al.,1983). The timing is coincident with intrusions ofsimilar rocks in the Rif in Morocco (Ben Othmanet al., 1984). The Dx+3 thermal peak in the easternInternal Zone is consequently an Early Miocenefeature. The regional importance of this event isindicated by widespread resetting of radioisotopesystems at the Oligocene-Miocene boundary inthe Kabylian Massifs in northern Algeria and theExternal Rif (MoniC et al., 1984a, b, 1988).

Timing of tectonic events in the Mulhacen Com-plex

In order to constrain the tectonic evolution ofthe Betic Zone we need to know the ages of thevarious tectonometamorphic phases. However,only very limited radiometric data on metamor-phic minerals are available. For the MulhacenComplex an average muscovite Rb/Sr age of 13.8Ma has been calculated from data reported byPriem et al. (1966) and Andriessen et al. (1989).The latter authors also report a 12.8 Ma biotiteRb/Sr age. In addition, in the context of a feasi-bility study of K/Ar dating of tourmalines threeages of 80-85 jI 8 Ma and one age of 116 f 10 Mahave been reported (Andriessen et al., 1989). Thesuggested blocking temperature for tourmaline ofabove 600 C (Andriessen et al., 1989) well ex-ceeds the maximum metamorphic temperature inthe Mulhacen Complex. These ages can thereforebe interpreted as metamorphic crystallization ages.The 80-85 Ma ages (averaging 83 Ma) are ob-tained from gneisses with a D, mylonite fabric, inwhich synkinematic growth of tourmaline has oc-

curred. This suggests an age of 83 Ma for D,. Forthe 116 Ma age, excess Ar might be put forward asan explanation as this component has a widespreadoccurrence (Hebeda et al., 1980); in addition, thisis an explanation which is difficult to disprove.Alternatively, this age can reflect pre-D, ~, or D, _.metamorphism. It will be shown that the 116 Maage is actually in accordance with the time scalefor the establishment of the Mulhacen ComplexP-T-t path, and can be geologically relevant.

A point which constrains the P-T-t path ofthe Mulhacen Complex is D,, which occurs at adepth of 31 km with an inferred age of 83 Ma.Another point is D, + 3, with an inferred age of 22Ma at about 7.5 km depth. Combination of thesetwo P-T-t points provides an exhumation of 23.5km in 61 Ma; which is an exhumation rate of 0.39km/Ma. As the depths of the other deformationphases are known (Fig. 3) their age can be esti-mated by using a first-order approach of uniformexhumation with time. As the Mulhacen Complexexperienced five phases of penetrative deforma-tion related to translations, this reasoning is cer-tainly a simplification: it should only be used asan initial guide. D, +1 and Dx+Z have an inferredage of 33 Ma and 26 Ma respectively. The age ofD x-1 is approximated at 99 Ma, using the post-D,exhumation rate for the D,-D,-, trajectory too.The 119 Ma age for initial HP/LT metamor-phism is obtained by adding 20 Ma to the age ofD x_,r which is consistent with estimates for iso-baric heating of HP/LT metamorphics frommodelling by Richardson and England (1979) andEngland and Thompson (1984). The 116 Ma ra-diametric age is thus in accordance with the ther-mal evolution of. HP/LT metamorphic crustalsegments. An alternative correlation of the 116Ma age with D,-, would result in a very slowuplift of 0.2 km/Ma between D, and D,- ,. Sucha slow uplift certainly would have erased all evi-dence of HP/LT metamorphism due to long-termrecrystallization near the thermal peak D,; thisdoes not accord with the observed partial over-printing. Another consequence of the uniform ex-humation approach is a period of about 50 Mawithout penetrative deformation between D, andD, + , This is the consequence of the rotation poleevolution, which will be discussed in the nextsection.


