1. metamorphic Geol., 1989, 7, 359-381

The geodynamic evolution of the Internal Zone of the Betic Cordilleras (south-east Spain): a model based on structural analysis and geothermobarometry H. E . BAKKER' Ceologisch Instituut, Universiteit van Amsterdam, Nieowe Prinsengracht 130, 1018 VZ Amsterdam, The Netherlands

K. DE JONC,' H. H E L M E R S AND C. B I E R M A N N lnstituut voor Aardwetenschappen, Vrije Universiteit, Postbus 7161, 1007 MC Amsterdam, The Netherlands

ABSTRACT The Internal Zone of the Betic Cordilleras consists of several superimposed major thrust sheets with different P-T-t evolutions. On the basis of an integrated field, miCroscopic and laboratory study, the tectono-metamorphic history of the Mulhacen Complex and Almanzora Unit has been reconstructed in detail. The Mulhacen Complex has been af€ected by at least five phases of penetrative deformation, which have been labelled Dx-,, D., DX+,, D,,, and DX++ D,-, and D, are related to continent-continent collision, which is indicated by high pressure-low temperature (WILT) and subsequent intermediate PIT metamorphic conditions. D,,, is related to crustal thinning and heterogeneous extension. During this event the ALmanzora Unit was juxtaposed against the Mulhacen Complex. This phase was succeeded by the establishment of low pressure-high temperature (LPIHT) conditions and at least two phases of folding and overthrusting. The Almanzora Unit shows a comparable tectono-metamorphic evolution p t D,,,. However, the P IT conditions prior to D.,, indicate a higher crustal position with respect to the Mulhacen Complex during the collisional event.

Key w& Alpine orogeny; Betic Cordilleras, Spain; continent-continent collision; crustal thickening; crustal thinning; deformation structures; geothermobarometry; P-T-t path.

.

INTRODUCTION

The Betic Cordilleras of southern Spain represent the westemmost part of the pen-mediterranean alpine orogenic belt of southern Europe. It has a northern External Zone of essentially non-metamorphic Triassic to Middle Miocene sediments, which were deposited onto the southern continental margin of the Iberian plate (Hemes, 1978; Garcia-Hernandez, Lopez-Garrido, Rivas, Sanz de Galdeano & Vera, 1980) and a southern Internal Zone, comprising mainly intensely deformed and metamorphosed sedimentary rocks of dominantly Triassic and Palaeomic age. The depositional realm of these metasediments is not well known (Bourrouilh & Gorsline, 1979; M2ke1, 1985).

The regional structure of the Internal Zone originates from large-scale overthrusting. The individual thrust units have been organized in four thrust sheet complexes on the basis of differences in tectono-metamorphic evolution.

The lowermost thrust complex is the Veleta Complex,

Present addrtss:%atec Geo-Exploration Consultants, Noor- dcrstraat 76"'. 1017 TW Amsterdam, The Netherlands.

* Reprint requests to K. de Jong.

characterized by low pressure/low temperature (LPILT) metamorphism (Puga & Diaz de Federico, 1978; Diaz de Federico, Gomez-Pugnaire, Puga & Sassi, 1979). It is overlain by the Mulhacen Complex, a series of thrust sheets with early high pressure/low temperature (HPILT) metamorphism, subsequently overprinted by medium- grade metamorphism of intermediate and low pressure type (Nijhuis, 1964; Puga & Diaz de Federico, 1978; Diaz de Federico et uf., 1979; Gomez-Pugnaire & Fernandez- Soler, 1987). Both complexes are commonly included into the ill-defined Nevada-Fiabride Complex (Egeler, 1964).

In the Internal Zone the thrust sheets of the Veleta and Mulhacen Complex have been overthrusted by a large number of thrust slices of mainly the Alpujarride Complex and the overlying Malaguide Complex. The thrust units of the Alpujarride Complex show a variable metamorphic grade indicating intermediate to low PIT ratios and locally W / H T conditions (Westerhof, 1975; Torres-Roldan, 1979; Akkerman, Maier & Simon, 1980). The Malaguide units consist of w n - or very low-grade metamorphic rocks (Roep & MacGillavry, 1962; Egeler & Simon, 1969).

In the northeastern part of the Internal Zone Alpujarride thrust units are overlying rocks of the Almagride Complex,

353

360 H. E. BAKKER ET AL.

QZ-

i ' ! i ' i l

I I !

-

0

0

0

0 0

0

0’

0 0

0 0

0 0

0 0

0

0

0

0

0

0

0

0

0

0

0

305

0

0

0

0

0 0

5 km

A

LE

GE

ND

A

LP

UJA

R R

I DE

CO

MP

LE

X

MU

LH

AC

EN

CO

MP

LE

X

TE

CT

ON

IC

CO

NTA

CTS

Hu

ert

ec

ico

s A

lto

s- - Dx-l

’ - Dx+

l A

lmo

caiz

or

Un

it - Qx+

~

Dx

+~

V

ori

ego

to U

nit

- - - -

- - -

I]

Qu

ote

rno

ry

se

dim

en

ts

v]

Up

pe

r N

eo

ge

ne

s

ed

ime

nts

E-1

Alm

on

zora

Unit

Mo

coel

-Ch

lve

Un

it - h

igh

an

gle

fa

ult

0

00

Up

pe

r N

eo

ge

ne

v

olc

on

ics

N

ev

ad

a-L

ub

rin

Un

it

pb. 2

. G

eolo

gica

l map

of

the

east

ern

Sier

ra d

e 10

s Fi

labr

es, a

fter H

elm

ers &

Voe

t (1%7),

Kam

psch

uur (

1975

) and

map

ping

by

H.E

.B. and K

.d.J

..

sz H. E. BAKKER Er A L .

which show strong stratigraphic affinities with Triassic rocks of the (Subbetic) External Zone (Kozur, Mulder- Blanken & Simon, 1985). Therefore these Almagride rocks can be interpreted as windows of the External Zone.

The lowermost of these thrust units is the Almanzora Unit. On the basis of lithostratigraphy and geochemical characteristics this unit shows close affinities with the thrust sheets of the Mulhacen Complex. However, because it has not suffered the same deformational and metamor- ph~c history during the alpine orogeny, it is regarded as a separate tectonic element.

The purpose of the present study is to establish the tectono-metamorphic history of the Mulhacen Complex and the Almanzora Unit and to discuss its implications for the geodynamic evolution of the Internal Zone of the Betic Cordilleras. For this reason detailed structural and meta- morphic analyses have been carried out in the eastern Sierra de 10s Filabres (Figs 1 and 2). The tectonic units in this area show a complicated small-scale structural ge- ometry produced by the superposition of several gener- ations of folds and foliations. The deformation phases are labelled D,-,, D,, D,,, etc. with the main deformation phase D. coinciding with a phase of regional metamorph- ism. In order to establish the physical conditions prevailing during the successive phases of deformation, extensive electron microprobe analyses were performed and the metamorphic reactions and time relations between mineral growth and deformation have been studied, using criteria developed by Zwart (1%2), Spry (1969) and Misch (1%9).

GEOLOGICAL SETTING A N D LITHOLOGY OF THE THRUST UNITS

The Sierra de 10s Fdabres forms the eastern extension of the Sierra Nevada. The topography of the mountain range is controlled by an east-west trending antiformal structure, dipping eastwards below essentially post-orogenic sedi- ments of Late Miocene and younger age (Volk, 1967; Voet, 1%7; Weijermars, Roep, Van den Eeckhout, Postma & Kleverlaan, 1985; Montenat, Ott d’Estevou & Masse, 1987).

In the investigated area the Mulhacen Complex comprises three thrust sheets. In ascending order they are the Nevado-Lubrin Unit (Nijhuis, 1964). the Macael-Chive Unit (Kampschuur, 1W5) and the Huertecicas Altas- Almacaizar Unit (Helmers & Voet, 1%7).

The lower part of the Nevado-Lubrin Unit consists of an alternation of dark-coloured mica schists and quartzites. It is covered by a Triassic sequence of schists and overlying marbles with intercalated schists, quartzites and locally gypsum. This sequence contains extrusive and effusive mafic and (minor) ultramafic rocks. During the alpine deformation and metamorphism most mafic rocks were transformed into amphibolites with a marked differentiated layering. Only locally the original igneous textures are preserved. Both the igneous and volcanic rocks show WPB-affinities (Pearce & Cann, 1973).

Age diagnostic fossils have not been preserved. However, based on striking lithostratigraphical and

geochemical similarities, the carbonate sequence has been correlated with Late Triassic carbonate series of the Almanzora Unit.

A similar lithostratigraphical sumssion is present in the overlying Macael-Chive Unit and the Huertecicas Altas- Almocaizar Unit. In these units the Triassic succession of clastics and overlying carbonates rests upon a suite of Palaeozoic rocks. In the Macael-Chive Unit this suite comists of a monotonous sequence of graphite-bearing mica schists with quartzite and marble intercalations, which has been intruded by granite. The intrusion was associated with skarn formation in adjacent carbonate rocks (Helmers, 1982). Due to severe alpine deformation the granite has been transformed into augen and even-grained gneisses. The 269Ma isochron age (Priem, Boelrijk. Hebeda & Verschure, 1966) for the metagranite indicates a pre-Early Permian age for the country rock.

The Palaeozoic sequence of the Huertecicas Altas- Almocaizar Unit is largely composed of darkcoloured biotite, epidote and amphibolite-rich gneisses, often with augen structures. The gneisses are associated with graphitic mica schists and quartzites and locally with marbles and calc-silicate rocks.

The thrust units of the Mulhacen Complex represent the oldest alpine thrust sheets which have been recognized in the area. The original nature of the contacts has been largely obliterated during subsequent defomation when the contacts were isoclinally folded and sheared. On the outcrop scale the contacts are parallel to the transposed foliation. Regional mapping, however, showed that the basal contact of the Macael-Chive Unit truncates lithological units of the underlying Nevado-Lubrin Unit.

