Alpine tectonics and rotation pole evolution of Iberia

Tectonophysics, 184 (1990) 219-296

Elsevier Science Publishers B.V., Amsterdam

279

Alpine tectonics and rotation pole evolution of Iberia

Koen 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)

ABSTRACT

De 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 plate

are provided by P-T-t data and regionally consistent stretching lineations from the metamorphic parts of the Betic

Cordilleras.

High-pressure low-temperature metamorphism in the Betic Cordilleras resulted from continent-continent collision which

caused subduction to a maximum depth of 37 km. A preliminary 116 + 10 Ma radiometric age for this event corresponds to

the initiation of seafloor spreading to the west of Iberia which lasted until about 80 Ma. Intracontinental thrusting in the

Betics between 99 Ma and 83 Ma took place after subduction ended. E-W to ESE-WNW trending stretching lineations

indicate the direction of thrusting, which resulted in extensional strains of 200-600%. The timing of thrusting in the Betics

coincides with a 95-80 Ma tectonic phase in northern Africa, during which E-W stretching lineations were formed. The

stretching Iineations are coincident with the 110-80 Ma motion vector of Africa-Iberia with respect to Eurasia. Thrusting in

the Betics and deformation in northern Africa was driven by convergence of Africa-Iberia and Eurasia. Cretaceous

deformation is further recorded by terrigeneous sedimentation in the Mauritanian Flysch and by the tectosedimentary

evolution of the Malaguide Complex. Crustal thinning, magmatism and metamorphism in the Pyrenees during the 110-85 Ma

period is governed by a left-lateral strike-slip of Africa-Iberia with respect to Eurasia around the same rotation pole as

thrusting in the Betics.

During the 80-54 Ma period the rotation pole was situated west of Gibraltar, near the previous active collision zone. This

inhibited 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 the

Campanian on wards, oblique convergence took place around the combined Gibraltar rotation pole. Deformation culminated

in the late Eocene, corresponding to spreading in the Norwegian-Greenland Sea at 55 Ma which induced an additional

compression in western Eurasia. During the Pyrenean collision, high-pressure metamorphic rocks in the Betic Cordilleras were

exhumed and they cooled substantially. The cooling trend was disturbed by Oligocene extensional deformation and

introduction of a transient heat source, which correlates with the mantle being uplifted during extension. Heating culminated

at the Oligocene-Miocene boundary in the Betic Cordilleras and in northern Africa. This evolution agrees with the

development of a plate boundary between Iberia and Africa at 30 Ma, after completion of the Pyrenean collision. The new

plate 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 two

phases of wrenching, juxtaposing crustal segments with different Moho depths inherited from the late Oligocene to Early

Miocene extension.

Introduction

The Iberian peninsula (Fig. 1) is bordered by

two Alpine foldbelts, the Pyrenees to the north

and the Betic Cordilleras to the south. These belts

separate the Iberian plate from, respectively, the

Eurasian 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 et

al., 1986). Until now the Betic Cordilleras has

most often been regarded as a Tertiary orogen,

mainly on the basis of important Tertiary defor-

mation in the non-metamorphic parts (Rondeel

and Simon, 1974; De Smet, 1984). A Mesozoic age

for the early deformation has, however, been sug-

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

280

gested by Kampschuur and Rondeel(l975) owing to the Mesozoic age of the flysch deposits in the western Betics. New data discussed in this paper also suggest important Cretaceous tectonics in the metamorphic Internal Zone of the Betic Cordilleras. Ceochronological studies in the Al- pine collision belt of northern Africa (Monie et al., 1984a; 1988) show a tectonic evolution which is comparable to that of the Betics-Cretaceous metamo~~c ages and an important Tertiary resetting. This paper aims at tying the new tectonic model for the Betic Zone and the thermotectonic evolution of the northern African belt to the well- constrained tectonic evolution of the Pyrenees. The tectonic evolution of the erogenic belts bordering the Iberian plate will be shown to be consistent with the Cretaceous and Tertiary rotation pole and kinematic evolution of the Iberian and African plates discussed by Savostin et al.

K. DE JONG

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

Regionally consistent stretching lineations which were formed during early Alpine thrusting at lower crustal levels are a salient feature of the tectonic evolution of the Internal Zone. They coincide with the mid-Cretaceous motion vector of

Savostin et al. (1986) of the African plate (including Iberia at that time) with respect to Eurasia. Because stretching lineations appro~mate the movement direction in shear zones (Esscher and Watterson, 1974) they probably also trace plate motion directions. A relationship between the direction of thrusting and plate motion has been suggested for an number of orogens (Shackleton and Ries, 1984), including the Alps (Baird and Dewey, 1986; Choukroune et al., 1986). During the later stages of the tectonic evolution of the arcuate western Alps, radial thrusting occurred

IBERIAN MESETA

GULF DE LION

Infernal Zones of the Betlc Cordilleras and Rlf

;:

D IAbne metamorphic rocks)

a

Fig. 1. Sketch map of the westernmost Mediterranean area (modified after Ricou et al., 1986) showing the major Alpine structural

provinces. The eastern Betic Cordilleras of southern Spain are delineated.

ALPINE TECTONICS AND ROTATION POLE EVOLUTION OF IBERIA 281

(Choukroune et al., 1986) at higher crustal levels,

this thrusting clearly bearing no relationship to

plate motion vectors. Relatively large finite dis-

placements and rotations during such a stage will

disturb the original pattern of older stretching

lineations formed at deeper levels. The eastern

part of the Betic Cordilleras does not demonstrate

an arcuate form, and therefore no pervasive re-

orientation of older structures is to be expected.

Palaeostress analyses in stable forelands do not

usually suffer the disadvantage of reorientation, as

finite strain is in general small. A clear relation-

ship between (successive) palaeostress directions

and plate motion vectors is therefore recorded in

the Alpine foreland (Letouzey, 1986; Bergerat,

1987). However, in Iberia the Mesozoic palaeos-

tress directions do not mimic the plate vector very

accurately (Malod, 1989). This is partly the result

of reorientation and heterogeneities induced by

pre-existing faults. Therefore, the regionally con-

sistent stretching lineations in the Internal Zone

are considered as an important constraint in the

early kinematic evolution of the Iberian plate.

Motions around different rotation pole positions

during orogeny will, due to overprinting and re-

orientation, not be recorded by successive genera-

tions of stretching lineations. Shifting of rotation

poles has, however, a marked effect on tectonics in

metamorphic belts, as will be discussed later.

Evolution of the tectonic zones bordering Iberia

The Alpine collision belts bordering Iberia are

characterized by Jurassic to Early Cretaceous ex-

tensional deformation and related strike-slip de-

formation. A Middle to Late Jurassic strike-slip

fault between Iberia-Africa is indicated by plate

reconstructions (Savostin et al., 1986; Klitgord

and Schouten, 1986). The occurrence of a frag-

ment of an ophiolite sequence of Late Jurassic age

in northern Africa (Bouillin et al., 1977) accords

with these reconstructions. Continuing motion into

the Cretaceous is indicated by flysch deposits

culminating in Aptian-Albian times in the Flysch

Domain (Bouillin et al., 1986). The non-metamor-

phic External Zones of Iberia and northern Africa

are palaeogeographically unrelated (Bouillin et al.,

1986) this also indicating their initial separation.

In the External Zone of the Betic Cordilleras an

algal platform broke up at the Middle to Late

Jurassic boundary (Geel, 1979) resulting in strong

palaeogeographical differentiation (Hermes, 1978).

An extensional tectonic regime is indicated by

pillow basalt intrusions in the Sub-Betic (Hermes,

1978; De Smet, 1984). Important hiatuses, turbi-

dite deposits and the occurrence of Middle Jurassic

lithoclasts in Albian-Aptian marls (Hermes, 1978)

indicate important vertical motions continuing into

the Cretaceous. Basaltic intrusion in the Internal

Zone 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 the

future 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 rifting

possibly occurred; important rifting started in the

Berriasian to earliest Valanginian (144-140 Ma,

Boillot et al., 1989). The end of emplacement of

ultramafic rocks by ductile normal faulting has

been 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 of

seafloor spreading (Boillot and Malod, 1988; Boil-

lot et al., 1989; Malod, 1989). Opening of the Bay

of Biscay occurred between the Aptian and

Campanian 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 and

Tapscott, 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 and

Michard-Vitrac, 1978; Montigny et al., 1986).