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The tentative 99 and 83 Ma ages for the firsttwo phases of penetrative deformation in the east-ern Internal Zone accord with well-established95-80 Ma 40Ar-39Ar ages for deformation innorthern Africa; this deformation is also char-acterized by important E-W stretching (Monie etal., 1984a, b). Stepwise heating experiments fur-ther record a thermotectonic event at 28-25 Ma(MoniC et al., 1984a, b, 1988), which accords withheating in the Betics after 26 Ma, culminating at22 Ma. The 13.8 Ma and 12.8 Ma Rb/Sr micaages indicate cooling after the Early Miocene ther-mal peak.

Rotation pole evolution and tectonics in the west-ernmost Mediterranean

Collision in the Internal Zone

Early Cretaceous rifting and rotation of Iberiaoccurred independent of the motion of Africa andNorth America (Savostin et al., 1986). Coinci-dence of the Iberian and African rotation poles inthe late Aptian has been explained by the collisionof these two plates (Savostin et al., 1986). Thetiming closely coincides with the proposed 116 +10 Ma age for initial HP/LT metamorphism andsubduction in the Internal Zone. Data from Klit-gord and Schouten (1986) and Srivastava andTapscott (1986) indicate that Iberia formed partof Africa directly from the initial rifting fromNorth America, after 123 Ma. HP/LT metamor-

5 OO

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phism in the Betics in this model would be ini-tiated by collision of Africa-Iberia with anothercontinental fragment. In both models initiation ofoceanic spreading to the west of Iberia, at about115 Ma (Boillot et al., 1989; Malod, 1989) iscoeval with collision and subduction to the east ofIberia. This timing of collision agrees with a num-ber of geological features indicative of tectonismat this time, and heavy terrigeneous sedimentationof Aptian age in the Mauritanian flysch unit innorthern Africa, which comprises turbidites with anorthern provenance (Dercourt et al., 1986)accords with collision to the north in the Betics.Widespread erosion and faulting characterizes theAptian to late Albian in the Malaguide Complex(Roep, 1980).Intracontinental thrusting in the Internal Zone

D x_ 1 and D, structures were tentatively formedat 99 Ma and 83 Ma respectively, during post-sub-duction intracontinental thrusting of segments ofvarious metamorphic grade and burial histories(Fig. 5a). Regionally consistent D,_, and D,ESE-WNW trending stretching lineations arecoincident with the 110-80 Ma motion vector inthe Betics around the combined African-Iberianrotation pole of Savostin et al. (1986) (Figs. 2 and6). Coincidence of the plate motion vector and themovement and stretching direction in a crustalscale imbricate stack suggests that thrusting atlower crustal depths is directly driven by plate

motion vector,P. 1 loo.

Fig. 6. The 110-80 Ma position of the combined rotation pole position of Africa and Iberia, motion vector around this pole forsoutheastern Spain indicated by the ESE-WNW trending bar, which coincides with similarly trending stretching lineations. Positions

of rotation poles (Ib = Iberia; A/= Afica) and continents at 80 Ma after Savostin et al. (1986), Apulia schematically indicated.


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ALPINE TEC-TONICS AND ROTATI ON POLE EVOLUTION OF IBERIA 289

convergence. The length of the 110-80 Ma platemotion vector (Savostin et al., 1986) indicatesappro~mately 600 km of motion at the latitude ofsouthern Spain. If motion has been steady throughtime this suggests 400 km of convergence between99 and 83 Ma. which constrains the maximumamount of overthrusting in the Internal Zone. Themap (Fig. 2) indicates a minimum D, overthrust-ing of 150 km of the Mulhacen Complex over theVeleta Complex parallel to the stretching linea-tion. To this estimate a few tens of kilometres ofD,_, thrusting within the Mulhacen Complex(Bakker et al., 1989) must be added. Total over-thrusting is about 200 km. The amount of over-thrusting of the Alpujarride Complex must besimultaneously considered. However, the originalrelationship between the Alpujar~de Complex andthe underlying Mulhacen Complex is disturbed bya low-angle ductile normal shear zone of Oligo-cene age (Fig. 5B). This makes it impossible toestablish the amount of Cretaceous overthrusting.Furthermore, part of the overthrusting of the Al-pujarride Complex might already have beenaccomplished during subduction. The balancingapproach is certainly a simplification, as it doesnot include the possible role of Cretaceous strike-slip motions between the Internal and ExternalZones of the Betic Cordilleras and in northernAfrica. It does indicate, however, the importanceand magnitude of westward directed overthrustingdriven by plate convergence.