The thrust units of the Mulhacen Complex are overlain by the Almanzora Unit. The contact is essentially a ductile shear zone, which transects the transposed lithological layering of the Mulhacen Complex on a regional scale. This contact has been reactivated during later deformation phases, especially in the eastern part of the area (Fig. 2).

The lower part of the Almanzora Unit consists of an alternation of quartzites and phyllites. The higher parts are characterized by carbonate rocks with locally intercalated phyllites and abundant gypsum with massive carbonate rocks at the top of the sequence. Metabasic rocks are present within the gypsum horizons. Pre-Triassic rocks are not present in the Almanzora Unit. An Early Kamian age can be attributed to the upper part of the carbonate sequence (Kozur, Mulder-Blanken & Simon, 1985).

DEFORMATION STRUCTURES

In the MulhGn Complex the main deformation structure is a D, transposition foliation (Sx). The foliation is a differentiated quartz-mica layering with a spacing of one millimetre or less. S, foliations enclose intrafolial folds, budins, augen structures and contain a pronounced striping and mineral lineation. Boudinaged crystals indicate that the lineation was formed parallel to an east-west to north-west-southeast trending stretching direction (Fig. 3b). Quartz c-axis fabrics demonstrate

N

linea

tion

0

Glu

coph

am L

imnt

ion

OX

-1 M

ULH

AC

EN

CO

MP

LEX

0

9-1

fol

iatio

n

N

+

Lx.1

itr

etch

ing

linoa

tion

0

SX

folia

tion

Dx

MU

LHA

CE

N C

OM

PLEX

N

I

t* C?O

sxts

nrio

nal a

mu

lati

on

ch

ap

D

~v

l MU

LHA

CE

N C

OM

PLE

X

0

SX.1

fo

liatio

n h

.2 M

ULH

AC

EN

CO

MPL

EX

linsi

tion

Dx

ALM

AN

ZOA

A U

NIT

N

lim

tio

n

0

stm

nd

mw

mtu

bs

o Sx

folia

tion

N

o

min

folia

tion

Dx.

3 M

ULH

AC

EN

CO

MP

LEX

Rg.

3.

Ster

eogr

ams (

low

er h

emisp

here

pro

ject

ion)

showing

stru

ctur

al el

emen

ts from s

elec

ted

area

s fo

r def

orm

atio

n ph

ases

D,-, to

M

acae

l-Chi

ve U

nit s

outh

of

Lubr

in; g

lauo

opha

ne lin

eatio

ns fr

om a

mph

ibol

itcs,

Ntv

ado-

Lub

rin

Uni

t nor

ther

n Si

erra

de

10s

Fdab

rcs.

(b) 4

from q

uart

zite

s, m

ica

schi

sts

and

mar

bles

, Ncv

ado-

Lubr

in U

nit n

orth

ern

Sier

ra d

e 10

s Fi

labr

es. (

c) L., from q

uart

zite

s, ph

yllit

es and marbles, A

lmm

ra U

nit n

orth

-eas

tern

Sier

ra dc

10s F

ilabr

es. (

d), (

e) a

nd (

f)

From m

arbl

es, q

uart

zite

s an

d m

ica

schi

sts,

Nev

ado-

Lubr

in U

nit n

orth

ern

Sier

ra d

c 10

s Fd

abre

s.

(a) S

tret

chin

g lineations i

n gn

ciss

es,

364 H. E. BAKKER ET AL.

dominant noncoaxial strain with west to north-west directed shear during D,. Locally S, is parallel to the axial plane of centimetre to metre scale, tight to isoclinal similar folds. The widespread Occurrence of folded quartz veins indicates that D, was preceded and accompanied by the formation of tension gashes and solution transfer processes.

S. is penetratively developed in all lithologies with the exception of some mafic bodies, which have either undeformed cores or contain a penetrative older S._, foliation. In gneiss bodies becomes progressively folded towards the contact zone with the metasedimentary country rock, where the gneisses contain a penetrative Sx foliation.

Across the contacts between the Mulhacen thrust slices no D, strain or metamorphic gradient is present. It is therefore concluded that the main thrust contacts within the Mulhacen Complex were formed prior to D,.

Gneissic, platy quartzite and mica foliations which are folded in the hinge zones of D, folds indicate that D. was preceded by D,_, deformation (Fig. 4a). Boudins, wrapped by S, contain an S,_, fabric and are characterized by minerals from the high-pressure assemblage. In amphibol- ites D,-I structures appear to be only locally preserved in relics. They comprise platy glaucophane-epidote-mica foliations, intrafolial folds, mineral and striping lineations, extensional crenulation cleavages (Platt & Vissers, 1980), boudins and tension gashes filled with quartz and aragonite. During formation of S,_, randomly orientated glaucophane sheaves were progressively rotated into a pronounced WNW-ESE trending mineral lineation (Fig. 3a).

The central part of the Macael-Chive gneiss body south of Lubrin (Fig. 2) contains a penetrative S.-, foliation with a WNW-ESE directed stretching lineation (Fig. 3a). Asymmetric tails around feldspar porphyroclasts indicate WNW-directed shear (Fig. 4b). In mica schists evidence for pre-D. deformation is, apart from local mica foliations within microlithons, usually restricted to the internal fabrics within pre- and synkinematic D, minerals. In these minerals internal S,_, foliations are constituted by boudinaged epidote crystals, graphite trails or shape fabrics of opaque minerals.

Dx-, structures have been observed at different levels in all three tectonic units of the Mulhacen Complex. It is therefore concluded that D,-, was also a phase of penetrative deformation. Differences in intensity of Dx-, deformation along and across the thrust sheet contacts within the Mulhacen Complex can not be established because D,-, structures were severely overprinted by the later phases of deformation.

Main phase D, structures were overprinted by a group of D,, , structures including isoclinal similar folds, mylonites, ultramylonites (in carbonate rocks) with approximately east-west trending mineral and stretching lineations, extensional crenulation cleavages (ECCs) and foliation boudinage structures (Fig. 4c). In general the ECCs form conjugate sets and show a variable spacing from one to several centimetres. They are present mainly

within the mica schists. On a regional scale the ECCs are prominently

developed in distinct zones. The most pronounced zone is located just below the contact with the Almanzora Unit. At this thrust contact the ECCs are developed in conjugate sets which define the most prominent planes in the rocks. Downwards, away from the contact, the ECCs gradually become broader spaced and less prominent.

On a regional scale the basal thrust contact of the Almanzora Unit truncates the S . foliation within the underlying Mulhacen Complex. This observation and the increasing intensity of D,,, ECCs towards the thrust contact indicates that the contact Mulhacen Complex- Almanzora Unit was formed during D,,,. The orientation of mineral and stretching lineations indicate east-west directed shear along this contact (Fig. 3d). Quartz c-axis fabrics demonstrate Edirected shear during D,,,.

Within the Almanzora Unit Dx+, structures have only been observed directly above the contact with the Mulhacen Complex. The structures in this zone are open to isoclinal folds, transposition foliations, intrafolial and sheath folds, mylonites and associated east-west trending mineral and stretching lineations, ECCs and foliation boudinage structures. The D,,, structures have over- printed the main phase structures in the Almanzora Unit which comprise transposition foliations (S,) with associated north-west-southeast trending striping lineations, sheared worm tubes (Figs 3c & 4d), boudinaged bedding planes and open to isoclinal folds. Penetrative amphibole-epidote S, fabrics have been locally developed in thin mafic bodies and locally in the contact zone with the metapelitic country rock. Folded quartz veins indicate that the main deforma- tion phase was preceded and accompanied by the forma- tion of tension gashes and solution transfer processes.

The main deformation structures indicate penetrative non-coaxial deformation of the Almanzora Unit with a north-west-southeast stretching direction. Based on their pre-D.,, age, comparable style, similar orientated stretching direction and metamorphic characteristics these structures are correlated with D. structures in the Mulhacen Complex (Fig. 5). Evidence for pre-D, deformation is only derived from microscopic observa- tions. Blue-green amphibole crystals, which constitute S., contain cores of crossite and glaucophane with an internal fabric of epidote s l . , rutile and opaques. This implies the existence of D,-, deformation.

During D,,, the previously formed transposition folia- tions were folded and the rocks were affected by south to south-west directed large-scale reverse faulting and over- thrusting. These structures are typically observed in the Mulhacen Complex.

The character of the Dx+Z deformation is strongly associated with the lithostratigraphic position. In the clastic sequence of the Nevado-Lubrin Unit D,,, deformation was penetrative and exclusively ductile. The structures here comprise inclined to upright open to very tight flexural slip folds, controlled by a broad to closely spaced crenulation cleavage (S,,,). In the overlying sequences ductile Dx+, deformation was restricted to thrust

GEODYNAMIC EVOLUTION OF THE BETICS 365

I I

Dx+

3

MU

L H

AC

EN

C

OM

PL

EX

I I

open

to ti

ght f

olds

, ax

ial p

lane

cle

avag

e, e

xten

sion

al fa

ults

,

wer

thru

sts,

rev

erse

and

tear

fau

!ts,

britt

le-d

uctil

e sh

ear

zone

s

N -

NE

war

d te

cton

ic tr

ansp

ort

t I I

Dxt

op

en to

tigh

t fo

lds,

axi

al p

lane

cle

avag

e

reve

rse

faul

ts,

over

thru

sts

S ~

SW w

ard

tect

onic

tran

spor

t

t I I

D~

-

I F

orm

atio

n of

Thr

ust S

heet

s in

Mul

hace

n C

ompl

ex

trans

pose

d fo

liatio

ns, i

ntra

folia

l fol

ds,

min

eral

and

str

etch

ing

linea

tions

, ext

ensi

onal

cre

nula

tion

clea

vage

WNW

war

d te

cton

ic tn

nrp

ort

I 1 I

isoc

linal

fold

s, t

rans

pose

d fo

liatio

ns,

(ultr

a) m

ylon

ites,

min

eral

and

stre

tchi

ng li

neat

ions

, ext

ensi

onal

cre

nula

tion

clea

vage

E w

ard

tect

onic

tran

spor

t ox

+ 1

t I I

DX

tig

ht t

o is

oclin

al fo

lds,

tran

spos

ed fo

liatio

ns

min

eral

and

str

etch

ing

linea

tions

W - N

W w

ard

tect

onic

tran

spor

t

AL

MA

NZ

OR

A

UN

IT

open

fold

s, a

xial

pla

ne c

lear

age

reve

rse

faul

ts a

nd o

vert

hrum

N - N

E w

ard

tect

onic

tran

spor

t

open

to

tigh

t fo

lds,

axi

al p

lane

cle

avag

e

~~

Em

plac

emen

t on

Mul

hece

n C

ompl

ex

open

to is

oclin

al fo

lds,

tra

nspo

sed

folia

tions

, m

ylon

ites,

min

eral

and

stre

tchi

ng li

neat

ions

, ext

ensi

onal

cre

nula

tion

clea

vage

E - W

str

etch

ing

dire

ctio

n

tight

to

isoc

linal

fold

s, b

oudi

nage

stru

ctur

es, s

hear

ed w

orm

tube

s,

trans

pose

d fo

liatio

ns,

min

eral

and

str

etch

ing

linea

tions

NW -

SE s

tretc

hing

dire

ctio

n

inte

rnal

mic

ro fa

bric

in g

lauc

ophr

nr a

nd c

rolt

itl c

ryltr

ls

a. 5. C

orre

latio

n ta

ble

show

ing

the

defo

rmat

ion

sche

mes

for

the

Mul

hace

n C

ompl

ex a

nd A

lman

zora

Unit.