During this period the North Pyrenean Fault zone

was characterized by high heat flow in response to

crustal thinning related to strike-slip tectonics

(Choukroune and Mattauer, 1978; Vielzeuf and

Kornprobst, 1984; Golberg et al., 1986). Early

Cretaceous metamorphic ages are also well re-

corded by the 4oAr-39Ar stepwise heating method

2x2 K. DE JDNG

of samples from mylonite zones in the eastern

Pyrenees (110-100 Ma and 90 Ma, Costa and

Maluski 1988) and in strike-slip basins at the

northwestern termination of the Iberic Cordillera

(100 Ma, Golberg et al., 1988).

The extensional regime in the Pyrenees changed

to compression in the latest Cretaceous (Vielzeuf

and Kornprobst, 1984); oblique convergence

started in the Campanian (PuigdefBbregas and

Souquet, 1986). Strike-slip ceased dr~atically

during the middle Eocene, when major thrusts

were developed parallel to the mylonite zones

(Soula et al., 1986). The northern Spanish passive

margin was converted into an active margin in the

Paleocene-Eocene interval as a result of plate

convergence (Boillot and Malod, 1988). During

convergence, Variscan and Early Cretaceous faults

were reactivated (Soula et al., 1986; McCaig and

Miller, 1986; Majoor, 1988). Radiometric ages in

mylonite zones indicate a latest Cretaceous to

middle Eocene age for reactivation (McCaig and

Miller, 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 and

Souquet, 1986). Piggy-back thrusting in the central

southern Pyrenees migrated southward with time

(Williams, 1985), concomitant with the progressive

southward development of molasse basins (Mat-

tauer and Henry, 1974) and their incorporation in

subsequently formed thrust units (Puigdef~bregas

and Souquet, 1986). Flexure modelling ildb shown

that the Ebro foreland basin formed as a result of

the Pyrenees, Catalan Coastal Range and Iberic

Cordillera 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 De

Weerd, 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 Jurassic

to Early Cretaceous period, which accord well

with phases of spreading in the North Atlantic

(Malod, 1989). Late Eocene to late Oligocene

northward compression reactivated Variscan and

Mesozoic faults (GuimerB, 1984; Viallard, 1985).

Raising of geotherms during extension caused ef-

fective weakening of the lithosphere, which is also

a si~ificant factor in the localization of in-

tracontinental Tertiary compression (Zoetemeijer

et al., 1990).

The External Zone of the Betic Cordilleras

records only relatively small increments of the

Pyrenean 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 to

Late Miocene boundary (Simon, 1987) during

northward thin-skinned tectonics with intermit-

tant strike-slip deformation (De Ruig et al., 1987).

Compression also began in the Middle Miocene in

the Tell and External Rif of northern Africa.

In the following I present indications of an

important phase of tectonism in the Internal Zone

of the Betic Cordilleras between the Barremian

and Campanian.

Tectonic evolution of the Internal Zone of the

Betic Cordilleras

Regionaf scale structure

The Internal Zone occurs to the south of the

External Zone, which constitutes the margin of

Iberia characterized by Mesozoic rifting (Hermes,

1978). These two zones are presently separated by

a strike-slip fault of Miocene age (Hermes, 1978;

Le Blanc and Olivier, 1984; De Smet. 1984). The

Triassic stratigraphy of the overthrust units in the

Internal Zone bears no resemblance to the Triassic

stratigraphy of the External Zone (Simon, 1987).

The Internal Zone can thus be considered as alloc-

hthonous to Iberia. (Very) low grade Triassic

metarno~~c rocks of the Almagride Complex

occur in windows below the Alpujarride Complex

of the Internal Zone (Fig. 2); they show striking

similarities 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 during

Mesozoic rifting. The Miocene strike-slip fault

between the Betic Zone and the External Zone can

not therefore be interpreted as a plate boundary,

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284 K. DE JONG

but as a later structure of relatively minor impor.

tance. The Internal Zone consists of four stacked

crustal segments. The lowest segment, the Veleta Complex, is characterized by low-pressure, low- temperature (LP/LT) metamorphism (Puga and Diaz de Federico; 1978), for which recently a pre-Alpine age has been suggested (Gomez- Pugnaire and Franz, 1989). This complex has been overthrust by the Mulhacen Complex, which bears evidence of early Alpine high-pressure, low-temperature (HP/LT) metamorphism partially over- printed by medium-grade metamorphism (Gomez -Pugnaire and Femandez-Soler, 1987; Bakker et al., 1989). The Alpujarride Complex occurs on top of the Mulhacen Complex and is also characterized by early H P/LT met~o~~srn (Bakker et al., 1989; GoffC et al., 1989). In the western Betics the Alpujarride Complex experienced high- temperature metamo~~sm at high to low pressure related to emplacement of ultramafic rocks (Westerhof, 1977) of Early Miocene age (Priem et al.. 1979; Zindler et al., 1983). The uppermost tectonic unit, the Malaguide (Ghoma~de in northern Africa) Complex, is almost entirely non-metamorphic (Egeler and Simon, 1969). Condensed, but continuous, Mesozoic and Paleogene stratigraphy (Roep, 1980) indicates that this crustal domain has always retained near the crustal surface. The Dot-sale calcaire and Pre-dorsalian Zone rep- resent the margin of this Malaguide/Ghomaride domain with dominant Triassic and Jurassic shelf and slope sedimentation (Bouillin et al., 1986).

3,

10

09

08

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05

04

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T(C”)

Fig. 3. P-T-r paths of the Mulhacen Complex (light stipple)

and the Almanzora Unit (dark stipple). Tourmahne K/Ar.ages

for D, and D,_t are indicated. Ages (Ma) of deformation

phases 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 (Newton

and 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 tectonic evolution of the Mulhacen Complex and the Al- manzora Unit of the Alpujarride Complex has been reconstructed by Bakker et al. (1989) by relating deformation and metamorphism and P-T determinations by microprobe analyses.

Both complexes are characterized by initial HP/LT metamorphism, indicating the disturbance of the pre-existing pattern of isotherms by subduction. The maximum metamorphic pres-

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

ALPINE TECTONICS AND ROTATION POLE EVOLUTION OF IBERIA

570” C in the Mulhacen Complex and 450 o C in

the Almanzora Unit (Fig. 3) (Bakker et al., 1989).

D, effectively transposed all earlier fabrics. In

the Almanzora Unit D,_, deformation fabrics

were only left as internal fabrics in glaucophane

and crossite crystals. In the Mulhacen Complex

penetrative D,._, fabrics were left unaffected only

locally in glaucophane schist facies amphibolites,

local conglomerate bodies and in the core of a 5.5

km3 gneiss body. D, _ 1 is characterized by E-W to

ESE-WNW trending stretching lineations (Fig. 2)