As the motion vector defines the relative mo-tion between Africa-Iberia and stable Eurasia,one of the crustal segments involved in thrustingwas attached to stable Eurasia. Two models canbe envisaged. In the first model the entire InternalZone belongs to Eurasia and the External Zone toIberia. In a second mode1 the Internal Zone alsobelongs to Iberia (and thus to Africa). In bothmodels the Eurasian promontory is connected viathe Sardinia-Corsica crust with main Eurasia. Thefirst model predicts strong differential motionsbetween the allochthonous Internal Zone and Ex-ternal Zone. ESE-WNW trending stretching lin-eations in the Almagride Complex indicate in-volvement of a rifted part of the External Zone inCretaceous thrustingIn the second model the In-ternal Zone was subducted and overthrust by a

segment of Eurasia; strong differential motion hastaken place between this upper plate, presumablya Kabylian type of segment, and Iberia. In thesecond model the differential motion between theInternal and External Zones of the BeticCordilleras is much more limited. Flysch depositsto the south of the Kabylian Massifs (Bouillin etal., 1986) indicate important motion between thesemassifs and Africa. The Kabylian Massifs experi-ence E-W stretching and shearing during the 95-80 Ma period (Mom& et al., 1984a, 1988). Thisdirection fits the plate motion vector in the area atthis time, indicating that the Kabylian segmentwas involved in the Cretaceous collision. Sub-marine faulting in late Turonian to early Senoniantimes and widespread erosion in the Cenomanianto early Turonian in the Malaguide Complex(Roep, 1980) occur contemporaneously withthrusting at depth.

The 110-85 Ma tectonic regime at the northernboundary of Iberia (the Galicia Margin and thefuture Pyrenees) resulted from left-lateral motionof Africa-Iberia with respect to Eurasia (Savostinet al., 1986; Srivastava and Tapscott, 1986; Klit-gord and Schouten, 1986; Malod, 1989). The NorthPyrenean Fault can be approximated by a smallcircle around the same rotation pole as that forthe movement direction of the collisional tectonicsin the Internal Zone, indicating a coupling be-tween both motions.Pyrenean collision and tectonic quiescence in theInternal Zone

A new position of the combined African-Iberian rotation pole after 80 Ma near the Straitsof Gibraltar (Savostin et al., 1986) (Fig. 7a) re-sulted in a dramatic change in the tectonic regimein the Pyrenees and in the Betics. The formerlythinned, heated and consequently weakenedPyrenean domain was thickened in compression.Flexure modelling of the Ebro Basin at the south-ern boundary of the Pyrenees confirmed that colli-sion was intracontinental in nature (Zoetemeijer etal., 1990). Oblique convergence started in theCampanian (Puigdefabregas and Souquet, 1986)mimicking the rotation pole evolution very accu-rately. Deformation culminated in the Eocene


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290 K. DE JONG

Fig. 7. Positions of the African ( Af )-Iberian (16) combined rotation pole after Savostin et al. (1986). (a) Quadrangles indicate the80-65 Ma and circles the 65-54 Ma position close to the previous collision zone: Apulia schemetically indicated. (b) Positions of the

separated African and Iberian poles for the 54-35 Ma period. Iberia is part of the Eurasian continent at 35 Ma.