GEODYNAMIC EVOLUTION OF THE BETICS J7

3E) H. E. BAKKER ET AL.

zones, in which a sharp increase 'in prominence of Dx+2 occurs. Open south to southwest vergent flexural slip folds (Figs 3d and 6a) become flattened and concurrently the spacing of the axial plane crenulation cleavage is strongly reduced. Near the thrusts the folds are tight to isoclinal and a closely spaced crenulation cleavage or differentiated layering is present. The thickness of these zones is a few tens of metres.

The sharp contrasting amount of flattening between the clastic sequence of the Nevado-Lubrin Unit and the overlying imbricated calcareous series suggests that an important detachment zone was present between both sequences during Dx+,. This interpretation is supported by the Occurrence of prominent south to south-west vergent D,+* folding just above and below this contact (Fig. 6b). Metre-thick chlorite-bearing quartz and carbonate veins along the thrusts indicate pronounced solution transfer at a macroscopic scale.

During D,+, the previously formed structures were overprinted by a group of structures including open to tight folds with associated axial plane cleavages, breccia zones, low and high angle extensional faults, reverse and tear faults, overthrusts and locally brittle-ductile shear zones with associated ultramylonites. Stretching lineations and asymmetric strain shadows in carbonate ultramylon- ites, slickensides, the orientation of SX+, cleavages and rotated older structures suggest a north to north-east directed tectonic transport (Fig. 30.

The folds are characteristically disharmonic with wave- lengths varying from one to one hundred metres. They have been preferentially formed within the contact zone of the clastic and carbonate sequences. S,+, cleavages are only prominently developed in the top of the clastic sequence of the Nevado-Lubrin Unit where they are locally developed as penetrative closely spaced crenulation cleavages.

The breccia zones are mainly developed in carbonate horizons, in particular in the basal part of the Nevado- Lubrin carbonate sequence. The zones originate as thin horizons parallel to the main foliation. During progressive deformation they extended and transected the main folia- tion. Large breccia wnes commonly show a lens shape which may attain lateral magnitudes of several kilometres.

Brittle-ductile shear zones have been observed in mica schists and carbonate rocks. In the carbonate rocks ductile and brittle deformation occurred intermittently.

The conspicuous D,+, strain gradient in the top of the clastic sequence of the Nevado-Lubrin Unit indicates that the D,+, detachment zone at this contact was reactivated during D,,,. In the Almanzora Unit Dx+, structures are most prominently developed in a zone in the northeastern part of the area where locally the Mulhacen Complex has overthrusted this unit.

METAM 0 RPH I S M

In 1% Nijhuis presented the first detailed description of the plurifacial metamorphism of the rocks of the Mulhacen

Complex for a small area south of Lubrin in the eastern Sierra de 10s Filabres. Since then his metamorphic scheme has been demonstrated to be valid, with minor additions, in most of the central and eastern Sierra de 10s Filabres (Langenberg, 1972; Kampschuur, 1975; Vissers, 1981).

In this section the most significant metamorphic parageneses are described in chronological order. Because the stacking of the tectonic units of the Mulhacen Complex has occurred before D, and no significantly different pre-D. mineral parageneses have been recognized between the three units, the metamorphic evolutions of the different units will be described together. The relationship between metamorphism and deformation is shown in Fig. 7a and b. So far only limited geothermobarometry has been

published (Helmers, 1983). For this study over 500 electron microprobe mineral-pair analyses were used to establish the P-T-t path of the Mulhacen Complex and the Almanzora Unit (Fig. 8). Electron microprobe analyses were performed with a Cambridge Instrument Co. Microscan-9 with an accelerating voltage of 20kV. Analysed mineral samples were used from standards and apparent concentrations were corrected using the online ZAFcorrection program. For further details of the microprobe data H. Helmers can be contacted at the Free University in Amsterdam.

Pre D,-, metamorphism

The first phase of alpine metamorphism is characterized by incipient eclogitization of dolerites and garnet- hedenbergite skams, by crystallization of quartz-jadeite, quartz-aragonite pairs and by the formation of albite- and epidote-bearing omphacitites at the contacts of mafic igneous rocks and carbonate layers.

In slightly deformed massive dolerite bodies incipient eclogitization is indicated by the development of garnet between magmatic plagioclase and clinopyroxene and by the formation of omphacite and rutile from augitic pyroxene. Eclogitization of skarn bodies mainly affected the outer zones. The central parts are usually still composed of the original skarn mineralogy: acmitic hedenbergite, intermediate grossular-almandine, titanite and apatite (Helmers, 1982). In boudinaged quartz-jadeite veins jadeite (Jdn-,,,,) is present as unorientated crystals, which are always separated from quartz by a later symplectitic rim of acmitic pyroxene and albite with inclusions of hematite and paragonite. Metagranites and gneisses locally contain omphacite relics enclosed in garnet and jadeite-aanite, low-albite, Si-rich phengite (3.52) and Al-rich titanite. In mica schists only pre-D.-, epidote crystallization has been recognized.

Temperature conditions prior to D.-, are indicated by garnet-omphaate K,, values (Rheim & Green, 1974; Ellis & Green, 1979), ranging from 41 to 24 and K,, values of garnet-phengite pairs (Krogh & Raheim, 1978; Green & Hellman, 1982) varying between 27 and 21.5. For the determination only compositions of unorientated inclusions of omphacite and phengite are used with compositions in

OM

PH

AC

ITE

GA

RN

ET

GLA

UC

OPH

AN

E S.

1

EPI D

OTE

CLl

NO

ZOlS

lTE

AM

pHIB

oLE

(blu

o9

AL

BilE

OLI

GO

CLA

SE

AC

llNO

LlTE

CH

LOR

ITE

BIO

TITE

OPA

QU

ES

TlTA

NlT

E

RU

TILE

TOU

RM

ALI

NE

AR

AG

ON

ITE

CALc

mlW

LoM

T"

M K

1A(co

tour

la

d

EP

lDO

TE

KYA

NlT

E

CH

LOR

ITO

ID

CH

LOR

ITE

BIO

TIT

E

LAZU

LlTE

A

LBIT

E

OL

looC

LA

sE

TlTA

NlT

E

OPA

OU

ES

QU

AR

TZ

MIC

Ab

Iou

rla

sd

m.7. R

elat

ive

timin

g of

def

orm

atio

n pha

scs a

nd m

etam

orph

ic cr

ysta

lliza

tion.

(a)

Mul

hace

n C

ompl

ex, m

ica

schi

sts a

nd g

neis

ses.

(b) M

ulha

cen

Com

plex

, maf

ic ro

cks.

(c)

Alm

anzo

ra

Uni

t, ph

yUite

s and

mic

a sc

hist

s.

3n H. E. BAKKER €T AL.

1

L

c

*

4 P

(kbar) f

E

7

0

9

10

1 1

T ("C) 300 4 00 500 600

10

Depth (km)

?O

30

10

b HIGH- PRESSURE^ Flg. 8. P-T-t paths for the Mulhaccn Complex (light shading) and the Almanzora Unit (dark shading). (1) albite = jadeite + quartz (Newton & Kennedy, 1968); (2) rutile + almandine = ilmcnite + kyanitc + quartz (Bohlen el d.. 1983); (3) aragonitein (Boettcher & Wyllie, 1%7); (4) glaucophanc + chloritoid = paragonite +chlorite (Kicnast & Triboulct, 19n); (5) glaucophanc stability (Maresch,

CEODYNAMIC EVOLUTION OF THE BETICS 311

the garnet core adjacent to the inclusions. Equilibration is indicated by the limited spread of values.

Pressure indicators are derived from jadeite(J&=)- albite pairs (Currie & Curtis, 1976; Kushiro, 1969; Holland, 1980), jadeite-quartz pairs (Newton & Kennedy, l W ) , rutile-almandine (Bohlen, Wall & Boettcher, 1983) and the presence of aragonite (Boettcher & Wyllie, 1%7).

These data indicate (Fig. 8) conditions of 350-400"C at 9.5-10.5 kbar. Taking into account the Cacontents of garnet (usually approximately Gr,) a slightly higher temperature range of 4oo-46o"C is obtained. This is appropriate, due to the non-ideal mixing character of Ca-substitution in garnet. Si-values of phengites (Mas- sonne & Schreyer, 1983) from undeformed metagranite (which contains K-feldspar) cluster around 3.52. Their stable occurrence against jadeite and jadeite-acmite and/or low-albite, in absence of garnet, suggests initial tempera- tures below 350°C in approximately the same pressure range. The P-T box based on these data is separately indicated in Fig. 8 at the beginning of the metamorphic path.