on flat-lying foliations. Strain analyses on con-

glomerate pebbles indicates extensional strains of

lOO-300% (Fig. 4A). Borradaile (1976) inferred

ESE-WNW extensional strains in excess of 150%

in the gneisses. Top-to-the-west shear is indicated

by asymmetric tails at the extremities of K-felds-

par porphyroclasts. Ductile D,_ 1 deformation was

the result of thrusting within the Mulhacen Com-

plex (Bakker et al., 1989). D, is also characterized

by important regionally consistent E-W to ESE-

WNW stretching (Fig. 2) paralleling axes of lo-

cally pronounced sheath-like folds. Plagioclase

porphyroclasts in amphibolites record extensional

strains of 200-600% (Fig. 4B) which, however,

might contain an unknown D,_, component. A

minimum D, stretching amount is provided by the

rotation of syn-D, garnets. Rotation angles of

90-110” indicate a shear strain of 3.5 using the

Rosenfeld (1970) equation. If this rotation is

B

285

formed during plane strain, the shear strain indt-

cates 250% extension. Vissers (1989) arrived at a

similar strain estimate. Because the contact be-

tween the internal garnet fabric and the external

fabric is lost most often during post-blastesis D,

deformation and garnets have been boundinaged,

this amount is a minimum estimate. Strain analyses

on pebbles in quartzites with a penetrative D,

fabric in the Alpujarride Complex indicate 300-

600% extension (Fig. 4C). Similarity of main phase

deformation fabrics in both complexes combined

with comparable amounts and directions of

stretching directions indicate a similar tectonic

history. Strain in the Alpujarride Complex is

much more heterogeneously developed, the core of

the 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 the

underlying Mulhacen Complex and heating to-

gether with loading by overlying crustal segments

during D,. In the last stages of D, the Mulhacen

Complex overthrust the Veleta Complex (Fig. 5B),

the latter never having experienced very deep

tectonic burial. The contact zone is characterized

by mylonites with E-W trending stretching linea-

tions (Fig. 2). Preferred orientations of quartz-c-

axes indicate top-to-the-west shear (De Jong, in

prep.). 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,_ 1 quartz pebbles, Mulhacen Complex, location (c) in Fig. 2, central Sierra de 10s Filabres. (B) D, elongated plagioclase porphyroclasts

in amphibolites, Mulhacen Complex, location (d) in Fig. 2, eastern Sierra de 10s Filabres. (C) D, quartz pebbles, high strain area

indicated by crosses, low strain area by dots, Alpujarride Complex, location (i) in Fig. 2, central Sierra de las Estancias; the single

square displays one determination from the Almanzora Unit near location (8).

2X6 K. DE JONG

A LD, CONFIGURATION 1 AU MC

i- T--r 80-85 Ma

6 Lb+, CONFIGURATION c T--r

‘i

w \

33 Ma

E

0

Fig. 5. (A) Crustal configuration during the 85-80 Ma phase of crustal scale imbrication a peak thermal conditions during D, (insets

arc 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, which

record top-to-the-west shearing by asymmetric

pull-apart structures.

After D,, decompression and cooling dominates

the thermal evolution, in the absence of any

penetrative deformation. The form of the P-T-t

trajectory precludes rapid exhumation of the In-

ternal Zone. Continued thermal equilibration to-

wards a new steady-state isotherm generates an

increasing apparent geothermal gradient, which

reaches about 35O/km. Erosional unroofing pre-

sumably played an important role during this

trajectory.

The Almanzora Unit was placed on top of the

Mulhacen Complex along a low-angle ductile nor-

mal shear zone during D,, ,, which represents a

phase of heterogeneous crustal thinning and ex-

tension (Bakker et al., 1989). The latter authors

estimated the excision of a 6 km thick crustal

section along a single shear zone during eastward

slip of the hangingwall. Thermal consequences of

extension were retarded; temperature increase

started during Dx+* and culminated during the

D X+3 second thermal peak at 525 o C (Bakker et

al., 1989). Retardation of heating at a particular

crustal level with respect to the timing of exten-

sion is also shown by thermomechanical models of

crustal extension (Crough and Thompson, 1976;

Thompson, 1981; Moretti and Froideveaux, 1986).

Crustal thickening by large-scale S-verging D, +Z


folding and associated thrusting enhanced this.

The second thermal peak may indicate the effect

of magma addition (Thompson, 1981) or introduc-

tion of magmatic fluids (England and Thompson,

1984) at higher crustal levels. The large amount of

heating, about 100 o C (Fig. 3) indicates the intro-

duction of a transient heat source in the Internal

Zone; this was the consequence of crustal and

lithospheric mantle extension and resulted in em-

placement of ultramafic rocks in the western Bet-

its (Platt, 1987; Bakker et al., 1989; Doblas and

Oyarzun, 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 of

similar rocks in the Rif in Morocco (Ben Othman

et al., 1984). The Dx+3 thermal peak in the eastern

Internal Zone is consequently an Early Miocene

feature. The regional importance of this event is

indicated by widespread resetting of radioisotope

systems at the Oligocene-Miocene boundary in

the Kabylian Massifs in northern Algeria and the

External Rif (MoniC et al., 1984a, b, 1988).

Timing of tectonic events in the Mulhacen Com-

plex

In order to constrain the tectonic evolution of

the Betic Zone we need to know the ages of the

various tectonometamorphic phases. However,

only very limited radiometric data on metamor-

phic minerals are available. For the Mulhacen

Complex an average muscovite Rb/Sr age of 13.8

Ma has been calculated from data reported by

Priem et al. (1966) and Andriessen et al. (1989).

The latter authors also report a 12.8 Ma biotite

Rb/Sr age. In addition, in the context of a feasi-

bility study of K/Ar dating of tourmalines three

ages of 80-85 jI 8 Ma and one age of 116 f 10 Ma

have been reported (Andriessen et al., 1989). The

suggested blocking temperature for tourmaline of

above 600” C (Andriessen et al., 1989) well ex-

ceeds the maximum metamorphic temperature in

the Mulhacen Complex. These ages can therefore

be interpreted as metamorphic crystallization ages.

The 80-85 Ma ages (averaging 83 Ma) are ob-

tained from gneisses with a D, mylonite fabric, in

which synkinematic growth of tourmaline has oc-

curred. This suggests an age of 83 Ma for D,. For

the 116 Ma age, excess Ar might be put forward as

an explanation as this component has a widespread

occurrence (Hebeda et al., 1980); in addition, this

is an explanation which is difficult to disprove.

Alternatively, this age can reflect pre-D, ~, or D, _. ,

metamorphism. It will be shown that the 116 Ma

age is actually in accordance with the time scale

for the establishment of the Mulhacen Complex

P-T-t path, and can be geologically relevant.

A point which constrains the P-T-t path of

the Mulhacen Complex is D,, which occurs at a

depth of 31 km with an inferred age of 83 Ma.

Another point is D, + 3, with an inferred age of 22

Ma at about 7.5 km depth. Combination of these

two P-T-t points provides an exhumation of 23.5

km in 61 Ma; which is an exhumation rate of 0.39

km/Ma. As the depths of the other deformation

phases are known (Fig. 3) their age can be esti-

mated by using a first-order approach of uniform

exhumation with time. As the Mulhacen Complex

experienced five phases of penetrative deforma-

tion related to translations, this reasoning is cer-

tainly a simplification: it should only be used as

an initial guide. D, + 1 and Dx+Z have an inferred

age of 33 Ma and 26 Ma respectively. The age of

D 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 of

D x_,r which is consistent with estimates for iso-

baric heating of HP/LT metamorphics from

modelling by Richardson and England (1979) and

England and Thompson (1984). The 116 Ma ra-

diametric age is thus in accordance with the ther-

mal evolution of. HP/LT metamorphic crustal

segments. An alternative correlation of the 116

Ma age with D,-, would result in a very slow

uplift of 0.2 km/Ma between D, and D,- ,. Such

a slow uplift certainly would have erased all evi-

dence of HP/LT metamorphism due to long-term

recrystallization near the thermal peak D,; this

does not accord with the observed partial over-

printing. Another consequence of the uniform ex-

humation approach is a period of about 50 Ma

without penetrative deformation between D, and

D, + , This is the consequence of the rotation pole

evolution, which will be discussed in the next

section.

288

The tentative 99 and 83 Ma ages for the first

two phases of penetrative deformation in the east-

ern Internal Zone accord with well-established

95-80 Ma 40Ar-39Ar ages for deformation in

northern Africa; this deformation is also char-

acterized by important E-W stretching (Monie et

al., 1984a, b). Stepwise heating experiments fur-

ther record a thermotectonic event at 28-25 Ma

(MoniC et al., 1984a, b, 1988), which accords with

heating in the Betics after 26 Ma, culminating at

22 Ma. The 13.8 Ma and 12.8 Ma Rb/Sr mica

ages 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 Iberia

occurred independent of the motion of Africa and

North America (Savostin et al., 1986). Coinci-

dence of the Iberian and African rotation poles in

the late Aptian has been explained by the collision

of these two plates (Savostin et al., 1986). The

timing closely coincides with the proposed 116 +

10 Ma age for initial HP/LT metamorphism and

subduction in the Internal Zone. Data from Klit-

gord and Schouten (1986) and Srivastava and

Tapscott (1986) indicate that Iberia formed part

of Africa directly from the initial rifting from

North America, after 123 Ma. HP/LT metamor-

‘5”’ ’ ‘OO’ ”

K DE JONG

phism in the Betics in this model would be ini-

tiated by collision of Africa-Iberia with another

continental fragment. In both models initiation of

oceanic spreading to the west of Iberia, at about

115 Ma (Boillot et al., 1989; Malod, 1989) is

coeval with collision and subduction to the east of

Iberia. This timing of collision agrees with a num-

ber of geological features indicative of tectonism

at this time, and heavy terrigeneous sedimentation

of Aptian age in the Mauritanian flysch unit in

northern Africa, which comprises turbidites with a

northern provenance (Dercourt et al., 1986)

accords with collision to the north in the Betics.

