[The Journal of Geology, 1999, volume 107, p. 399–419] q 1999 by The University of Chicago. All rights reserved. 0022-1376/1999/10704-0002$01.00

399

Mafic Precursors, Peraluminous Granitoids, and Late Lamprophyresin the Avila Batholith: A Model for the Generation of

Variscan Batholiths in Iberia

Fernando Bea, Pilar Montero, and Jose F. Molina

Department of Mineralogy and Petrology, Fuentenueva Campus, University of Granada, 18002 Granada, Spain(e-mail: fbea@goliat.ugr.es)

A B S T R A C T

The Avila batholith of central Spain is composed of upper Carboniferous peraluminous granitoids that were precededby volumetrically insignificant bodies of mafic-ultramafic hybrid magmas and postdated by several dike swarms ofcamptonitic lamprophyres. Rb-Sr dating indicates continuous magmatic activity from 350 Ma to 280 Ma, starting∼ ∼with the mafic precursors and a few midcrustal anatectic leucogranites, followed by massive autochthonous andallochthonous granodiorites and granites, and ending with the camptonitic lamprophyres. Early hybrid mafic magmas( ; ) were produced in small batches during or immediately after the main deformation340 Ma 340 Ma« Sr ∼ 25 « Nd ∼ 21.5phase, probably by the partial melting of a mixture of 8 : 2 mantle and biotite-bearing crustal rocks at the crust-∼mantle interface. These magmas were emplaced in the middle crust at 340 Ma, advecting a negligible amount of∼heat. The generation of crustal granites during the main deformation phase was very scarce and limited to highlyfertile protoliths, rich in heat-producing elements, affected by strong shear zones. The generation of crustal granitoidson a batholithic scale took place from 330 Ma to 290 Ma, during the main extensional period. Granites∼ ∼( –150; ) were produced by the partial melting of fertile crustal rocks310 Ma 310 Ma« Sr ∼ 45 « Nd ∼ 22.1 to 2 9( –218; ), characterized by high heat production ( 2.5–3 mW m23). The zone of310 Ma 310 Ma« Sr ∼ 48 « Nd ∼ 22.2 to 2 9 ∼partial melting, 15–22 km in depth, was heated by thermal conduction from below after crustal thinning, but the∼contribution of radiogenic heat and the fertility of source rocks would have been essential for anatexis. The fastthinning of the crust from 310 Ma to 285 Ma released lithostatic pressure in the upper mantle and caused decom-∼ ∼pressional melting of the metasome layer at 60–85 km in depth, producing camptonitic melts dated at 283 Ma.∼ ∼The existence of a fertile metasome layer implies that the lithospheric mantle beneath central Iberia was not activelyinvolved in subduction during the Variscan orogeny.

Introduction

The Avila batholith, one of the largest and mostcomplicated plutonic complexes in central Iberia,is mainly composed of upper Carboniferous biotite5 cordierite 5 muscovite 5 aluminosilicate mon-zogranites, granodiorites, and leucogranites. Thesegranites were preceded by volumetrically insignif-icant bodies of mafic-ultramafic calc-alkalinerocks, hereafter called the mafic precursor, andpostdated by several dike swarms of camptoniticlamprophyres (Bea 1985; Bea and Corretge 1986;Fuster and Villaseca 1987). Field relationships in-dicate that the mafic precursors are only slightlyolder than the spatially related granites. Campton-itic lamprophyres, previously thought to be Triassic

Manuscript received October 28, 1998; accepted March 18,1999.

(Bea and Corretge 1986), owing to their analogieswith dikes in Portugal dated at 235–205 Ma (Fer-reira and Macedo 1979), are roughly contempora-neous or only slightly younger than the latest gran-ites. Such a close space-time association ofmantle-derived calc-alkaline mafic rocks, crust-de-rived strongly peraluminous granites, and mantle-derived alkaline rocks, repeatedly found in centralIberia, cannot be accidental; instead, it reflects asystematic interaction between the continentalcrust and upper mantle in this region during theVariscan orogeny.