Complete absence of lawsonite crystals or pseudo- morphs in this early high pressure assemblage indicates the presence of a partial CO, pressure (Nitsch, 1972). This conclusion is supported by the presence of carbonate veins with aragonite relics. A low P is indicated by abundant development of anhydrous mnerals and absence of glaucophane s.1. and phengite in mafic rocks during this stage.

"P.

Synkinematic D,-, metamorphism

In mafic rocks synkinematic D.-l metamorphism is in general characterized by an increasing Ph0. Growth of glaucophane s.f., garnet, epidote s.f., pure low-albite, paragonite and locally omphacite has occurred. In boudinaged carbonate veins aragonite pseudomorphs have been found.

Glaucophane has been formed in necks of boudinaged omphacite crystals (Fig. 6c). Locally glaucophane has been crystallized to omphacite. At other localities however eclogitic parageneses have remained stable. Zonal omphacite prisms growing from wall to wall indicate rapid growth in open 0uid-Ned veins discordant to s-,. This indicates varying and Buctuating Pw conditions.

The epidote crystals are commonly zoned with a gradually decreasing Fecontent towards the rim, locally reaching a clinozoisite composition. The Fe-rich (b) cores contain equidimensional, random orientated opaques. Towards the Fe-poorer rims a progressively stronger internal fabric is defined by stretched and folded opaques.

This observation indicates that the epidote composition changed during Dx-l. Synkinematically formed clinomisite locally fills the necks of Dx-l garnet boudins.

In tourmaline gneisses stretched phengite-epidote- garnet aggregates crystaked at the expense of magmatic biotite. In mica schists synkmematic formation of glauco- phane (Fig. a), phengite, epidote, chloritoid, Ti- hematite and locally garnet. has been observed. In mica schists and amphibolites chloritoid-glaucopbane pairs are present as arrnoured relics in garnet.

The K,,-values of garnet-clinopyroxene pairs (18 to 14) of schistose eclogites, garnet-phengite pairs (14.5 to 12) in gneisses and the Jd-values of omphacites (45 to 35) point towards temperatures of 475-525" C at 9-11 kbar. Taking into account the Cacontents of the garnets (up to Grm) a temperature range of 485-540"C is indicated (Fig. 8). However, we question the uppermost temperature because it is calculated using the outer rim of the garnet porphyroblast, which may have grown post-Dx-l. An increase in temperature during D,-l is indicated by the decreasing Fecontent in epidote s.1. towards the rim, locally reaching clinozoisite composition.

In the Miyashiro-F'RGM-diagram glaucophane and crossite compositions near to the glaucophane S.S.

end-member indicate high pressure (Carman & Gilbert, 1983) and a temperature below 550°C (Maresch, 1977). The relative large gap in Jdcontent of clinojyroxenes (Fig. 9) is in accordance with the pressure indicated by the geobarometers (Mottana, 1983). Further high-pressure indicators are Mg-rich chloritoid (Fig. 10; Bearth, 1963; Ganguly, 1972; Chopin & Schreyer, 1983), almandine- rutile intergrowth, aragonite stability and the Si-values (clustering around 3.37) and b,,-values of phengites (Fig. 11).

Synkinematic D, metamorphism

Synlunematic D. reactions in ma6c rocks are characterized by a further increase in P H p . Omphacite, glaucophane and crossite have recrystallized to blue-green amphibole, usually accompanied by albite growth. This recrystalliza- tion process has been nearly complete in rocks which contain a penetrative S, foliation. S. is further constituted by clinozoisite, zoisite and paragonite. Garnet and plagioclase (albite and oligoclase) OCCUT as synkinematic porphyroblasts. For the plagioclase both increasing and decreasing Ancontents towards the rim have been observed (Ans,). Aragonite, occurring in carbonate veins in mafic rocks, has recrystallized to calcite together with blue-green amphibole growth at the expense of omphacite.

In mica schists the D, mineral assemblage consists of

Fig. 8. (Continued.) 19nk (6) staurolite-in (Hoschek. 1969): (7) antieorite = forsterite + talc + H,O (Evans et d.. 1976): 18) aoorthite + H,O = &tC+ misite + q& (New& & K e k d y , i963); (9) barroisite stability-(E&t, 1979); (10) Ah&tc triple poini(Holdaway, 1971); (11) stilpnomelane + phengite = biotite + chlorite + quartz (Nitsch, 1970;) (12) pyrophyllite = Al-silicate + quartz + H,O (Chatterjee et d., 1984). Si-values according to MWMC & Schreyer (1983). Jd-percentages according to Currie & Curtis (1976). Boxes indicate the established P-Tconditions for each deformation phase D. Pre-D,-,-A of the Mulhacen P-T-t path points to the precclogite stability box, prc-D,-,-B to the early dogite stability box. Rc-D,-A of the Almanzora P-T-t path indicates the mlourttss amphibole pressure box en route to pn-Dx-B, the blue amphibole stability.

372 H. E. BAKKER ET AL.

A I

. ,* .+ : . * \

\ ' / -

-6

-4

-2

Fig. 9. Compositions of diopside, acmite-jadcite, omphacite and jadeite in central and eastern Sierra de 10s Filabres. Calculation of end-members from electron microprobc a n a l p by method of Papike et ul. (1974). Augite end-members always a diopside-hedenbergite solid solution with less than 1% wollastonite. Ca-tschcrmakite or cnstatite. Arrows indicate core and rim compositions in one crystal. Below: determination of Jd-contents by method of Esscne & Fyfe (1967) using the value of (221) re&ction of X-ray photographs. Both methods agree within 5% and frequently within 3%.

phengite, quartz, chloritoid. kyanite, staurolite, almandine-rich garnet and (clino)zoisite. Locally blue- green amphibole or taramite (Linthout & Kieft, 1970) replaces glaucophane. Staurolite (Fig. 10) contains a considerable amount of Zn and Mg. In the absence of biotite the staurolite is probably formed from kyanite+ chloritoid. This reaction indicates a decreasing H,O content. Its occurrence is restricted to graphite-bearing mica schists surrounding metabasite bodies or ortho-

gneisses. The reason for this distribution pattern can probably be found in a lower PHI,, and Ps in the immediate surroundings of the former igneous bodies.

In calcitic and dolomitic marbles the parageneses tremolite + misite and colourless mica + quartz + Mg-rich chlorite (penninite) have been formed. This indicates a low P? in these rocks. In ultramafic rocks chlorite+ anbgorite are stable. Neither forsterite nor pseudomorphs after this mineral have been observed. In gneisses taramite

CEODYNAMIC EVOLUTION OF THE BETICS 373

Tzn+++Mn++ t

I I Mg++ 0 20 50 80

Psg. 10. Composition of chloritoid (circles from mica schists; triangles from amphmolltes; crosses from Almanzora mica schists) all low in Mn2' and with up to 4% Fe3+ in Mg-rich members, and of staurolite (points) all low in MI*+. Arrows indicate change in composition to the rim, tie-lines are dram between coexisting doritoid and staurolite.

was locally formed from jadeite-acmite, concurrently with the recrystallization of K-feldspar. High Si-phengite has remained stable.

According to the Brown (1977) diagram the composition of taramite, in an ill-defined part of the diagram, and of blue-green hornblende indicate a pressure of 7 kbar (Fig.

gar, 9.Ox) 9.020 9.030 9.042 9050 *coo

Fig. 11. Cumulative frequency lines of b,-values of phengite according to methad of Sassi & Scolari (1974). (1) mica schists of the Mulhac.cn Complex in the Sierra de 10s Fiabres (S,,,); (2) mica schists of the Mulhaccn Complex in eastern Sierra de 10s Fiiabres (S,); (3) mica schists of the Mulhaccn Complex in the Central Sierra de 10s Filabres (SA; (4) mica schists of the Almanzora Unit (Sx); (5) phyllites of the ALmanzora Unit (S,); (6) tourmaline gneisscs (Sx-,) of the Mulhacen Complex, castern Sierra de 10s F'iiabres.

12). Kelyphitic rims of blue-green amphibole and albite around omphacite (Jd-) point to a maximum pressure of 9 kbar. In the marbles relatively high pressures are indicated by the presence of tremolite+zoisite (Franz & Spear, 1983). In the mica schists high pressures are indicated by the moderate Mg-content of chloritoid (Bearth, 1963; Ganguly, 1972; Chopin & Schreyer, 1983) and by the association kyanite+zoisite (Newton & Kennedy, 1963) in absence of margarite. The occurrence of kyanite+ilmenite+quartz (Bohlen et al., 1983) indicates a lower pressure and a higher temperature than during D,- , .

The garnet-hornblende geothermometer (Graham & Powell, 1984) indicates temperatures of 535-595"C, using the rim of synlunematically grown garnet porphyroblasts. The formation of staurolite (Hoschek, 1969) indicates a minimum temperature of 560°C at 8kbar pressure, disregarding its Mgcontent of 12-27%. A maximum temperature of about 580°C can be inferred from the absence of forsterite in ultramatic rocks (Evans, Johanna, Oterdoom & Trommsdorff, 1976), never containing brucite.

Zoning profiles of syn-D, garnet (Lmthout & Westra. 1968), showing an increasing Mg/Fe ratio and a decreasing Mn/Fe ratio from core to rim, suggest an increasing temperature (Thompson, 1976) during D,. The b,,-values of phengite (Guidotti & Sassi, 1976) constituting S, (Fig. 11) indicate a change in gradient to intermediate temperature conditions. The Sicontent of these phengites ranges from 3.24 to 3.32.

The P-T conditions during D, are indicated in Fig. 8 at 8.25 kbar at about 570" C.

Metapelitic rocks containing chloritoid or pseudomorphs after this mineral intercalate at the decimetre scale with nonchloritoid-bearing rocks. This intercalation is believed to reflect chemical differences in the on@ sedimentary

374 H. E . BAKKER Er AL.

2.0

1.8

1.6

1.4

Na ML 1.2

1 .o

0.8

0.6

0.4

0.2

0

I I I I 1 I I I

\ GLAUCOPHANE S.L.

TARAMITE

-.