Widespread erosion and faulting characterizes the

Aptian to late Albian in the Malaguide Complex

(Roep, 1980).

Intracontinental thrusting in the Internal Zone

D x_ 1 and D, structures were tentatively formed

at 99 Ma and 83 Ma respectively, during post-sub-

duction intracontinental thrusting of segments of

various metamorphic grade and burial histories

(Fig. 5a). Regionally consistent D,_, and D,

ESE-WNW trending stretching lineations are

coincident with the 110-80 Ma motion vector in

the Betics around the combined African-Iberian

rotation pole of Savostin et al. (1986) (Figs. 2 and

6). Coincidence of the plate motion vector and the

movement and stretching direction in a crustal

scale imbricate stack suggests that thrusting at

lower 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 for

southeastern 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.

ALPINE TEC-TONICS ‘AND ROTATION POLE EVOLUTION OF IBERIA 289

convergence. The length of the 110-80 Ma plate

motion vector (Savostin et al., 1986) indicates

appro~mately 600 km of motion at the latitude of

southern Spain. If motion has been steady through

time this suggests 400 km of convergence between

99 and 83 Ma. which constrains the maximum

amount of overthrusting in the Internal Zone. The

map (Fig. 2) indicates a minimum D, overthrust-

ing of 150 km of the Mulhacen Complex over the

Veleta Complex parallel to the stretching linea-

tion. To this estimate a few tens of kilometres of

D,_, 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 be

simultaneously considered. However, the original

relationship between the Alpujar~de Complex and

the underlying Mulhacen Complex is disturbed by

a low-angle ductile normal shear zone of Oligo-

cene age (Fig. 5B). This makes it impossible to

establish the amount of Cretaceous overthrusting.

Furthermore, part of the overthrusting of the Al-

pujarride Complex might already have been

accomplished during subduction. The balancing

approach is certainly a simplification, as it does

not include the possible role of Cretaceous strike-

slip motions between the Internal and External

Zones of the Betic Cordilleras and in northern

Africa. It does indicate, however, the importance

and magnitude of westward directed overthrusting

driven 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 thrusting

was attached to stable Eurasia. Two models can

be envisaged. In the first model the entire Internal

Zone belongs to Eurasia and the External Zone to

Iberia. In a second mode1 the Internal Zone also

belongs to Iberia (and thus to Africa). In both

models the Eurasian promontory is connected via

the Sardinia-Corsica crust with main Eurasia. The

first model predicts strong differential motions

between 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 in

Cretaceous thrustingIn the second model the In-

ternal Zone was subducted and overthrust by a

segment of Eurasia; strong differential motion has

taken place between this upper plate, presumably

a Kabylian type of segment, and Iberia. In the

second model the differential motion between the

Internal and External Zones of the Betic

Cordilleras is much more limited. Flysch deposits

to the south of the Kabylian Massifs (Bouillin et

al., 1986) indicate important motion between these

massifs and Africa. The Kabylian Massifs experi-

ence E-W stretching and shearing during the 95-

80 Ma period (Mom& et al., 1984a, 1988). This

direction fits the plate motion vector in the area at

this time, indicating that the Kabylian segment

was involved in the Cretaceous collision. Sub-

marine faulting in late Turonian to early Senonian

times and widespread erosion in the Cenomanian

to early Turonian in the Malaguide Complex

(Roep, 1980) occur contemporaneously with

thrusting at depth.

The 110-85 Ma tectonic regime at the northern

boundary of Iberia (the Galicia Margin and the

future Pyrenees) resulted from left-lateral motion

of Africa-Iberia with respect to Eurasia (Savostin

et al., 1986; Srivastava and Tapscott, 1986; Klit-

gord and Schouten, 1986; Malod, 1989). The North

Pyrenean Fault can be approximated by a small

circle around the same rotation pole as that for

the movement direction of the collisional tectonics

in the Internal Zone, indicating a coupling be-

tween both motions.

Pyrenean collision and tectonic quiescence in the

Internal Zone

A new position of the combined African-

Iberian rotation pole after 80 Ma near the Straits

of Gibraltar (Savostin et al., 1986) (Fig. 7a) re-

sulted in a dramatic change in the tectonic regime

in the Pyrenees and in the Betics. The formerly

thinned, heated and consequently weakened

Pyrenean 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 et

al., 1990). Oblique convergence started in the

Campanian (Puigdefabregas and Souquet, 1986)

mimicking the rotation pole evolution very accu-

rately. Deformation culminated in the Eocene

290 K. DE JONG

Fig. 7. Positions of the African ( Af )-Iberian (16) combined rotation pole after Savostin et al. (1986). (a) Quadrangles indicate the 80-65 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 and

Souquet, 1986), during continuing convergence be-

tween Africa and Eurasia around the Gibraltar

pole. Savostin et al. (1986) indicate that motion

around this pole resulted in about 100 km of

shortening in the Pyrenees; this accords with the

analysis of the ECORS profile (Rome et al., 1989).

Thermal disturbances at 60-55 Ma in the Pyrenean

shear zones {Costa and Malt&i, 1988) and at the

northwestern termination of the Iberic Cordillera

(Golberg et al., 1988) coincide with initiation of

spreading in the Norwegian-Greenland Sea

(Srivastava and Tapscott, 1986; Klitgord and

Schouten, 1986; Savostin et al., 1986). This in-

duced an additional compressional component in

western Eurasia.

Because Iberia was attached to Africa it acted

as an African promontory during collison with

Eurasia in the Pyrenees. It suffered relatively in-

tense deformation by reactivation of Variscan and

Mesozoic faults (Guimera, 1984; Viallard, 1985) in

a weakened lithosphere by Cretaceous extension

(Zoetemeijer et al., 1990). Convergence resulted in

limited riot-later~ motion of main Iberia with

respect 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 direction

u1 in the Catalan Coastal Range (Guimera, 1984)

coincides with the motion direction around the

Gibraltar poie.

During the Pyrenean collision the Internal Zone

was tectonically relatively quiet, as is evident from

the absence of penetrative deformation for about

50 Ma, between 83 Ma and 33 Ma. This is in close

agreement with the geochronological evolution of

the Kabylian Massifs in northern Algeria. The

location of the African-Iberian rotation pole near

the Cretaceous Betic collision zone (Fig. 7a) in-

bibited large-scale motions in this zone after 80

Ma. Minor tectonics is documented in the Exter-

nal Zone of the Betics in the form of relative uplift

between 68 Ma and 60 Ma; this probably resulted

from compression caused by African-Eurasian

convergence (Kenter et al., 1990).