To understand the genetic link among theserocks, we carried out a detailed field, petrographic,geochemical, and isotopic study in the Gredos sec-tor of the Avila batholith (fig. 1), where the mafic

Figure 1. Geological sketch of the Avila batholith. The size of the mafic precursor bodies is exaggerated. Numbers representthe location of samples for the isochrons shown in figure 7. See also table 10e (available from The Journal of Geology office uponrequest).

Journal of Geology V A R I S C A N B A T H O L I T H S I N I B E R I A 401

precursors and late camptonitic lamprophyres areparticularly abundant and well exposed. We deter-mined the Rb-Sr age of representative plutons ofeach rock type and estimated, from geobarometriccalculations, the loci of melt production. Scanningelectron microscopy (SEM) studies of accessorymineral assemblages allowed us to identify assim-ilated crustal components in the mafic precursors.Using major and trace elements, and Sr and Ndisotope geochemistry, we have outlined the pet-rogenesis of each rock type. On the basis of thesedata, we propose a model for the petrogenesis ofthe batholith and discuss its most important con-sequences for the evolution of this segment of theVariscan belt.

Samples and Methods

Two hundred fifty-three samples from the best ex-posures of representative plutons and facies in theGredos sector were collected for this study. Slidesfor thin sections were cut from each sample foroptical, SEM, and microprobe analyses. Five to 10kg of rock were crushed to a grain size of !5 mmwith a jaw crusher. Powders for chemical analysiswere obtained by grinding about 50 grams per sam-ple of crushed rock in a tungsten carbide jar. Thisprocedure produced no detectable contamination oftrace elements, except W and Co. All samples wereanalyzed for major elements and 40 trace elements,including 14 rare earth elements (REEs). A subsetof 83 samples was analyzed for Sr isotopes, and 63of them were also analyzed for Nd isotopes. Sam-ples for each isochron were collected as close aspossible spatially in order to minimize the heter-ogeneity of the initial isotope ratios. In the case oflamprophyres, each sample is from a different dike,but all were collected along a 100-m cross sectionwhere the dike swarm is densest. In total, we con-structed 10 Rb-Sr isochrons, the locations of whichare shown in figure 1.

Major element determinations were performedby X-ray fluorescence after fusion with lithium te-traborate. Typical precision was better than 51.5%for an analyte concentration of 10 wt %. Zirconwas determined by X-ray fluorescence on pressedpellets, with a precision better than 54% for 100ppm Zr. Trace element determinations were doneby ICP-mass spectrometry (ICP-MS) after

digestion of 0.1000 g of sample powderHNO 1 HF3

in a Teflon-lined vessel at 1807C and 200 psi for∼ ∼30 min, evaporation to dryness, and subsequent dis-solution in 100 mL of 4 vol % HNO3. Instrumentmeasurements were carried out in triplicate witha PE SCIEX ELAN-5000 spectrometer using rho-

dium as internal standard. Precision was betterthan 52% and 55% for analyte concentrations of50 and 5 ppm, respectively.

Samples for Sr and Nd isotope analyses were di-gested in the same way, using ultraclean reagents,and analyzed by thermal ionization mass spectrom-etry in a Finnigan Mat 262 spectrometer after chro-matographic separation with ion exchange resins.Normalization values were 88 86Sr/ Sr 5 8.375209and . Blanks were 0.6 and 0.09146 144Nd/ Nd 5 0.7219ng for Sr and Nd, respectively. The external pre-cision (2j), estimated by analyzing 10 replicates ofthe standard WS-E (Govindaraju et al. 1994), wasbetter than 0.003% for 87Sr/86Sr and 0.0015% for143Nd/144Nd. The 87Rb/86Sr and 147Sm/144Nd were di-rectly determined by ICP-MS, following themethod developed by Montero and Bea (1998), witha precision better than 1.2% and 0.9% (2j),respectively.