ACTINOLITE

I I I I

2 0.2 0:s ole 1 lo A l X 1.4 1.6 1.8 2.0 Fig. 12. Composition of over Z M electron microprobe analyses of amphiboles, plotted in the diagram of Brown (19n) . Analyses recalculated on 0, to an uncharged formulae manipulating Fe2+/Fe3+. Shaded areas contain analyses of blue-green amphiboles in kelyphitic intergrowths around omphacite. Arrows point to outer rims of crystals. Double lines indicate amphibole compositions in Ahanzora Unit metabasites.

layering. Minerals like chloritoid and kyanite are unable to Na-poor, chloritoid-bearing layers from neighbouring form in rocks with a high Nacontent (Hoschek, 1969). Na-rich layers, where the element has been stored within After D, but prior to Dr+,, chloritoid, kyanite and the mica lattice. This process reflects a decreasing staurolite have become unstable and have been altered to temperature (Cipriani, Sassi & Scolari, 1971; Guidotti & colourless mica (usually a mixture of paragonite and Sassi, 1976). In amphibolites Na-liberation by re- phengite), with or without chlorite. This non-isochemical crystallization of glaucophane to blue-green amphibole breakdown shows the introduction of Na into previously causes the breakdown of chloritoid.

CEODYNAMIC EVOLUTION OF THE BETICS 375

Synkinematic Dx+, metamorphism

During this phase no pronounced mineral blastesis has taken place. The metamorphism is characterized by non-isochemical retrograde reactions, which are spatially associated with D,+, extensional crenulation cleavages.

In mica schists chlorite was formed at the expense of phengite and garnet. In tourmaline gneisses biotite has formed during dephengitization. Locally D,+, tension gashes filled with albite have been formed in these rocks. In m a k rocks the most sipficant reactions include partial crystallization of blue-green hornblende to colourless mica, chlorite and locally actinolite and biotite.

The composition of rims of some amphiboles indicates a decrease in pressure to about 5-6 kbar (Fig. 12). according to the calibration of Brown (1977). A few outermost rims indicate a pressure decrease to about 4 kbar. The mutual presence of blue-green hornblende and, in other rocks, ripidolitic chlorite (locally pseudomorphing amphibole) indicates a temperature path that approximates the margin of the barroisite stability field (Ernst, 1979), to about 4 kbar and a temperature below 450" C. The combination of the hornblende-plagioclase (@) and garnet-phengite thermometers indicate 5.5 kbar and 505" C along this path. The pressure drop is also indicated by dephengitization. This trajectory reflects a post-D, cooling and decompression. The D,+, event is located at the lower P-T end of this path (Fig. 8).

Synkinematic metamorphism

During D,+, local blastesis of garnet, kyanite, chloritoid, biotite and stilpnomelane has occurred. The composition of the garnet has become richer in Mn and Ca and poorer in Fe and Mg (Linthout & Westra, 1968). The outer r ims of a few garnet porphyroblasts show a renewed increase in almandine content, with decreasing Ca and Mn (Lmthout & Westra, 1968). In the waning stages of this metamorphic phase however, all garnet have started to form chlorite (and quartz). Albite growth is striking in mica schists.

In mica schists pre- to syn-D, grown kyanite and chloritoid porphyroblasts have been partially crystallized into masses of paragonite, muscovite and chlorite. Colourless mica or albite pseudomorphs have locally been formed after these minerals. These reactions together with the local occurrence of albite coronas around chloritoid crystals indicate non-isochemical conditions. Albite and chlorite porphyroblasts have also grown at the expense of the main foliation mica. X-ray data indicate lower &-values for the &+, mica (Fig. 11).

In Fe-rich mafic rocks blue-green hornblende has remained stable. Locally major transformation into intergrowths of chlorite and albite together with minor epidote or d a t e is observed in Mg-rich types. In ultramafic rocks tremolite was decomposed into antigorite + carbonate and/or talc.

Locally stilpnomelane was formed (Nijhuis, 1%4), whereas normally chlorite and biotite growth has occurred. This indicates a temperature close to the upper stability

temperature of stilpnomelane (Nitsch, 1970). In combina- tion with the local kyanite growth of this stage, the P-T conditions are estimated at 425°C and 3-4 kbar (Fig. 8). The zoning profiles of garnet indicate that during D,,, a temperature decrease changed into an increase.

Synkinematic Dx+3 metamorphism

During D,+3 earlier grown albite porphyroblasts or aggregates have been surrounded by an anhedral rim of plagioclase (up to An,), suggesting a solid-fluid reaction type!. Small euhedral prisms of staurolite have nucleated in the colourless micachlorite reaction rims around syn-D, grown chloritoid, kyanite and staurolite or in chlorite- mica pseudomorphs after these minerals. Chlorite masses have partially been transformed into oxychlorite, along the cleavage or at the margins. Locally chlorite has been recrystallized into colourless mica or phlogopitic biotite. The Si-content of these micas is 3.07-3.10. In carbonate breccias and marbles euhedral plagioclase porphyroblasts show an increasing Ancontent towards the rim, with a maximum of An,. Locally sodic plagioclase-Na-rich scapolite pairs (b-Mei,) are present. A temperature of 510-525°C is indicated by the presence of Fe-rich staurolite (Hoschek, 1%9). Absence of forsterite limits the temperature to about 550°C (Evans et 1, 1976). The strongly varying plagioclase composition in combination with the absence of late zoisite, grossular and wollastonite indicates a comparable temperature range, in combination with a partial CO, pressure (Storre & Nitsch, 1972).

The stability of staurolite in the absence of cordierite indicates pressures above 2-3kbar (Richardson, 1968). This is also suggested by the stability of chlorite (clinochlore and ripidolite) against quartz in the absence of garnet (Hirschberg & Winkler, 1%) together with the decomposition of spessartine-grossularite-rich garnet dur- ing the waning stages of the preceding phase. A similar pressure is indicated by the Sicontent of Sx+3 micas (Fig. 8).

During Dx+2 and D.+3 mineral parageneses are developed locally and incompletely. Mineral reactions are restricted to zones of pronounced deformation. Deforma- tion may have offered the required activation energy to start up metamorphic reactions, but the distribution can also reflect local heating by hydrothermal processes. The activity of fluid migration during D,,, is indeed suggested by a significant Fe- and Mg-enrichment in fracture zones. So, in' our opinion, the maximum recorded temperatures for D,+, might have been reached in distinct zones only.

The metamorphic evolution of the Almanzora Unit

The Almanzora Unit shows its own distinct structural and metamorphic development. This paper presents the 6rst data on the relationship between deformation and metamorphism (Fig. 7c) and the geothermobarometry (Fig. 12) of the unit.

Bodies of blastophytic metabasites, wrapped by the main foliation S, in the metapelites, show hardly any

376 H. E. BAKKER ET AL.

deformation structures. The metamorphic minerals of the mafic rocks are albite, amphibole s.l., biotite, chlorite, epidote s.l. , garnet and phengite. These minerals generally OCCUT in a random fabric. Penetrative amphibole-epidote S, fabrics have been developed only locally. Blue-green amphibole crystals, which constitute this fabric, contain cores of glaucophane and crossite with an internal planar fabric correlated with S,-,.

The mafic rocks contain relics of magmatic minerals like sodic plagioclase, augitic pyroxene, and brown horn- blende. Blue and green amphibole crystals and aggregates are mainly formed from former plagioclase crystals. Locally they surround brown magmatic hornblende. The amphiboles display the following zoning pattern: magnesioriebeckite-crossite-blue-green amphibole. Sepa- rate aggregates of actinolite inside magmatic augite show an increasing content of Na and A1 towards the rims (Fig. 12, double arrow). According to the Brown (1977) diagram this composition points towards a pressure up to 5 kbar. Their presence inside magmatic augite, which is locally surrounded by blue and blue-green amphibole, suggests an early origin of the actinolite.

In the Miyashiro diagram for glaucophane s.1. the magnedoriebeckite shows a volume change value (Muir Wood, 1980) of about -2, and the crossite a value of -2.5. This is explained by an increase in pressure. The crossite, containing limited tetragonal Al, plot at 7 kbar (Fig. 12) in the diagram of Brown (1977). Blue-green amphibole indicates the same pressure, although their CI content up to 3.6wt% may influence this figure. Garnet contains up to 30% grossular. Its stability against the zonal blue amphibole in absence of omphacitic pyroxene indicates a temperature below W C , the minimum temperature for eclogite equilibrium. The garnet-phengite thermometer (Green & Hellman, 1982) indicates 440- 450" C at 6-8 kbar. The garnet-hornblende thermometer (Graham & Powell, 1984), applied to the rim of the garnet crystals and C1-poor amphibole, indicates 400°C. However, this value may still be affected by the Fe-enrichment effect of C1-content of the amphibole.

In the metapelites phengite, chloritoid, kyanite (ex- tremely rare) and epidote s.1. have been formed pre- to syn-D, (Fig. 7c). The Mg-content of chloritoid (4445%) points to a relatively high pressure (Fig. 10; Bearth, 1%3; Ganguly. 1972; Chopin & Schreyer, 1983). This is supported by the b,-values of S,-phengites (Fig. 11) and their Si-content of 3.27. This value closely coincides with the values for S, micas in the Mulhacen Complex. Epidote s.1. crystals lying with a body preferred orientation within S, become less Fe-rich towards the rim. This indicates an increasing temperature during D.. These characteristics, together with local growth of kyanite and biotite in the absence of stilpnomelane (Nitsch, 1970) are in accordance with the P-T conditions for D, derived from the geothermobarometry of the mafic rocks (Fig. 8).

After D, blastesis of chlorite and colourless mica occurred at the expense of kyanite and chloritoid. Often only pseudomorphs were left. The Occurrence of post-D. albite indicates Na-homogenitization and non-isochemical

conditions, as in the Mulhacen Complex. The Si-content (3.15-3.11) of phengites cross-cutting S. coincides with the lowest Si-values within the Mulhacen Complex during Dx+3. In mica schists and marbles late growth of plagioclase (max. An,,,, in part rimmirag albite) and biotite indicates a similar late temperature increase as ex- perienced by the Mulhacen Complex (Fig. 8).