After 54 Ma the African rotation pole was

separated from the Iberian pole and shifted north-

ward (Savostin et al., 1986) (Fig. 7b). Shifting

might be partly explained by the Pyrenean colli-

sion, forcing the African plate to pivot around a

different pole. The effect of this new rotation pole

position is the starting of limited differential mo-

tion between Africa and Iberia, resulting in the

tectonism which has been well documented in the

Betic Cordilleras. The External Zone records a

second period of uplift at about 50 Ma (Kenter et

al., 1990). The Malaguide Complex shows an ero-

sional contact between the Maastrichtian and the

Early Eocene (Roep, 1980). Because the metamor-

phic zones of southern Spain and northern Africa

do not record penetrative deformation and meta-

morphism, this tectonic activity is probably of


minor importance. Northerly to north-northeast-

erly compression is detected in the late Eocene in

the Catalan Coastal Range (Guimera, 1984) and

the External Betics (De Ruig, this issue) and in

other areas in Spain (Letouzey, 1986; Bergerat,

1987). The main compression direction is coinci-

dent with the motion vector around the 54-35 Ma

African rotation pole of Savostin et al. (1986)

(Fig. 7b). The mountain ranges of northern Africa

also experienced compression during this episode

(Letouzey, 1986; Dercourt et al., 1986) presuma-

bly connected with the Pyrenean collision. The

Pyrenean collision ended at about 35 Ma, after

which no independent lberian rotation pole may

be detected (Savostin et al., 1986). Klitgord and

Schouten (1986) concluded that the final

amalgamation of Iberia with Eurasia occurred at

about 30 Ma.

Extensional tectonics

Savostin et al. (1986) demonstrate a dramatic

shift of the African rotation pole to a new position

in the South Atlantic at 35 Ma as a consequence

of 40% decrease in spreading rate between the

African and North American plates; no change

occurred in Eurasia-North America spreading.

This kinematic pattern is, however, not supported

by the central Atlantic magnetic anomaly pattern

(Klitgord and Schouten, 1986) or by North

Atlantic and Arctic data (Srivastava and Tapscott,

1986). Klitgord and Schouten (1986) indicate a

shift of the African-Eurasian plate boundary from

the northern side of Iberia to the southern side of

it, separating Iberia from Africa by the Azores-

Gibraltar fracture zone after 30 Ma. The establish-

ment of a new plate boundary between Iberia and

Africa accords with timing of extensional defor-

mation and heating in the Internal Zone and

northern Africa. Regionally consistent southeast-

erly slip of the hangingwall towards ultramafic

rocks in the western Betics (Tubia and Cuevas,

1986) and the Rif (Saddiqi et al., 1988) suggests a

SE-NW extensional component on the plate

boundary. Extension is further attested by an Early

Miocene basaltic dike-swarm (Torres-Roldan et

al., 1986). The timing of extension and magmatism

accords well with the tectonic evolution of the

Gulf of Lion and the Balearic Basin, which opened

in Oligocene to Aquitanian times (Alvarez et al.,

1974; Rehault et al., 1984). Deformation in this

area spread southward with time (Mauffret and

Gennesseaux, 1989). Extension resulted in pro-

gressive crustal thinning in northeastern Spain to-

wards the Balearic Basin (Banda, 1987). Strong

crustal thinning in the eastern part of the Internal

Zone (Banda and Ansorge, 1980) is also attributed

to this Oligo-Aquitanian crustal thinning and ex-

tension. Extensional deformation to the east of

Iberia links the African-Eurasian plate boundary

with the western European rift system (Bresse-

Rhine graben) which experienced extension up to

the latest Oligocene (Bergerat, 1987). The shift of

the African rotation pole at 35 Ma accords well

with renewal of deformation in the Betics and

northern Africa. However, the motion around this

pole (Savostin et al., 1986) generates northwesterly

compression between Africa and Eurasia and

hence cannot explain the observed extension.

The 20 Ma position of the African rotation

pole to the northwest of Portugal (Savostin et al.,

1986) resulted in cessation of extensional deforma-

tion and induced a northeasterly relative motion

of Africa with respect to Eurasia, resulting in

NE-SW directed compression (Letouzey, 1986;

Bergerat, 1987). Compression caused thickening of

the 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 reactivation

of the original contact between the Alpujarride

and Mulhacen Complexes. The Alpujarride Com-

plex has moved northwards on a shear zone with

respect to the underlying Mulhacen Complex (Platt

et al., 1983) which already contains the imprints

of the 22 Ma Dx+3 thermal peak. Behrmann (1984)

concluded on the basis of palaeostress and strain

rate estimates that ductile deformation in this

shear zone lasted 4 Ma. Consequently, ductile

thrusting stopped during the Burdigalian. This

accords with the superposition of the Malaguide

Complex onto the already quite well developed

Alpujarride Complex around the Aquitanian-

Burdigalian boundary (Makel, 1981) and with

sealing of thrusts in the Malaguide Complex by

Burdigalian sediments (MacGillavry et al., 1963;

292 K DE JONG

Torres-Roldan et al., 1986). Burdigalian sedimen-

tation in large parts of the Internal Zone occurred

in a tectonically quiet environment (Volk, 1967).

The compression direction after 20 Ma enables

early left-lateral motion on ENE-WSW to NE-

SW trending strike-slip faults crossing the Betic

Zone. Middle Miocene strike-slip motion has been

described by Sanz de Galdeano et al. (1985) and

by Bon et al. (1989). The 10 Ma to present rota-

tion pole position of Africa, to the west of

Gibraltar (Savostin et al., 1986) enables the

northwesterly to northerly compression detected

by Montenat et al. (1987) in the Betics. The posi-

tion of the main compression axes in the obtuse

angle of a set of NE-SW and NNE-SSW trend-

ing strike-slip faults (Montenat et al., 1987) points

to a reactivation of an earlier fault system. The

position of the rotation pole relatively close to the

strike-slip fault system only allows limited dis-

placements. About 20 km of left-lateral slip has

been 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 late

Oligocene to Aquitanian extension. Present-day

compression and intraplate seismicity (Udias et

al., 1976) is localized in the thinned Alboran Sea

and southern Spain.

Conclusions

The tectonic evolution of the Alpine collision

belts bordering Iberia is in accordance with

kinematics of the Iberian and African plates and

their rotation pole evolution.

Subduction and HP/LT metamorphism in the

Betics at 116 + 10 Ma coincide with initiation of

seafloor spreading at the northwestern margin of

Iberia and the Bay of Biscay. This phase also

coincides with heavy terrigeneous sedimentation

in the Mauritanian flysch in northern Africa and

the tectosedimentary evolution of the Malaguide

Complex.

A period of intracontinental thrusting between

99 Ma and 83 Ma in the Internal Zone of the

Betics occurred during continuing spreading to the

west of Iberia and progressive opening of the Bay

of Biscay. Regionally consistent ESE-WNW

trending stretching lineations in the Internal Zone,

which were formed during thrusting at lower

crustal levels, are coincident with the motion di-

rection of Africa-Iberia with respect to Eurasia.

At least 200 km of thrusting in the Internal Zone

was driven by the convergence of Africa-Iberia

with respect to Eurasia. Strike-slip motion of

Africa-Iberia with respect to Eurasia along the

North Pyrenean Fault took place around the same

rotation pole as thrusting in the Betics.

At 80 Ma, shifting of the African-Iberian rota-

tion pole to a position near the Betic collision

zone 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 African

promontory by motion around the rotation pole

near Gibraltar. Pyrenean collision culminated at

60-55 Ma by an additional compression generated

by seafloor spreading in the Norwegian-Green-

land Sea.

After completion of the Pyrenean collision at

35-30 Ma, penetrative deformation and heating

again took place in the Internal Zone and north-

ern Africa by establishment of the Azores-

Gibraltar fracture zone, which interacted with the

Cretaceous collision system. The new plate

boundary was connected via the Balearic rift sys-

tem with the western European rift. Deformation

was initiated by late Oligocene crustal and litho-

spheric mantle extension, resulting in Early

Miocene emplacement of ultramafic rocks with

associated 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 previously

extended and heated and consequently weakened

Internal Zone. This process resulted in overthrust-

ing in the Internal Zone, which was completed

before the end of the Burdigalian. Middle and

Late Miocene strike-slip motion was initiated dur-

ing continuing convergence and juxtaposed crustal

segments of differing Moho depths, the latter hav-

ing been inherited from late Oligocene to Early

Miocene extension.

Acknowledgements

I thank Sierd Cloetingh for constructive criti-

cism on the manuscript and Henk Helmers for

ALPINE TECTONK’S AND ROTATION POLE EVOLUTION OF IBERIA 293

placing the strain data on plagioclase porphyro-

clasts at my disposal. Jacques Malod, Paul

Andriessen and Reini Zoetemeier are thanked for

providing preprints. The technical assistance of

Fred Kievits during the preparation of the figures

is aiso acknowledged. Part of the field work was

financed by a grant from the “Stichting

Molengraaff-fonds”.