(Mattauer and Henry, 1974; Puigdefabregas andSouquet, 1986), during continuing convergence be-tween Africa and Eurasia around the Gibraltarpole. Savostin et al. (1986) indicate that motionaround this pole resulted in about 100 km ofshortening in the Pyrenees; this accords with theanalysis of the ECORS profile (Rome et al., 1989).Thermal disturbances at 60-55 Ma in the Pyreneanshear zones {Costa and Malt&i, 1988) and at thenorthwestern termination of the Iberic Cordillera(Golberg et al., 1988) coincide with initiation ofspreading in the Norwegian-Greenland Sea(Srivastava and Tapscott, 1986; Klitgord andSchouten, 1986; Savostin et al., 1986). This in-duced an additional compressional component inwestern Eurasia.

Because Iberia was attached to Africa it actedas an African promontory during collison withEurasia in the Pyrenees. It suffered relatively in-tense deformation by reactivation of Variscan andMesozoic faults (Guimera, 1984; Viallard, 1985) ina weakened lithosphere by Cretaceous extension(Zoetemeijer et al., 1990). Convergence resulted inlimited riot-later~ motion of main Iberia withrespect to the Iberic Cordillera and the Ebro Basin(Malod, 1982) and left-lateral motion in the Cata-lan Coastal Range (Guimera, 1984). A subhori-zontal NW-directed main compression directionu1 in the Catalan Coastal Range (Guimera, 1984)coincides with the motion direction around theGibraltar poie.

During the Pyrenean collision the Internal Zonewas tectonically relatively quiet, as is evident fromthe absence of penetrative deformation for about50 Ma, between 83 Ma and 33 Ma. This is in closeagreement with the geochronological evolution ofthe Kabylian Massifs in northern Algeria. Thelocation of the African-Iberian rotation pole nearthe Cretaceous Betic collision zone (Fig. 7a) in-bibited large-scale motions in this zone after 80Ma. Minor tectonics is documented in the Exter-nal Zone of the Betics in the form of relative upliftbetween 68 Ma and 60 Ma; this probably resultedfrom compression caused by African-Eurasianconvergence (Kenter et al., 1990).

After 54 Ma the African rotation pole wasseparated from the Iberian pole and shifted north-ward (Savostin et al., 1986) (Fig. 7b). Shiftingmight be partly explained by the Pyrenean colli-sion, forcing the African plate to pivot around adifferent pole. The effect of this new rotation poleposition is the starting of limited differential mo-tion between Africa and Iberia, resulting in thetectonism which has been well documented in theBetic Cordilleras. The External Zone records asecond period of uplift at about 50 Ma (Kenter etal., 1990). The Malaguide Complex shows an ero-sional contact between the Maastrichtian and theEarly Eocene (Roep, 1980). Because the metamor-phic zones of southern Spain and northern Africado not record penetrative deformation and meta-morphism, this tectonic activity is probably of


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ALPINE TECTONICS AND ROTATI ON POLE EVOLUTION OF IBERIA 291

minor importance. Northerly to north-northeast-erly compression is detected in the late Eocene inthe Catalan Coastal Range (Guimera, 1984) andthe External Betics (De Ruig, this issue) and inother areas in Spain (Letouzey, 1986; Bergerat,1987). The main compression direction is coinci-dent with the motion vector around the 54-35 MaAfrican rotation pole of Savostin et al. (1986)(Fig. 7b). The mountain ranges of northern Africaalso experienced compression during this episode(Letouzey, 1986; Dercourt et al., 1986) presuma-bly connected with the Pyrenean collision. ThePyrenean collision ended at about 35 Ma, afterwhich no independent lberian rotation pole maybe detected (Savostin et al., 1986). Klitgord andSchouten (1986) concluded that the finalamalgamation of Iberia with Eurasia occurred atabout 30 Ma.Extensional tectonics