DISCUSSION

The reconstructed P-T-t paths of the Mulhacen Complex and the Almanzora Unit are shown in Fig. 8. The HPILT conditions prior to D,-l indicate that the Mulhacen Complex has been subjected to a phase of rapid tectonic burial, to a depth of approximately 37 km. This process is considered to represent a phase of crustal underthrusting of relatively cold continental crustal material, which caused deformation of the original pattern of subhorizontal isotherms in the upper part of the lithosphere (e.g. England & Thompson, 1984). No deformation structures have been observed that document this tectonic process.

On the trajectory towards D,-I the P-T-r path indicates that the rocks were subjected to an increase in temperature at approximately constant pressure condi- tions. This pattern clearly illustrates that the Mulhacen rocks were heated while they were kept at constant depth.

These conditions continued during and after Dx-,, the oldest recognized deformation phase. This isobaric heating pattern is caused by the cessation of the rapid crustal-scale underthrusting, enabling the start of the restoration of the disturbed thermal structure of the lithosphere. The cessation of large-scale underthrusting is probably best explained by the property of the continental crust, provided that a crust-mantle detachment has been established, that it is only allowed to descend to a limited depth during a process of crustal doubling, because the forces associated with the underthrusting are balanced by the buoyancy of the underthrust segment.

The relationship between the penetrative D,-l deforma- tion and formation of thrust sheets within the Mulhacen Complex is completely obliterated due to severe D, overprinting. Therefore, it is not possible to establish whether the thrust sheets were formed prior to or during D,-l. However, we consider it likely that the formation of the thrust sheet contacts has been associated with penetrative deformation. Hence we propose that the contacts were formed during Dx-l and that the original tectonic transport direction has been parallel to the glaucophane lineation and stretching lineation in gneisses during west-north-west shearing. This D,-, pattern of imbrication within the Mulhacen Complex, with present thrust sheet thicknesses of about one kilometre, is explained by continued collision. Further crustal-scale underthrusting is hampered by buoyancy SO that the original underthrust segment starts to imbricate.

The temperature increase lasted up to D,. This prograde trajectory is characterized by the recrystallization of several HP minerals to intermediate pressure parageneses with a decrease in density of several per cent. Locally

CEODYNAMIC EVOLUTION OF THE BETICS 377

eclogites hydrated to glaucophane schists, but most often glaucophane has grown at the expense of magmatic minerals. This indicates that the underthrusted crustal segment, containing the Mulhacen Complex, was not transformed to a high-density eclogite segment. Subse- quently glaucophane recrystallized to intermediate pressure blue-green amphiboles. The main effect of this compositional change is the reduction of the HzO wt. % of the amphiboles, only giving rise to a small increase in density. However, liberation of HzO, also indicated by extensive quartz veining in mica schists before and during D,, leads to an increase in volume and therefore to a decrease in density of a unit rock volume. This density reversion might be a pre- requisite for uplift (cf. Richardson & England, 1979).

The P-T-r path for the Mulhacen Complex indicates that after D,-, the rocks were consistently brought to a higher level in the crust. From D,-, to D. the rocks were uplifted some 7 km while the temperature increased by an amount of 70°C. The P-T conditions indicate that D, took place a depth of 2&34km (Fig. 8). D, itself has been associated with an upward movement of the affected crustal segment towards a higher level in the crust. Because this phase proceeded under peak temperature conditions it has resulted in the most penetrative and homogeneous deformation. Structural data point to dominant noncoaxial deformation and suggest an important component of lateral transport. Quartz c-axis fabrics indicate west to north-west directed shear during D,.

Recently, several models have been put forward concern- ing the processes which may be responsible for the uplift of high-pressuremetamorphic rocks (e.g. Draper & Bone, 1980; Davy & Gillet, 1986; Platt, 1986, 1987). Platt has presented a model in which uplift is essentially the effect of extension in the upper rear part of a mechanically continuous wedge above an active subduction zone. An alternative model is that the rocks are actively

transported over the footwall. of a previously formed crustal-scale overthrust zone during D,. The rocks are brought to the surface by progressive crustal-scale footwall collapse combined with erosion during continued collision (Fig. 13a). Thus D, is considered as a continued imbrication of a crustal segment unable to descend because of buoyancy.

A similar process has been described by Davy & Gillet (1986) in the western Alps. They described a process of discontinuous crustal-scale stacking and transport towards higher crustal levels of previously underthrusted sheets. We suggest that D,-l and D, are related to a similar process and represent distinct pulses during progressive continent-continent collision.

The initial part of the P-T-t path of the Almanzora Unit indicates that this unit has also been involved in a process of continent-continent collision. This unit experienced lower maximum pressure and temperature conditions than the Mulhacen Complex. This indicates that the Almanzora Unit has been buried to a lesser degree and consequently has been located at a higher crustal level than the Mulhacen Complex before and during D,. The most penetrative structures were formed during D, under peak

temperature conditions. Because of the lower thermal conditions recrystallization has been less pervasive, com- pared with the Mulhacen Complex. Because of this a part of the burial process is still visible in the mineralogy of the malic rocks (Fig. 8).

Given the striking stratigraphic similarities of the lower part of the Almanzora Unit with the top of the Nevado-Lubrin Unit of the Mulhacen Complex, the Almanzora Unit might have formed part of the same major thrust sheet but has merely been buried to a lesser degree (Fig. 13a; cleaved arrow). However, the apparent thermal gradient for the Almanzora Unit is approximately twice the value for the Mulhacen Complex. This is more readily explained by assuming that the Almanzora Unit has initially been included into a higher thrust sheet (Fig. 13a; cleaved arrow). In any case the Almanzora Unit also formed part of a crustal-scale hanging wall during D,. The lower maximum P-T conditions, but higher thermal gradient and the heating after the initial burial can be explained by a combination of a screening effect of the underlying rocks (the Mulhacen Complex) and a heating by an overlying (hypothetical) thrust sheet (Fig. 13a), according to the models of Davy & Gillet (1986).

The subsequent post-D, trajectory in both crustal segments is characterized by a marked decompression under decreasing temperature conditions (Fig. 8). At the end of this retrograde trajectory the Almanzora Unit has been placed on top of the Mulhacen Complex. Since the Almanzora Unit was previously located at a higher crustal level compared to the Mulhacen Complex, this juxtaposi- tion must have involved heterogeneous extension and crustal thinning. The thickness of the excized crust amounts to approximately 6km. The very low angle of truncation of the S . in the underlying Mulhacen Complex, the orientation of stretching lineations and quartz c-axis fabrics demonstrate an eastward slip during D,+, extension (Fig. 13b).

The metamorphic data indicate that during and after D.+* a significant temperature increase took place (Fig. 8). The data show that locally a geothermal gradient of approximately 7" C/km has been established during DX+,. T h i s high geothermal gradient might have been the conse- quence of the crustal thinning process initiated during

In the western part of the Internal Zone, metasediments of the Alpujarride Complex have been intruded by ultramafic rocks, for which an K/Ar age of 22 Ma has been derived (Priem, Boelrijk, Hebeda, Oen, Verdurmen & Verschure, 1979). It is envisaged that this intrusion initially took place during a phase of crustal thinning, although the actual contacts indicate that the ultramafic rocks were involved in a later overthrusting event (Westerhof, 1975; Tubia & Cuevas, 1986). Correlation of the establishment of the high geothermal gradient in the eastern Sierra de 10s Filabres with this intrusion event might suggest an Oligocene to Early Miocene age for the D,,, to D,,, trajectory. A phase of crustal stretching followed by oceanic spreading during this period is well documented in the western Mediterranean (Alvarez, Cocozza & Wezel,

Dx+,.

(a)

D,

CONF

IGUR

ATIO

N

WN

W

AU

M

C

T--

r I

ES

E

- th

rus

t s

he

et

co

nta

ct

iso

the

rm

t t

t I

f v

ert

ica

l fo

rce

s d

ue

to

grl

av

ita

tio

na

l re

sp

on

se

I

(b)

D,,

CONF

IGUR

ATIO

N

W

T+

I

P 1

t

f f

D,+1

ex

ten

sio

na

l fa

ult

-

--- p

re D

x+

l th

rus

t sh

eet

co

nta

ct

iso

the

rm

ve

rtic

al

forc

es

du

e t

o d

ev

elo

pin

g

ma

ntl

e a

no

ma

ly

P P

f Q

, U. C

arto

on-li

ke s

ectio

ns i

llust

ratin

g cru

stal

str

uctu

re (

MC

: M

ulha

cen

Com

plex

, dur

ing:

(a)

D,;

(b) D,,,.

Inse

t P - T -

r pat

hs in

dica

te

ther

mal

evo

lutio

ns. A

U:

Alm

anzo

ra U

nit)

.

E

GEODYNAMIC EVOLUTION OF THE BETICS 379

1974; Bellon, Coulon & Edel, 1977; Rehault, Boillot & hhufiet, 1984). It is obvious that this process af€ected the whole continental crust and underlying mantle. This regional phase of extension and oceanic spreading also affected parts of the crust which have not been thickened considerably during alpine orogeny (e.g. the Balearic Islands and Sardinia). Therefore this extension is not related to the geometry and dynamics of a crustal wedge (Platt, 1986; 1987) but to processes within the mantle.

D,+, and D,+, structures, which reflect south to south-west vergent and north to northeast vergent large-scale overthrusting respectively, postdate the phase of extensional deformation in the Internal Zone. These structures have been imprinted on a crust that had locally been thinned wnsiderably and that was obtaining an irregular thermal structure. The simultaneous Occurrence of a locally developing high geothermal gradient, while extensional deformation has ceased, has been predicted by conductive thennomechanical models of rifting (Moretti & Froideveaux, 1986). These models show that once a thermal anomaly is generated, for instance by regional extension and crustal thinning, the further creation of a high geothermal gradient becomes a self-propagating process. So while the compressional D,+* and Dx+3 structures suggest crustal thickening the metamorphic data might reflect that crustal thinning has still been active during these phases.