References

Albarede, F. and Michard-Vitrac, A., 1978. Age and signifi-

cance of the north Pyrenean metamorphism. Earth Planet.

Sci. Lett., 40: 327-332.

Alvarez, W., Cocozza, T. and Wezel, F.C., 1974. Fragmenta-

tion of the Alpine erogenic belt by microplate dispersal.

Nature, 248: 309-314.

Andriessen, P.A.M., Hebeda, E.H. and Verschure, R.H., 1989.

Tourmaline dating in Alpine anatectic leucosomes and

polymetamorpbic rocks: a feasibility study in the Cyclades

and southern Spain. Chem. Geol., in press.

Baird, A.W. and Dewey, J.F., 1986. Structural evolution in

thrust belts and relative plate motion: the Upper Pennine

Piemont Zone of the Internal Alps, southwestern Switzer-

land and northwest Italy. Tectonics, 5: 375-387.

Bakker, H.E., De Jong, K., Helmers, H. and Biermann, C.,

1989. The geodynamic evolution of the Internal Zone of the

Betic Cordilleras (SE Spain): a model based on structural

analysis and geothermobarometry. J. Metamorph. Geol., 7:

359-381.

Banda, E., 1987. Crustal parameters in the Iberian Peninsula.

Phys. Earth Planet. Inter., 51: 222-225.

Banda. E. and Ansorge, J., 1980. Crustal structure under the

central and eastern part of the Betic Cordillera. Geophys.

J.R. Astron. Sot., 63: 515-532.

Bebrmann, J.H., 1984. Patterns of palaeostress patterns be-

neath the Aquilon nappe. Z. Dtsch. Geol. Ges., 135: 293-

305.

Ben Othman, D., Polve, M. and Allegre, C.J., 1984. Nd-Sm

isotope composition of granulites and constraints on the

evolution of the lower continental crust. Nature, 307: 510-

515.

Bergerat, F.. 1987. Stress fields in the European platform at the

time of Africa-Eurasia collision. Tectonics, 5: 99-132.

Besems, R.E. and Simon, O.J., 1982. Aspects of Middle and

Late Triassic palynology. 5. On the Triassic of the Subbetic

Zone in the province of Murcia. Verh. K. Ned. Akad. Wet.,

Ser. B, 85: 29-51.

Bon, A., Biermann, C., Koenen, D.B. and Simon, O.J., 1989.

Middle Miocene strike-slip tectonics in the Aguilas-Mazar-

ron region, SE Spain. hoc. K. Ned. Akad. Wet., Ser. B, 92

(2): 143-157.

Boillot, G. and Malod, J., 1988. The North and Northwestern

Spanish Continental Margin: A review. Rev. Sot. Geol.

Esp., 1: 295-316.

Boillot, G., Girardeau, J. and Komprobst, J.. 1989. Rifting of

the West Galicia Continental margin: a review. Bull. Sot.

Geol. Fr., 8: 393-400.

Borradaile, G.J., 1976. A strain study of a granite-granite

gneiss transition and accompanying schistosity formation

in the Betic erogenic zone, SE Spain. J. Geol. Sot. London,

132: 417-428.

Bouillin, J., Komprobst, J. and Raoult, J., 1977. Don&es

preliminaires sur le complexe volcano-stdimentaire de Re-

kkada Metletine (ex-Texenna), en Petite Kabylie (Algtrie).

Bull. Sot. Geol. Fr., 4: 805-813.

Bouillin, J., Durand Delga, M. and Olivier, Ph., 1986. Betic-Rif

and Tyrrhenian: distinctive features. genesis and develop-

ment stages. In: F.C. Wezel (Editor), The Origin of Arcs.

Elsevier, Amsterdam, pp. 281-304

Choukroune, P. and Mattauer, M., 1978.Tectoniques de plaques

et Pyrenees: sur le fonctionnement de la faille transfor-

mante nord-pyreneenne; comparaison avec les modeles ac-

tuels. Bull. Sot. GCol. Fr., 5: 689-700.

Choukroune, P., Ballevre, M., Cobbold, P., Gautier, Y., Merle,

0. and Vuichard, J-P., 1986. Deformation and motion in

the western Alpine arc. Tectonics, 5: 215-226.

Costa, S. and Maluski, H., 1988. Use of the 40Ar-‘9Ar stepwise

heating method for dating of mylonite zones: An example

from the St. Barth&my Massif (northern Pyrenees, France).

Chem. Geol., 72: 127-144.

Crough, S. and Thompson, G.A., 1976. Numerical and ap-

proximate solutions for lithosphetic thickening and thin-

ning. Earth Planet. Sci. Lett., 31: 397-402.

Davy, Ph. and Gillet. Ph., 1986. The stacking of thrust slices in

collision zones and its thermal consequences. Tectonics. 5:

913-929.

Dercourt, J., Zonenshain, L.P.. Ricou, L.-E., Kazim, V.G., Le

Pichon, X., Knipper, A.L.. Grandjaquet, C., Sbortshikow,

I.M., Geyssant, J., Lepvrier, C., Pechersky. D.H., Boullin,

3.. Sibuet, J.-C., Savostin, L.A., Sorokhtin. O., Westphal,

M., Bazbenov. M.L.. Lauer, J.P. and Biju-Duval, B., 1986.

Geological evolution of the Tethys belt from the Atlantic to

the Pamirs since the Lias. Tectonophysics. 123: 241-315.

De Ruig, M.J.. 1990. Fold trends and stress deviation in the

Alicante fold belt, southeastern Spain. Tectonophysics. 184

(this issue): 393-403.

De Ruig, M.J., Mier, R.M. and Stel, H., 1987. Interference of

compressional and wrenching tectonics in the Alicante re-

gion, SE Spain. Geol. Mijnbouw, 66: 201-212.

De Smet, M.E.M., 1984. Wrenching in the External Zone of

the Betic Cordilleras, Southern Spain. Tectonophysics. 107:

57-79.

Doblas, M. and Oyarzun, R., 1989. Neogene extensional col-

lapse in the western Mediterranean (Betic-Rif Alpine oro-

genie belt): Implications for the genesis of the Gibraltar

Arc and magmatic activity. Geology, 17: 430-433.

Egeler, C.G. and Simon, O.J., 1969. Sur la tectonique de la

Zone betique (Cordilleres bttiques, Espagne). Verh. K.

Ned. Akad. Wet., 25: 90 pp.

England, P.C. and Thompson, A.B., 1984. Pressure-tempera-

ture-time paths of regional metamo~hism I. Heat transfer

294 K DE JONG

during the evolution of regions of thickened continental

crust. J. Petrol.. 25: 894-928.

Esscher, A. and Watterson, J.S., 1974. Stretching fabrics. folds

and crustal shortening. Tectonophysics, 22: 223-231.

Feraud, G., Girardeau, J., Beslier, M.-O. and Boillot, G., 1988.

Datation ‘“Ar-“OAr de la mise en place des ptridotites

bordant la marge de la Galice (Espagne). C.R. Acad. Sci.

Paris. 307: 49-55.

Geel. T.. 1979. Malaguide and Subbetic Jurassic sediments

between Velez Rubio and Caravaca (Eastern Betic

Cordilleras. SE Spain). La sedimentation Jurassique W.

Europeen. Assoc. Stdimentol. Fr. Publ. Spec., 1: 335-344.

GoffCe, B. Michard, A., Garcia-Dueiias, V., Gonzalez-Lodeiro,

F., MoniC, P.. Campos, J., Galindo-Zaldivar, J.. Jabaloy.

A.. Martinez-Martinez. J.M. and Simancas, J.F., 1989. First

evidence of high-pressure. low-temperature metamorphism

in the Alpujarride nappes. Betic Cordilleras (SE. Spain).

Eur. J. Mineral., 1: 1399142.

Golberg, J.M.. Maluski, H. and Leyreloup, A.F., 1986. Petro-

logical and age relationships between emplacement of mag-

matic breccia, alkaline magmatism, and static metamor-

phism in the North Pyrenean Zone. Tectonophysics, 129:

275-290.