Savostin et al. (1986) demonstrate a dramaticshift of the African rotation pole to a new positionin the South Atlantic at 35 Ma as a consequenceof 40% decrease in spreading rate between theAfrican and North American plates; no changeoccurred in Eurasia-North America spreading.This kinematic pattern is, however, not supportedby the central Atlantic magnetic anomaly pattern(Klitgord and Schouten, 1986) or by NorthAtlantic and Arctic data (Srivastava and Tapscott,1986). Klitgord and Schouten (1986) indicate ashift of the African-Eurasian plate boundary fromthe northern side of Iberia to the southern side ofit, separating Iberia from Africa by the Azores-Gibraltar fracture zone after 30 Ma. The establish-ment of a new plate boundary between Iberia andAfrica accords with timing of extensional defor-mation and heating in the Internal Zone andnorthern Africa. Regionally consistent southeast-erly slip of the hangingwall towards ultramaficrocks in the western Betics (Tubia and Cuevas,1986) and the Rif (Saddiqi et al., 1988) suggests aSE-NW extensional component on the plateboundary. Extension is further attested by an EarlyMiocene basaltic dike-swarm (Torres-Roldan etal., 1986). The timing of extension and magmatismaccords well with the tectonic evolution of the

Gulf of Lion and the Balearic Basin, which openedin Oligocene to Aquitanian times (Alvarez et al.,1974; Rehault et al., 1984). Deformation in thisarea spread southward with time (Mauffret andGennesseaux, 1989). Extension resulted in pro-gressive crustal thinning in northeastern Spain to-wards the Balearic Basin (Banda, 1987). Strongcrustal thinning in the eastern part of the InternalZone (Banda and Ansorge, 1980) is also attributedto this Oligo-Aquitanian crustal thinning and ex-tension. Extensional deformation to the east ofIberia links the African-Eurasian plate boundarywith the western European rift system (Bresse-Rhine graben) which experienced extension up tothe latest Oligocene (Bergerat, 1987). The shift ofthe African rotation pole at 35 Ma accords wellwith renewal of deformation in the Betics andnorthern Africa. However, the motion around thispole (Savostin et al., 1986) generates northwesterlycompression between Africa and Eurasia andhence cannot explain the observed extension.

The 20 Ma position of the African rotationpole to the northwest of Portugal (Savostin et al.,1986) resulted in cessation of extensional deforma-tion and induced a northeasterly relative motionof Africa with respect to Eurasia, resulting inNE-SW directed compression (Letouzey, 1986;Bergerat, 1987). Compression caused thickening ofthe previously thinned and weakened crustal do-main to the south of Iberia, this being demon-strated by NNE-directed thrusting in the Mulha-ten Complex (Bakker et al., 1989) and reactivationof the original contact between the Alpujarrideand Mulhacen Complexes. The Alpujarride Com-plex has moved northwards on a shear zone withrespect to the underlying Mulhacen Complex (Plattet al., 1983) which already contains the imprintsof the 22 Ma Dx+3 thermal peak. Behrmann (1984)concluded on the basis of palaeostress and strainrate estimates that ductile deformation in thisshear zone lasted 4 Ma. Consequently, ductilethrusting stopped during the Burdigalian. Thisaccords with the superposition of the MalaguideComplex onto the already quite well developedAlpujarride Complex around the Aquitanian-Burdigalian boundary (Makel, 1981) and withsealing of thrusts in the Malaguide Complex byBurdigalian sediments (MacGillavry et al., 1963;


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292 K DE JONG

Torres-Roldan et al., 1986). Burdigalian sedimen-tation in large parts of the Internal Zone occurredin a tectonically quiet environment (Volk, 1967).