Major element analyses of minerals were ob-tained by wavelength dispersive analyses with aCamebax SX-50 electron microprobe, using syn-thetic standards. Accelerating voltage was 20 kV,and beam current was 15 nA. Coefficients of var-iation were close to 1%, 2.5%, and 5% for 10 wt%, 1 wt %, and 0.25 wt % analyte concentration,respectively. REE analyses of garnet were done witha laser ablation ICP-MS probe, following themethod described in Bea et al. (1996, 1997). Thecoefficient of variation in determining Gd/Dy ra-tios was better than 1%.

Analytical results are available electronically ei-ther from the The Journal of Geology or from theauthors; see a summary of available data intable 1.

Geological Setting

The Central Iberian Zone makes up the middle partof the Variscan belt of the Iberian Peninsula (Juliv-ert et al. 1972). It consists of Proterozoic and lowerPaleozoic metasediments and orthogneisses, in-truded by numerous syn- or late-kinematic Varis-can granitoids. The first phase of Variscan defor-mation at 360–350 Ma (e.g., Ferreira et al. 1987;∼Serrano Pinto et al. 1987) nearly doubled the thick-ness of the crust. Crustal thickening seems to havebeen caused mainly by the underthrusting of theOssa Morena lower crust beneath central Iberiaduring the lower Carboniferous, as recently docu-mented by Azor et al. (1994), Aller (1996), and Ex-posito Ramos et al. (1998). The main compressivephase was followed by a period of extensional tec-tonics, characterized by the development of sub-horizontal shear zones (Escuder Viruete et al. 1994),

402 F . B E A E T A L .

Table 1. Summary of Data Tables Available Electronically upon Request, Either from The Journal of Geology Officeor from the Authors

Table number Contents

1e Electron microprobe composition of representative minerals of all facies2e Laser ablation ICP-MS REE analyses (ppm) and pressure estimations with the Gd/Dy garnet

geobarometer (Bea et al. 1997) of garnets from three granulitic xenoliths scavenged bycamptonites

3e Whole-rock composition of selected samples of Pena Negra migmatites and orthogneisses4e Whole-rock composition of selected samples of the mafic precursors5e Whole-rock composition of selected samples of camptonites and related plagiophyres6e Whole-rock composition of selected samples of HO granodiorites and monzogranites7e Whole-rock composition of selected AL granodiorites and monzogranites8e Whole-rock composition of selected samples of PL granites9e Whole-rock composition of selected FA leucogranites10e Sr and Nd isotope composition of selected samples of the Avila batholith

Note. These tables are numbered in the text with a number followed by the character “e” (electronic). earth element;REE 5 rare; ; ; del Alberche.HO 5 Hoyos AL 5 Alberche PL 5 Plasencia FA 5 Fuente

during which most of the granites were produced(Bea 1985). The last Variscan tectonic movementswere compressional and produced subvertical shearzones that affected the youngest granite bodies (Lo-pez Plaza and Martınez Catalan 1987).

The Avila batholith (fig. 1) crops out over an areaof 13,000 km2 within the axial part of the Central∼Iberian Zone. It consists of two sectors: Gredos tothe west (Bea 1985) and Guadarrama to the east(Fuster and Villaseca 1987). The Gredos sector iscomposed of coalesced plutons of strongly peralu-minous granites, together with large low-pressureanatectic complexes preferentially located alongthe axial zone of the batholith (Bea and Pereira1990; Pereira 1993; fig. 1). Numerous small bodiesof cortlandite, gabbro, diorite, and tonalite appeareither as large enclaves inside the peraluminousgranitoids or as intrusions into migmatites. Thesebodies are more abundant near the northern bound-ary of the batholith, where mafic plutons arealigned in several lineaments parallel to the Var-iscan foliation (fig. 1). Camptonitic lamprophyresare concentrated in the middle part of the Gredossector, within a roughly rectangular region of about10 km E-W and 30 km N-S (fig. 1), where 1100 smallN-S or NNE-SSW lamprophyric dikes occur.