Golberg. J.M., Guiraud, M., Maluski, H. and Seguret, M.,

1988. Caracttres petrologiques et Lge du mttamorphisme

en contexte distensif du bassin sur dtcrochement de Soria

(C&ace inferieur, Nord Espagne). C.R. Acad. Sci. Paris,

307: 521-527.

Gomez-Pugnaire, M.R. and Femandez-Soler, J.F., 1987. High

pressure metamorphism in metabasites from the Betic

Cordilleras (S.E. Spain) and its evolution during the Alpine

orogeny. Contrib. Mineral. Petrol., 95: 231-244.

Gomez-Pugnaire, M.T. and Franz, G., 1988. Metamorphic

evolution of the Palaeozoic series of the Betic Cordilleras

(Nevado-Filabride complex, SE Spain) and its relationship

with the Alpine orogeny. Geol. Rundsch., 77: 619-640.

Guimera. J., 1984. Paleogene evolution of the deformation in

the northeastern Iberian Peninsula. Geol. Msg., 121: 413-

420.

Hebeda. E.H.. Boelrijk, N.A.I.M.. Priem, H.N.A., Verdurmen.

E.A.Th., Verschure, R.H. and Simon, O.J., 1980. Excess

radiogenic Ar and undisturbed Rb-Sr systems in basic

intrusives subjected to Alpine metamorphism in southeast-

em Spain. Earth Planet. Sci. Lett., 47: 81-90.

Hermes, J.J.. 1978. The stratigraphy of the Subbetic and south-

ern Prebetic of the Velez Rubio-Caravaca area and its

bearing on the transcurrent faulting in the Betic Cordilleras

in southern Spain. Verh. K. Ned. Akad. Wet., Ser. B, 87:

319-333.

Holdaway, M.J., 1971. Stability of andalusite and the

aluminium-silicate diagram. Am. J. Sci., 271: 97-131.

Hoschek, G., 1969. The stability of staurolite and chloritoid

and their significance in metamorphism of pelitic rocks.

Contrib. Mineral. Petrol., 22: 208-232.

Kampschuur. W. and Rondeel, H.E., 1975. The origin of the

Betic Orogen, southern Spain. Tectonophysics, 27: 39-56.

Kenter, J.A.M., Reymer, J.J.G.. Van der Straaten, H.C. and

Peper, T., 1990. Facies patterns and subsidence history of

the Jumilla-Cieza region (south-eastern spain). Sediment.

Geol., 67: 263-280.

Klitgord, K.D. and Schouten. H.. 1986. Plate kinematics of the

Central Atlantic. In: B.E. Tucholke and P.R. Vogt (Editors),

The Geology of North America: the Western Atlantic

Region. (Decade of North American Geology, Ser. 1.)

Geol. Sot. Am., Boulder. Colo.. pp. 351-378.

Leblanc. D. and Ohvier, P.. 1984. The role of strike-slip faults

in the Betic-Rifian orogeny. Tectonophysics, 101: 3455355.

Le Pichon, X., Bonnin. J.. Francheteau, J. and Sibuet, J.C.,

1971. Une hypothese d’evolution tectonique du Golfe de

Gascogne. In: J. Debyser et al. (Editors). Histoire Structur-

ale du Golfe du Gascogne. Technip., Paris. Vol. VI.1 1. pp.

l-44.

Letouzey, J.. 1986. Cenozoic paleo-stress pattern in the Alpine

Foreland and structural interpretation in a platform basin.

Tectonophysics. 132: 215-231.

MacGillavry, H.J., Geel, T., Roep, Th. B. and Soediono. H..

1963. Further notes on the geology of the Betic of Malaga,

the Subbetic and the zone between these two units in the

region of Velez Rubio (southern Spain). Geol. Rundsch.,

53: 623-625.

Majoor, F.J.M., 1988. A geochronological study of the Axial

Zone of the Central Pyrenees, with emphasis on Variscan

events and Alpine resetting. Verh. Zuiver Wet. Onderz.

Lab. Isot. Geol., Amsterdam, 6: 117 pp.

Mlkel, G.H., 1981. Differences in tectonic evolution of super-

posed Malaguide and Alpujarride tectonic units in the

Espuiia area (Betic Cordilleras, Spain). Geol. Mijnbouw,

60: 203-208.

Malod. J.. 1982. Comparaison de I’evolution des marges con-

tinentales au N at au S de la peninsula iberique. Thesis,

Univ. Paris, 235 pp.

Malod, J.A., 1989. Ibetides et plaque iberique. Bull. Sot. GCol.

Fr.. 8: 927-934.

Maresch, W.V., 1977. Experimental studies on glaucophane:

An analysis of present knowledge. Tectonophysics. 43:

1099125.

Mattauer, M. and Henry, J., 1974. Pyrenees. In: A.M. Spencel

(Editor), Mesozoic-Cenozoic Belts: Data for Orogenic

Studies. Geol. Sot. London Spec. Publ., 4: 3-11.

Mauffret, A. and Gennesseaux, M., 1989. Compression,

decrochements et distension sur le pourtour mediterranten

nord-occidental, CR. Acad. Sci. Paris, 308: 961-967.

McCaig, A.M. and Miller, J.A., 1986. ‘9Ar-40Ar age of

mylonites along the Merens fault, central Pyrenees.

Tectonophysics, 129: 149-172.

MoniC, P., Caby, R. and Maluski, H., 1984a. ‘9Ar-40Ar investi-

gations within the Kabylie Massif (northern Algeria): evi-

dence for its Alpine structure. Eclogae Geol. Helv.. 77:

115-141.

Monie, P., Frizon de Lamotte, D. and Leikine, M., 1984b.

Etude gechronologique preliminaire par la methode

39Ar/40Ar du metamorphism alpin dans le rif externe

ALPINE TECTONI<‘S AND ROTATION POLE EVOLUTION OF IBERIA 295

(Maroc). Precisions sur le calendrier tectonique tertiare.

Rev. GCogr. Phys. Gtol. Dyn., 25: 307-317.

MoniC, P., Malt&i, H.. Saadaliah, A. and Caby, R.. 1988. New

“‘A’A~-~“A~ ages of Hercynian and Alpine thermotectonic

events in Grande Kabylie (Algeria). Tectonophysics, 152:

53-69.

Montenat. Chr.. Ott d’Esstevou, Ph. and Masse. P., 1987.

T~tonic-sedi~l~ntary characters of the Betic Neogene

basins evolving in a crustal transcurrent shear zone (SE

Spain). Bull. Cent. Rech. Explor.-Prod. Elf-Aquitaine, 11:

Montigny. R.. Azambre, B., Rossy, M. and Thuizat. R., 1986.

K-Ar study of Cretaceous magmatism and metamorphism

in the Pyrenees: age and length of rotation of the Iberian

peninsula. Tectonophysics. 129: 257-273.

Moretti, J. and Froideveaux. C.. 1986. The~omechanical

models of active rifting. Tectonics. 5: 501-511.

Nagtegaal. P.J.C. and De Weerd, J.T., 1985. Provenance of

Cambro-Ordovician to Oligocene sandstones in the south-

ern Pyrenees. Geol. Mijnbouw, 64: 25-40.

Newton, M.S. and Kennedy. G.C.. 1968. Jadeite, analcite,

nepheline and albite at high temperatures and pressures.

Am, J. Sci., 266: 728-735.

Newton, R.C. and Kennedy, G.D., 1963. Some equilib~um

reactions in the Join CaAl,Si,O,-H,O. J. Geophys. Res.,

68: 2967-2983.

Platt. J.P.. 1987. The uplift of high-pressure low-temperature

metamorphic rocks. Philos. Trans. R. Sot. London. Ser. A.

321: 87-103.

Platt. J.P., Van den Eeckhout. B., Janzen, E. Konert, K.,

Simon. O.J., and Weijermars, R., 1983. The structure and

tectonic evolution of the Aguilon Fold-nappe, Sierra Al-

hamiila. Betic Cordiileras. SE Spain, J. Struct. Geol., 5:

519-538.