The compression direction after 20 Ma enablesearly left-lateral motion on ENE-WSW to NE-SW trending strike-slip faults crossing the BeticZone. Middle Miocene strike-slip motion has beendescribed by Sanz de Galdeano et al. (1985) andby Bon et al. (1989). The 10 Ma to present rota-tion pole position of Africa, to the west ofGibraltar (Savostin et al., 1986) enables thenorthwesterly to northerly compression detectedby Montenat et al. (1987) in the Betics. The posi-tion of the main compression axes in the obtuseangle of a set of NE-SW and NNE-SSW trend-ing strike-slip faults (Montenat et al., 1987) pointsto a reactivation of an earlier fault system. Theposition of the rotation pole relatively close to thestrike-slip fault system only allows limited dis-placements. About 20 km of left-lateral slip hasbeen argued by Veeken (1983) for the NNE-trending Palomares Fault. Both phases of strike-slip motion juxtapose crustal segments with differ-ent Moho depths, the latter inherited from lateOligocene to Aquitanian extension. Present-daycompression and intraplate seismicity (Udias etal., 1976) is localized in the thinned Alboran Seaand southern Spain.Conclusions

The tectonic evolution of the Alpine collisionbelts bordering Iberia is in accordance withkinematics of the Iberian and African plates andtheir rotation pole evolution.

Subduction and HP/LT metamorphism in theBetics at 116 + 10 Ma coincide with initiation ofseafloor spreading at the northwestern margin ofIberia and the Bay of Biscay. This phase alsocoincides with heavy terrigeneous sedimentationin the Mauritanian flysch in northern Africa andthe tectosedimentary evolution of the MalaguideComplex.

A period of intracontinental thrusting between99 Ma and 83 Ma in the Internal Zone of theBetics occurred during continuing spreading to thewest of Iberia and progressive opening of the Bayof Biscay. Regionally consistent ESE-WNWtrending stretching lineations in the Internal Zone,

which were formed during thrusting at lowercrustal levels, are coincident with the motion di-rection of Africa-Iberia with respect to Eurasia.At least 200 km of thrusting in the Internal Zonewas driven by the convergence of Africa-Iberiawith respect to Eurasia. Strike-slip motion ofAfrica-Iberia with respect to Eurasia along theNorth Pyrenean Fault took place around the samerotation pole as thrusting in the Betics.

At 80 Ma, shifting of the African-Iberian rota-tion pole to a position near the Betic collisionzone resulted in cessation of penetrative deforma-tion in the Internal Zone and northern Africa.Collision of Africa-Iberia with Eurasia was trans-ferred to the Pyrenees. Iberia acted as an Africanpromontory by motion around the rotation polenear Gibraltar. Pyrenean collision culminated at60-55 Ma by an additional compression generatedby seafloor spreading in the Norwegian-Green-land Sea.

After completion of the Pyrenean collision at35-30 Ma, penetrative deformation and heatingagain took place in the Internal Zone and north-ern Africa by establishment of the Azores-Gibraltar fracture zone, which interacted with theCretaceous collision system. The new plateboundary was connected via the Balearic rift sys-tem with the western European rift. Deformationwas initiated by late Oligocene crustal and litho-spheric mantle extension, resulting in EarlyMiocene emplacement of ultramafic rocks withassociated heating.

Compression after 20 Ma was initiated by mo-tion around a rotation pole to the west of Iberia.Crustal thickening was localized in the previouslyextended and heated and consequently weakenedInternal Zone. This process resulted in overthrust-ing in the Internal Zone, which was completedbefore the end of the Burdigalian. Middle andLate Miocene strike-slip motion was initiated dur-ing continuing convergence and juxtaposed crustalsegments of differing Moho depths, the latter hav-ing been inherited from late Oligocene to EarlyMiocene extension.Acknowledgements

I thank Sierd Cloetingh for constructive criti-cism on the manuscript and Henk Helmers for


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ALPINE TECTONKS AND ROTATION POLE EVOLUTION OF IBERIA 293

placing the strain data on plagioclase porphyro-clasts at my disposal. Jacques Malod, PaulAndriessen and Reini Zoetemeier are thanked forproviding preprints. The technical assistance ofFred Kievits during the preparation of the figuresis aiso acknowledged. Part of the field work wasfinanced by a grant from the StichtingMolengraaff-fonds.ReferencesAlbarede, F. and Michard-Vitrac, A., 1978. Age and signifi-

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