Field Relations and Petrography

Mafic precursor bodies range from a few meters toa few hundreds of meters in size. The largest aretypically roughly zoned, with cortlandites or horn-blendites in the core and tonalites or granodioritesin the rim. Contacts with host granites are com-plex, and some mingling is locally evident (Mor-enoventas et al. 1995), producing breccioid mixedfacies with an overall granodioritic compositioncharacterized by the simultaneous presence of K-

feldspar megacrysts with large, locally hollow, crys-tals of zoned amphibole and rare orthopyroxene.The mixing zone is usually no more than a fewmeters in width. Mafic-ultramafic bodies also in-truded high-grade metapelites, causing an aureoleof intense migmatization (Franco Gonzalez andGarcıa de Figuerola 1986).

Petrographically, mafic precursors consist ofphlogopite and/or amphibole-bearing peridotites,olivine-amphibole-orthopyroxene 5 clinopyroxenegabbros, and amphibole-biotite diorites, quartz di-orites, and granodiorites. They commonly have anisotropic fabric, but in some cases the outer bordersof plutons may also have a deformational foliation.Textures are granular, ophitic, or subophitic, lo-cally cummulatic, especially in ultramafic mem-bers (Franco Gonzalez and Garcıa de Figuerola1986; Franco Gonzalez and Sanchez Garcıa 1987).Amphibole and biotite can be slightly porphyritic.Coronitic textures around olivine and titanian am-phibole are common in gabbros and cortlandites.Olivine (Fo80–70) is usually mantled by orthopyrox-ene (En75–70). Titanian pargasitic hornblende isrimmed by an Fe-Mg amphibole, probably cum-mingtonite (table 1e). Mantled amphiboles coexistwith unzoned crystals of actinolite. Augite is com-mon in some gabbros, where it also appears rimmedby actinolitic amphibole. Diorites and quartz dio-rites rarely have mafic silicates other than biotiteand actinolitic hornblende. Both minerals usuallyform clusters that pseudomorphize original horn-blende or clinopyroxene.

The most abundant rock types in Gredos are, byfar, peraluminous granites and granodiorites. Bea(1985) identified four main types or “superfacies”according to field relationships and composition.The HO superfacies (named after the Hoyos gran-odiorite) accounts for 35% of the exposed surface;∼

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it is composed of the oldest granodiorites and mon-zogranites, which form thick subhorizontal plu-tons, showing gradational contacts to anatecticcomplexes. Rocks consist of a granodioritic, some-times tonalitic, medium-grained groundmass inwhich abundant, usually oriented, K-feldspar me-gacrysts stand out. The groundmass containsquartz 1 plagioclase (cores An40–35, rims An25–20) 1aluminous biotite 5 K-feldspar 5 cordierite (seecomposition in table 1e). Microgranular enclaves,biotite schlierens, and K-feldspar megacrysts usu-ally define a planar subhorizontal foliation subpar-allel to pluton contacts.

The AL superfacies (after the Alberche grano-diorite) accounts for 40% of the exposed surface;∼it consists of biotite 5 cordierite monzogranitesand granodiorites that form allochthonous plutonswith subvertical contacts and no internal fabric orpoorly defined subvertical foliation; they may con-tain enclaves of HO granodiorites. Under the mi-croscope, AL rocks differ from HO rocks mainly inthe conspicuous oscillatory zoning and the widercompositional range of plagioclase (cores An52–40,rims An25–15). Both HO and AL rocks contain abun-dant dark microgranular enclaves that can begrouped into three types: (1) biotite-sillimanite en-claves, nearly identical to the melanosomes of re-lated migmatites; (2) biotite-quartz-plagioclase(An55–25) granoblastic enclaves of an uncertain ori-gin; and (3) enclaves of mafic rocks, identical intexture and mineralogy to mafic precursors.