Priem, H.N.A.. Boelrijk, N.A.I.M., Hebeda, E.H. and

Verschure. R.N.. 1966. Isotopic age determinations on

tourmaline granite-gneisses and a metagranite in the east-

ern Betic Cordilleras (southern Sierra de 10s Filabres), SE

Spain. Geol. Mijnbouw, 45: 184-187.

Priem. H.N.A., Btrelrijk, N.A.I.M. Hebeda, E.H., Oen, IS.,

Verdurmen, E.A.Th and Verschure. R.H., 1979. Isotopic

dating of the emplacement of the ultramafic masses in the

Serrania de Ronda, southern Spain. Contrib. Mineral. Pet-

rol.. 70: 103-109.

Puga, E. and Diaz de Federico, A.. 1978. Metamorfismo

polifasico y deformaciones alpinas en el Complejo de Sierra

Nevada (Cordillera Betica). Implicaciones geodinamicas.

Proc. Reun. Geodinamica de la Cordillera BCtica y Mar de

Albodn (Univ. Granada), pp. 79-111.

Puigdefabregas, C. and Souquet, P.. 1986. Tecto-sedimentary

cycles and depositional sequences of the Mesozoic and

Tertiary from the Pyrenees. Tectonophysics, 129: 173-203.

Rehault, J-F.. Bouillot, G. and Mauffret, A., 1984. The western

Mediterranean basin, geological evolution. Mar. Geol., 55:

447-477.

Richardson. S.W. and England, P.C.. 1979. Metamorphic con-

sequences of crustal eclogite production in overthrust oro-

genie zones. Earth Planet. Sci. Lett., 42: 183-190.

Ricou, L.E.. Dercourt, J., Geysant. J.. Grandjaquet, C., Lepv-

tier, C. and Biju-Duval. B., 1986. Geological constraints on

the Alpine evolution of the Mediterranean Tethys.


Roep, Th.B., 1980. Condensed Cretaceous limestones in a

section near Xiquena, Betic of Malaga, SE Spain. I. Gen-

eral description and comparison with other areas in SE

Spain. Verh. K. Ned. Akad. Wet., Ser. B, 75: 183-200.

Rondeel. H.E. and Simon. O.J., 1974. Betic Cordilleras. In:

A.M. Spencer (Editor), Mesozoic-Cenozoic Belts: Data for

Orogenic Studies. Geol. Sot. London Spec. Publ.. 4: 23-35.

Rosenfeld. J.L., 1970. Rotated garnets in metamorphic rocks.

Geol. Sot. Am. Spec. Pap., 129: 102 pp.

Roure, F.. Choukroune. P. and the ECORS Team, 1989.

ECORS deep seismic data and balanced cross section:

geometric constraints on the evolution of the Pyrenees.

Tectonics, 8: 41-50.

Saddiqi, O., Reuber, 1. and Michard, A., 1988. Sur la tectonique

de denudation du manteau infracontinental dans les Beni

Boursera, Rif septentrional, Maroc. C.R. Acad. Sci. Paris,

307: 657-662.

Sam de Galdeano, C.. Rodriguez-Fernandez, J. and Lbpez-

Garrido, A.C., 1985. A strike-slip fault corridor within the

Alpujarra Mountains (Betic Cordilleras. Spain). Geol.

Rundsch., 74: 641-655.

Savostin. L.A., Sibuet, J., Zonenshain, P., Le Pichon, X. and

Roulet, M.. 1986. Kinematic evolution of the Tethys belt

from the Atlantic Ocean to the Pamirs since the Triassic.

Tectonophysics, 123: I-35.

Shackleton, R.M. and Ries, A.C., 1984. The relation between

regionally consistent stretching lineations and plate mo-

tions, J. Struct. Geol., 6: 111-117.

Simon, O.J., 1987. On the Triassic of Betic Cordilleras (South-

ern Spain). Cuad. Geol. lber., 11: 385-402.

Soula, J.C., Lamouroux, C.. Viallard, P.. Bessiere, G.. Debat. P.

and Ferret, B.. 1986. The mylonite zones in the Pyrenees

and their place in the Alpine tectonic evolution.


Souquet, P., Debroas, E.J., Boiric. J.M., Pons. P., Fixari, G.,

Roux. J.C.. Dol. J., Thieuloy, J.P.. Bonnemaison, M.,

Manivit, H. and Peybernes, B., 1985. The black ffysch

(Albian-Early Cenom~i~) from the Pyrenees. Bull. Cent.

Rech. Explor.-Prod. Elf-Aquitaine, 9: 183-252.

Srivastava. S.P. and Tapscott, C.R., 1986. Plate kinematics of

the North Atlantic. In: B.E. Tucholke and P.R. Vogt (Edi-

tors), The Geology of North America: the Western Atlantic

Region. (Decade of North American Geology. Ser. 1.)

Geol. Sot. Am., Boulder, Colo.. pp. 3799404.

Thompson. A.B., 1981. The pressure--temperature plane of

regional metamo~hism. Terra Cog&a, I : 10-20.

Torres-Roidan, R.L.. Poli, G. and Peccerillo. A., 1986. An

Early Miocene arc-tholeiitic magmatic dike event from the

296 K. DE JONEi

Alboran Sea-evidence for pre-collisional subduction and

back-arc crustal extension in the westernmost Mediter-

ranean. Geol. Rundsch., 75: 219-234.

Tubia, J.M. and Cuevas, J., 1986. High temperature emplaee-

ment of the Los &ales peridotite nappe (Betic Cordilleras,

Spain). J. Struct. Geol., 8: 473-482.

Udias, A., Lopez Arroyo. AL. and Mezcua, J., 1976. Seismo-

tectonics of the Azores-Alboran region. Tectonophysics,

31: 259-289.

Van de Fliert, J.R., Hebeda, E.H., Priem, H.N.A., De Smet,

M.E.M. and Vetdurmen, E.A.Th., 1979. Pillow lavas and

basic intrusives of Early Cretaceous age in the Cantar area

(prov. Murcia), Spain. Estud. Geol., 35: 605-608.

Veeken, P.C,H.. 1983. Stratigraphy of the Neogene-Quaternary

Pulpi Basin provinces Murcia and Almeria (SE Spain).

Geol. Mijnbouw, 62: 255-265.

Velilla. N. and Fen011 Ha&Ah, P., 1986. The composition and

zoning of metabasite garnets in the Sierra Nevada (Betic

Cordilleras, SE Spain). Neues Jab&. Mineral. Petrol_ A&.,

154: 41-56.

Viallard, P., 1985. lb&ides et IbCrie: un exemple de relations

entre tectognese intracontinentate et tectonique des plaques.

CR. Acad. Sci. Paris, 300: 217-222.

Vielzeuf. D. and Kornprobst, J., 1984. Crustal splitting and the

emplacement of Pyrenean lherzolites and granubtes. Earth

Planet. Sci. Lett., 67: 87-96.

Vissers, R.L.M., 1989. Asymmetric quartz c-axis fabrics and

flow vorticity: a study using rotated garnets. 3. Struct.

Geol., 11: 231-244.

Viilk, HR., 1967. Zur Geologic und Strati~aphi~ des

Neogenbeckens von Vera. Sudost-Spanien. Thesis. Univ.

Amsterdam, Amsterdam, 164 pp.

Westerhof. A.B., 1977. On the contact relations of high temper-

ature peridotites in the Serrania de Ronda, southern Spain.


Williams, G.D., 1985. Thrust tectonics in the south central

Pyrenees. J. Struct. Geol., 7: 11-17.

Wood. D.S., 1974. Current views on the development of slaty

cleavage. Annu. Rev. Earth Planet. Sci.. 2: 369-401.

Zindler, A., Staudigel, H., Hart, S.R. Enders, R. and Goldstein.

S., 1983. Nd and Sr isotope study of a mafic complex.

Nature, 304: 226-230.

Zoetemeijer, R., Desegaulx, I?., Cloetingh, S., Route, F. and

Moretti, l., 1990. Lithospheric dynamics and tectonic-stra-

tigraphic evolution of the Ebro Basin. J. Geophys. Res., 95:

2701-2711.

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Alpine tectonics and rotation pole evolution of Iberia