The established pressure conditions (Fig. 8) indicate that the structures have been formed at a depth of about 10-12 km (D.+,) and 5-8 km (D,+J. In comparison to the older deformation structures these structures reflect more localized and more heterogeneous deformation. These structures largely determine the present morphology of the eastern Sierra de 10s Fdabres.

ACKNOWLEDGEMENTS

Electron microprobe analyses were performed at the electron microprobe laboratory of the Institute for Earth Sciences of the Free University, Amsterdam, with finanical and personnel support by NWO-WACOM (research group for analytical chemistry of minerals and rocks subsidized by the Dutch Organization for the Advancement of Pure Research). We thank Dr P. Maaskant, Dr C. Kieft and W. J. Lustenhouwer for performing the analyses.

We also thank the technical staffs of the Free University and the University of Amsterdam for the preparation of many hundreds of thin sections, Alwine Prinsen for typing the manuscript and Fred Kievits for drawing the figures.

One of us (HEB) benefited from a grant from Het Vakgroepfonds Structurele Geologie van de Universiteit van Amsterdam. Professor Dr B. R. Frost and Dr J. P. Platt have reviewed this paper thoroughly and quickly. Their reviews and subsequent discussions with Dr J. P. Platt helped to improve the text and clarified our way of thinking.

Last but not least we would like to thank Dr 0. J. Simon and K. Linthout for their criticism and many stimulating discussions on the geology of the Betic Cordilleras.

The order of the first two authors was determined by flipping a win.

REFERENCES

Alvarez, F. & Aldaya, F., 1985. Las unidades de la Zoaa Bttica en la region de Aguilas-Mazarr6n (Rov. de Murcia.) Esludiar Geologiws, 11,139-146.

Alvarez, W., Coamza, T. & Wezcl, F. C.. 1974. Fragmentation of the Alpine orogenic belt by miuoplate dispersal. Nafwe, 248, 309-314.

Akkennan, J. H., Maicr. G. & S i n , 0. J., 1980. On the geology of the Alpujarride Complex in the western Sierra de las Estaacias (Betic Cordilleras, SE Spain). Geologie en Mynbouw, 99, -374.

B d , P., 1%3. chloritoid und Paragonit am der Ophiolith- Zone von Zermatt-Saas Fee. Schweizerische Minemlogische und Perrogmphirche Miruilungen, 43,269-286.

Bellon, H., Coulon, C. & Edel, J. P., lw7. Le dtplaement de la Sardaigne: sync de la dondes gCochronologique, mag- matiques et palComagdtiques. Bulleiin & la Soeictc gCobgique

Boettcher, A. L. & Wyllie, P. J., 1967. Revision of the calcite-aragonite transition, with the location of a triple point between calcite I, calcite Il and aragonite. Nature, 2l3, 792-793.

Bohlen, S. R., Wall, V. J. & Boettcher. A. L.. 1983. Experimental investigations and plogical applicatiom of equilibrium in the system Fe0-li02-A1203-Si0,-H,0. Amrricon Minerdogist, 68,1049-1058.

tectonics of continental blocks in the western mediterrancan. Bulledn of the Geological Society of America, 9@, 1074-1083.

Brown, E. H., lw7. Tbe Crossite content of Ca-amphibole as a guide to pressure of metamorphism. J o d of Petrology, 18, 53-72.

Carman. J. H. & Gilbert, M. C.. 1983. Experimental studies on glaucophane stability. Amcricon J o d of Science, =A, 414-437.

Chatterjee, N. D., Johannes. W. & Lciitner, H., 1984. The system Ca0-&03-Si02-H20: new phase equilibria data, some calculated phase relations and their petrological applications. Conhibutions to Mineralogy and Petrology, 88,

Chopin, C. & Schreyer, W., 1983. Magnesio carpholite and magnesio chloritoid: two index minerals of pclitic blueschists and their preliminary phase relatiom in the model system MgO-&O,-SiO,-H,O. American J o d of Science, =A, 72-96.

Cipriani, C., Sassi, F. P. & Scolari. A.. 1971. Metamorphic white micas: definition of paragenetic ficlds. Schweizerirche Miner&- girchc und Perrogmphirchc M W u n g r n , 51,259-302.

Currie, K. L. & Curtis, L. W., 1976. An application of multiamponent solution theory to jadeitic pyroxenes. J o d of Geology, 84,179-194.

Davy, Ph. & m e t , Ph., 1986. The stacking of thrust slice in collision mnes and its thermal consequences. Tectonics, 5, 91S-929.

Diaz de Federico. A., Gomez-Pugnaice, M. T.. Puga, E. & Sassi, F. P., 1979. New problems in the Sierra Nevada Complex (Betic Cordilleras, Spain). N e w JcJvbuch f i r Geologie und

Draper, G . & Bone, R., 1980. Denudation rates, thermal evolution, .and preservation of blueschist terrains. Journal of Geology, 89,601613.

Egeler, C. G., 1964. On the tectonics of the eastern Betic Cordilleras. Geobgische Rundscluuc, 53,260-269.

Egeler, C. G. & Simon, 0. J . , 1969. Sur la tectonique de la Zone

Koninklijke Ne&rk?ndse Akudemie van Wetcnrdmppm. Afde- ing Naruurkundc, eente reeks 25, pt. 3,90 pp.

& Fmn~r, 119,825-831.

B0~1~0uilh. R. & Gonline, D. S., 1979. RC-TriaSsic fit and a l p k

1-13.

~ h ~ n t o i ~ g i c . M O ~ I U ~ ~ , 10,511-585.

Bttique (Cordillbres Bctiqucs, Espagne). Vrrtrandelu, . gen der

31) H. E. BAKKER ET AL.

Ellis, D. J. & Green, D. H., 1979. An experimental study of the effect of Ca upon garnet-clinopyroxene Fe-Mg exchange equilibria. Conhiburions to Mineralogy and Petrology, n, 13-22.

England, P. C. & Thompson, A. B., 1984. Pressure- temperature-time paths of regional metamorphism I. Heat transfer during the evolution of regions of thickened continental crust. Journal of Petrology, 25,894-928.

Ernst, W. G., 1979. Coexisting sodic and calcic amphiboles from high-pressure metamorphic belts and the stability of barroisitic amphibole. Mineralogical Magazine, 43,269-278.

Essene, W. J. & Fyfe, W. S., 1%7. Omphacites in Californian metamorphic rocks. Contributions to Mineralogy and Petrology, B, 125-140.

Evans, B. W., Johannes, W., Oterdoom, H. & Trommsdorff, V., 1976. Stability of Chrysotile and Antigorite in the serpentine multisystem. Schweizerische Mineralogische und Petrograph- irche Mitiedungen, 56,79-93.

Franz, G. & Spear, F. S., 1983. High pressure metamorphism of siliceous dolomites from the central Tauern Window, Austria. American Journal of Science, =A, 396-413.

Ganguly, J., 1972. Staurolite stability and related parageneses: theory, experiments and applications. Journal of Petrology, U, 335-365.

Garcia-Hernandez, M., Lopez-Gamdo, A. C., Rivas, P., Sanz de Galdeano, C. & Vera, J. A., 1980. Mesozoic paleogeographic evolution of the External Zone of the Betic Cordilleras. Geologie en Mijnbouw, 59, 155-168.

Gomez-Pugnaire, M. T. & Fernandez-Soler, J. M., 1987. High pressure rnetamoxphism in metabasites from the Betic Cordilleras (S.E. Spain) and its evolution during the Alpine orogeny. Contributions to Mineralogy and Petrology, 95, 231-244.

Graham, C. M. & Powell, P., 1984. A garnet-hornblende thermometer: calibration, testing and application to the Pelona Schists, S . California. Journal of Metamorphic Geology, 2,

Green, T, H. & Hellman, P. L., 1982. Fe-Mg partitioning between cozxisting garnet and phengite at high pressure and comments on a garnet-phengite geothermometer. Lithos, 15,

Guidotti, C. V. & Sassi, F. P., 1976. Muscovites as a petrogenetic indication mineral in pelitic schists. N e w Jahrbuch f i r Mineralogie Abhandlungen, 1zI, 97-142.

Helmers, H., 1982. Eclogitization of Hedenbergite skarns in the Sierra de 10s Filabres, SE Spain. Terra Cog&, 2-3,320.

Helmers, H., 1983. Alpine eclogite-blueschist-greenschist transi- tion in sheets flattened during ascent, Sierra de 10s Filabres, SE Spain. Terra Cognita, 3,185.

Helmers, H. & Voet, H. W., 1S67. Regional extension of the NevadwFilabride nappes in the eastern and central Sierra de 10s Fiiabres (Betic Cordilleras, SE Spain). Pruceedings of the Koninklijke Nederlnndse Akademie van Wetenschappen, Series

Hermes, J. J., 1978. The stratigraphy of the Subbetic and southern Prebetic of the Velez Rubio-Caravaca area and its bearing on the transcurrent faulting in the Betic Cordilleras in southern Spain. Proceedings of the Koninklijke Nederlandse Ahiademie van Wetenschappen, Series B, 87, 319-333.

Hirschberg, A. & Winkler, H. G. F., 1968. Stabilitatsbeziehungen zwischen Chlorit, Cordirit und Almandin bci der Metamor- phose. Contributions to Mineralogy and Petrology, 18,17-42.

Holdaway, M. J., 1971. Stability of andalusite and the aluminium-silicate diagram. American Journal of Science, 271,

Holland, J. T. B., 1980. The reaction albite= jadeite+quartz determined experimentally in the range of 600-12WC. American Mineralogirt, 65, 129-134.

Hoschek, G., 1969. The stability of staurolite and chloritoid and their significance in metamorphism of pclitic rocks. Contributions to Mineralogy and Petrology, 22,208-232.

Kampschuur, W., 1975. Data on thrusting and metamorphism in

13-32.

253-266.

B, 70, 239-253.

97-131.

the eastern Sierra de 10s Filabres: higher Nevado-Filabride units and the glaucophanitic greenschist facies. Tectonophysics, 27,

Kienast, J. R. & Triboulet, C., 19n. Le chloritoide dans le paragencsc glaucophane. albite ou paragonite. Societe Fruqaise de Mineralogie et Crisrcrllographie Bdetin. 95,565-573.