The PL superfacies (after the Plasencia granite)accounts for 20% of the exposed surface; it is com-∼posed of two-mica monzogranites and leucograni-tes that form allochthonous plutons of variousshapes that are commonly intrusive in HO and ALgranodiorites. PL granites are either equigranularor porphyritic, with phenocrysts of K-feldspar in agroundmass of quartz, K-feldspar, plagioclase (coresAn30–20, rims An10–5), aluminous biotite, and mus-covite. The FA superfacies (named after the Fuentedel Alberche leucogranite) accounts for 5% of the∼exposed surface; it contains leucogranites that mayhave any combination of biotite, muscovite, cor-dierite, almandine garnet, sillimanite, and anda-lusite. Plagioclase has a composition in the rangeAn12–0 and is usually not zoned. FA leucogranitesform small bodies that are either subautochthon-ous, closely connected to migmatites, or allo-chthonous, intruding into HO and, less commonly,AL granodiorites. PL and FA granites have no en-claves of mafic precursors.

Gredos granitoids are locally strongly deformedby both extensional and compressional structures.Gently dipping shear zones, corresponding to the

main extensional phase, affect the HO granitoidsand the oldest FA anatectic leucogranites, espe-cially those appearing as small satellites of themain batholith. Late subvertical shear zones affectall the granites (e.g., Bea and Morenoventas 1985;Bea et al. 1994).

In the central part of the Gredos sector, the hugePena Negra Anatectic Complex crops out (Bea andPereira 1990; Bea 1991; Pereira 1992; see fig. 1). Ithas abundant nebulitic or schlieren migmatites,minor metatexitic orthogneisses, and several au-tochthonous bodies of anatectic granitoids. Theleucosomes of diatexitic migmatites have hypidi-omorphic textures and contain quartz 1 plagioclase(cores An33–28, rims An28–12) 1 cordierite 1 biotite 1K-feldspar, with minor sillimanite and rare tour-maline. Melanosomes consist of alternating gra-noblastic (cordierite 5 quartz 5 plagioclase 5 K-feldspar) and schistose (sillimanite 1 biotite 1 il-menite 5 cordierite) layers. In a few exceptionalcases, migmatites bear large idiomorphic or sub-idiomorphic almandine-rich garnet crystals, gen-erally mantled by cordierite plus biotite. Orthog-neisses, consisting of quartz, plagioclase (coresAn48–32, rims An35–23), biotite, and K-feldspar, werestrongly affected by all the major phases of Variscandeformation and have been migmatized partially,developing fine-grained leucosomes with almostthe same mineralogical and chemical compositionas the mesosomes (Bea 1989). Small irregular bodiesof cordierite and leucogranite and large, subhori-zontal, concordant sheets of porphyritic, cordierite-bearing granodiorite and monzogranite, nearlyidentical to those of the HO superfacies, grade tran-sitionally into migmatites (Pereira and Bea 1994).On the basis of field, petrographic, geochemical,and isotopic evidence, we have assumed that ma-terials similar to the Pena Negra migmatites likelyrepresent the source rock for nearly all Gredos gran-ites (Bea and Morenoventas 1985; Bea 1985, 1991).

The lamprophyric dike swarm of central Gredosis made up of more than 100 N-S or NNE-SSWsubvertical dikes with a thickness of 5 cm to 2 mand little longitudinal continuity. They are com-monly emplaced in HO and AL granodiorites.About 90% of the dikes are composed of campton-ite, usually ocellar, while the rest consists of dif-ferentiated albite-bearing reddish porphyry, thecomposition of which is nearly identical to that ofthe ocelli (Bea and Corretge 1986). Camptoniteshave large (up to a few cm) idiomorphic crystals ofkaersutite and Ti-augite, rarely calcite and Ti-phlogopite, within a groundmass of minute idio-morphic crystals of kaersutite, Ti-augite, plagio-clase (An50–10), K-feldspar, Ti-magnetite, ilmenite,

404 F . B E A E T A L .

and abundant pyrite, with minor chalcopyrite andsphalerite, epidote, titanite, and calcite. The ocellirange from droplike to veinlike, are from 5 to 30mm in size, and have a composition of albite 1 K-feldspar 1 epidote 1 calcite 1 sulfides, presumablyoriginating from the segregation of an immisciblefelsic liquid. A few camptonitic dikes do containsmall rounded xenoliths of garnet-sillimanitegranulite.

Accessory Mineral Assemblages

SEM studies have revealed that the accessory min-eral assemblage of mafic precursors is amazinglyexotic. Cortlandites and gabbros, in addition tophases characteristic of mafic rocks (Fe-Ni-,Fe-Cu-, Pb-, Zn-, and Mo-sulfides, Cr-spinel, mag-netite, chromian magnetite, ilmenite, baddeleyite,zirconolite, barite, apatite, and zircon), contain nu-merous crystals of Th-rich monazite, thorite, ura-ninite, xenotime, wolframite, scheelite, and arse-nopyrite, phases commonly limited toperaluminous granites and metapelites (e.g., Bea1996). Remarkably, the exotic accessories appearincluded within minerals that seem to be earlycrystallizing phases, such as olivine, clinopyrox-ene, orthopyroxene, and Ti-amphibole (fig. 2). Zir-con crystals, especially the largest ones, are xeno-morphic, with skeletal or embayed contours,locally partially transformed to baddeleyite, com-monly containing inclusions of U-rich thorite, ura-ninite, and xenotime (fig. 3). Huge poikilitic crys-tals of apatite, either automorphic or xenomorphic,containing a plethora of minute inclusions of mon-azite, cheralite, and xenotime, are also common(fig. 4).

Diorites and quartz diorites, however, have veryrare monazite, thorite, uraninite, xenotime, W-minerals, and arsenopyrite. They do, however, con-tain huge crystals of other REE minerals neverfound in cortlandites and gabbros, such as a REE-Ca borosilicate (melanocerite?), allanite, and a lightrare earth element (LREE)—Ca arseniate (gaspar-ite?). Melanocerite may contain minute idiomor-phic inclusions of amphibole, biotite, and plagio-clase, which suggests it could have precipitateddirectly from the melt (fig. 5). Allanite and gaspar-ite, in contrast, seem to have formed by a reactionof the magma with early crystals of melanocerite(fig. 5) and arsenopyrite (fig. 6), respectively. Otheraccessories include zircon of various morphologies,apatite (never with the textures displayed in gab-bros and cortlandites), ilmenite (usually mantled bytitanite), rare magnetite, and a variety of Ni-Fe andFe-Cu sulfides.

Camptonites contain a great abundance of ac-cessory sulfides (pyrite, chalcopyrite, galena, sphal-erite, and rare petlandite) and oxides (Ti-magnetite,magnetite, and ilmenite). Titanite is common butalways appears partially pseudomorphized by mag-netite. Zircon and baddeleyite are very rare. Despitethe elevated whole-rock REE contents (table 5e),there are only minute amounts of a mineral withappreciable REE contents, identified under the SEMas a Zr-Ti-Ca phase, probably zirconolite.

Peraluminous granites, migmatites, and orthog-neisses contain nearly the same accessory assem-blage, usually consisting of ilmenite, pyrite, zircon,apatite, monazite, rare thorite and xenotime, andsporadic uraninite, chalcopyrite, arsenopyrite,wolframite, and scheelite. It is noteworthy that thisis the same association that appears as inclusionsin early phenocrysts of gabbros and cortlandites.Granites may also contain secondary titanite andallanite.

Rb-Sr Ages

The occurrence of REE-saturated accessories in allthe rock types, as well as the presence of likelyxenogenous zircon in mafic precursors and the vir-tual absence of this mineral in camptonites, led usto choose the Rb-Sr method for age determinations.We constructed 10 isochrons (see locations in fig.1), one for the mafic precursors, one for the camp-tonites, two for the HO granodiorites, two for theAL granodiorites, two for the PL granites, and twofor the FA granites, one strongly deformed and theother nearly isotropic (fig. 7; raw data in table 10e).