Kozur, H., Mulder-Blanken, C. W. H. & Simon, 0. J., 1985. On the Triassic of the Betic Cordilleras (S. Spain), with special emphasis on holothurian scleretis. Proceedings of the Konink- lijke Nederlandre Akademie van Wetenschappen, Series B , 88,83-110.

Krogh, E. J. & RBheim, A., 1978. Temperature and pressure dependence of Fe-Mg partitioning between garnet and phengite with particular reference to eclogites. Contributions to Mineralogy and Petrology, 66,75-80.

Kushiro, I., 1969. Clinopyroxene solid solutions formed by reaction between diopside and plagioclase at high pressure. In: prroxenes and Amphiboles: Crystal chemistry and phase petrology, Mineralogical society of America, Special Papers, 2,

Langenberg, C. W., 1972. Polyphase deformation in the eastern Sierra de 10s Fiabrcs north of Lubrin, SE Spain, GUA Papers of Geology, Series 1, 2.81 pp.

Linthout. K. & Kieft, C., 1970. Preliminary note on a mboziite of metamorphic origin. Mineralogical Magazine, 37, 629-631.

Linthout, K. & Westra, L., 1968. Compositional zoning in almandine-rich garnets and its relation to the metamorphic history of their host rocks. Proceedings of the Koninklijke

. Nederlandse Akademie van Wetenschappen Series B, 71,

Makel, G. H., 1985. The geology of the Malaguide Complex and its bearing on the geodynamic evolution of the Betic-Rif Orogen (southern Spain and northern Morocco). GUA Papers of Geology, Series 1, 22. 263 pp.

Mares&, W. V., 1977. Experimental studies on glaucophane: An analysis of present knowledge. Tectonophysics, 43, 109-125.

Martinez-Martinez, J. M., 1985. Las sucesiones Nevado-FilBbride en la Sierra de 10s Filabres y Sierra Nevada. Correlaciones. Cuadernos de Geologia, l2, 127-144.

Massonne, H . 4 . & Schreyer, W., 1983. A new experimental phengite barometer and its application to a variscan subduction zone at the southern margin of the Rhenohercynicum. Terru

Misch, P., 1969. Paracrystalline microboudinage of zoned grains and other criteria for synkinematic growth of metamorphic minerals. American Journal of Science, 267,43-63.

Montenat. Chr., Ott d'Estevou, Ph. & Massc, P., 1W. Tectonic-sedimentary characters of the Betic Neogene basins evolving in a crustal transcurrent shear zone (SE Spain). Bulletin Centrm Recherche Exploration-Production. Erf

Moretti. J. & Froideveaux, C.. 1986. Thermomechanical models of active rifting. Tectonics 5,501-51 1.

Mottana, A.. 1983. Orderdisorder in omphacites: practical determination and petrological significance. Terra Cognita, 3, 188.

Muir Wood, R., 1980. Compositional zoning in sodic amphiboles from the blueschist facies. Mineralogical Magazine, 43,

Newton, R. C. & Kennedy, G. D., 1963. Some equilibrium reactions in the join CaAI,Si,08-H20. Journal of Geophysical Research, 68, 2967-2983.

Newton, M. S. & Kennedy, G. C., 1968. Jadeite. analcite, nepheline and albite at high temperatures and pressures. American J o d of Science, 266, 728-735.

Nijhuis, H. J.. 1964. Plurifadal alpine metamorphism in the southcastern Sierra de 10s Fiabres south of Lubrin, SE Spain. Unpubl. PhD thesis University of Amsterdam, 151 pp.

Nitsch. K.-H., 1970. Experimentele Bestimmung der oberen Stabilitiitsgrenzc von Stilpnomelaan. Fortschritte der Mineralogie, 47,48-49.

57-81.

179-191.

297-312.

Cog&, 2-3, 187.

Aquirainc, 11, 1-22.

741-752.

CEODYNAMIC EVOLUTION OF THE BETICS 38l ~~

Niwh, K.-H., 1972. Das P-T-X.CO,-Stabilitatsfeld von Lawsonit. Conhibutions to Mineralogy and Petrology, 34,116-134.

PapiLC, J. J., Cameron, K. L. & Baldwin, K., 1974. Amphiboles and Pyroxenes: charederization of other than quadrilateral components and estimates of femc iron from microprobe data. Geologicul Society of America Abstracts with programs, 6 , 1053-1054.

Pearce, J. A. & Cann, J. R., 1973. Tectonic setting of basic volcanic rocks determined using trace elements analysis. Ehrrh and Planetary Science Lctters, 19,290-330.

Platt, J. P., 1986. Dynamics of orogenic wedges and the uplift of high-pressure metamorphic rocks. Buflefin of the Geologid

Platt, J. P., 1981. Tbe uplift of high-pressure low-temperature metamorphic rocks. Philosophical T m a c i i o m of the Royal Society of London, A321,87-103.

Platt, J. P. & Vissers, R. L. M., 1980. Extensional structures in anisotropic rocks. Journal of Structural Geology, 2, 397-410.

Priem, H. N. A., Boelrijk, N. A. I. M., Hebeda, E. H. & Verschure, R. H., 1966. Isotopic age determinations on tourmaline granite-gneisses and a metagranite in the eastern Betic Cordilleras (southeastern Sierra de l a Filabres), S.E. Spain. Geologie en Mijnbouw, 45,184-187.

Priem, H. N. A., Boelrijk, N. A. I. M., Hebeda, E. H., Oen, I. S., Verdurmen, E. Th. & Verschure, R. H., 1979. Isotopic dating of the emplacement of the ultramafic masses in the Serrania de Ronda, southern Spain. Confributions to Mineral- ogy and Petrology, 70,103-109.

Puga, E. & Diaz de Federico, A., 1978. Metamorfismo polifaSico y deformaciones alpinas en el Complejo de Sierra Nevada (Cordillera Bttica). rmplicaaones geodiniimicas. Proceedings Reunidn sobre la geodhmica de kI Cordillera B&rica y Mar de Alborbn, 79-111.

Rehault, J-P., Boillot, G. & Mauffret, A., 1984. The western Mediterranean basin, gwlogical evolution. Marine Geology, 55, 447477.

Raheim, A. & Green, D. H., 1974. Experimental determination of the temperature and pressure dependence of the Fe-Mg padtion coefficient for coexisting clinopyroxene and garnet. Contributions to Mineralogy and Petrology, 48,179-203.

Richardson, S . W., 1968. Staurolite stability in a part of the system Fe-Al-Si-0-H. Joumal of Petrology, 9,467-488.

Richardson, S . W. & England, P. C., 1979, Metamorphic consequenas of crustal eclogite production in overthrust orogenic zones. Earth and Planetary Science Letters, 42, 183-190.

Roep, Th. B. & MacGillavry, H. J., 1962. Preliminary note on the presence of distinct tectonic units in the Betic of Malaga of the Velez Rubio region (SE Spain). Geologic en Mijnbouw, 41, 423-429.

sociery of A h , 97,1037-1053.

Sassi, F. P. & Saolari, A., 1974. The b,-value of potassic white micas as a barometry indicator in low-grade metamorphism of pelitic schists. Contributions to Mineralogy and Petrology, 45, 143-152.

Simon, 0. J. & V i h e r , H., 1983. El PCnnico de las Cordilleras BCticas. In: Carbodfen, y Ptnnico de Espatia. X Congreso Interncrcional de Estratigrajb y Geologh Carbonffero. (ed. Martinez Diaz, C.). Madrid, 1983,453-499.

Spry, A., 1969. Metumorphic Texwes. Pergamon Press, Oxford,

Storre, B. & Nitsch, K.-H., 1972. Die Reaktion 2 Zoisit + lC02 = 3 Anorthit + 1 Calcit + 1H20. Conrribulionr to Mineral- ogy and Petrology, 35,l-10.

Thompson, A. B., 196. Mineral reactions in pelitic rocks: I Prediction of P-T-X (Fe-Mg) phase relations. American Journal of Science, 276,401-424.

Torres-Roldan, R. L., 1979. The tectonic subdivision of the Betic Zone (Betic Cordilleras, Southern Spain): its signilicance and one possible gwtectonic scenario for the westernmat alpine belt. American Journal of Science, 279, 19-51.

Tubia, J. M. & Cuevas, J., 1986. High temperature emplacement of the Los Reales peridotite nappe (Betic Cordilleras, Spain). Journal of Structural Geology, a, 473-482.

Vissers, R. L. M., 1981. A structural study of the central Sierra de la Filabres (Betic Zone, SE Spain), with emphasis on deformational processes and their relation to the alpine metamorphism. GUA Papers of Geology, Series 1, 15,154 pp.

Voet, H. W., 1%7. Geological investigations in the northern Sierra de 10s Fiiabres around Macael and Cobdar, southeastern Spain. Unpubl. PhD thesis University of Amsterdam, 122 pp,

Viilk, H. R., 1967. Zur Geologie und Stratigraphie des Neogenbeckens von Vera, Siidost Spanien. Unpubl. PhD thsis, University of Amsterdam. 160 pp.

Weijermars. R., Roep, Th. B., Van den Eeckhout, B., Postma, G. & Kleverlaan, K., 1985. Uplift history of a Betic fold nappe inferred from Neogene-Quaternary sedimentation and tectonics (in the Sierra Alhamilla and Almeria, Sorbas, and Tabernas Basins of the Betic Cordilleras, S E Spain). Geologic en Mijnbouw, 64,397-411.

Westerhof. A. B., 1975. Genesis of magnetite ore near MarbcUa, southern Spain: formation by oxidation of silicates in polymetamorphic gedrite-bearing and other rocks. GUA Papers of Geology, Series 1, 6,216 pp.

Zwart, H. J., 1962. On the determination of polymetamorphic mineral associations and its applications to the Bosost area (Central Pyrenees). Ceologische Rundschau, 53,38-65.

350 pp.

Received 15 February 1988; reviswn accepted 6 July 1988.