Mafic precursors give an age of Ma,340 5 18nearly the same as a strongly deformed FA leuco-granite, which yields Ma. HO granodiorites344 5 8range from Ma to Ma. AL gran-327 5 8 317 5 13odiorites range from Ma to Ma.310 5 9 306 5 8An undeformed FA leucogranite pluton intrusivein HO granodiorites yields Ma. PL gran-305 5 16ites yield to Ma. The campton-297 5 26 295 5 13itic dike swarm yields Ma. These data283 5 30confirm the picture inferred by field geology ofnearly continuous magmatic activity from 350∼Ma to 280 Ma, starting with the mafic precursors∼and a few leucogranites, followed by voluminousautochthonous and allochthonous granodioritesand granites, and ending with the camptoniticlamprophyres.

Geobarometric Estimations

The PT conditions of anatexis recorded in the PenaNegra Complex are 4–4.5 kbar and 7507–8007C, cor-

Journal of Geology V A R I S C A N B A T H O L I T H S I N I B E R I A 405

Figure 2. Inclusions of exotic accessories, probably xenocrysts derived from crustal materials, in primocrysts of amagnesian gabbro. Top, Uraninite (white) in Ti-amphibole (light gray). Bottom, Monazite (light gray) in clinopyroxene(light gray). BSE microphotographs.

responding to the formation of autochthonous bod-ies of cordierite-bearing granodiorites and monzo-granites dated at Ma (Pereira et al. 1992).310 5 6In other metamorphic areas of the Avila batholith,where the anatexis was not so intense, there arerelics of former Barrovian metamorphism (Franco

Gonzalez and Sanchez Garcıa 1987). Pressure es-timations with the GASP geobarometer and withthe garnet-only Gd/Dy geobarometer (Bea et al.1997) on xenoliths of felsic granulites dragged upby camptonites (likely representing the crustal unitunderlying the Avila batholith) give pressure esti-

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Figure 3. BSE microphotograph of a xenomorphic zircon (white gray) with an inclusion of thorite (white) in a gabbro.The zircon is partially transformed to baddeleyite, likely as a consequence of reaction with the Si-poor melt.

mates in the range of 5.5–7 kbar (table 2e). As dis-cussed below, these estimations indicate that thesource for peraluminous granites was a layer

15–22 km in depth.∼In the case of camptonites, pressure estimates

with the clinopyroxene-melt barometer (Putirka etal. 1996) are in the range of 13.5–17 kbar. In thecase of mafic precursors, pressure estimates withthe same barometer are near 3–5 kbar for all bodiescurrently found inside granitoids, but near 7.5–11kbar for bodies emplaced outside the anatexis do-main, in spite of the fact that they developed a low-pressure contact aureole. It seems, therefore, thatwhen mafic precursors intruded in relatively coolterrains that later did not undergo anatexis, someprimocrysts had insufficient time to reequilibrateat lower pressures and still record the pressure ofintratelluric crystallization. This pressure maytherefore be used as a rough estimate for the min-imum depth of melting, which, according to theabove-mentioned data, would have been close to 40km for mafic precursors and near 60 km forcamptonites.

Chemical and Isotopic Composition

Pena Negra migmatites have a composition similarto that of a peraluminous granodiorite, with SiO2

commonly in the range of 64–68 wt %, CaO ∼

wt %, wt %, wt0.8–1.5 Na O ∼ 2.0–2.7 K O ∼ 3–42 2

%, and the aluminum saturation index (ASI 5) ∼1.2–1.7 (tablemol Al O /[CaO 1 Na O 1 K O]2 3 2 2

3e). Their primitive mantle-normalized trace ele-ment patterns are characterized by the enrichmentin incompatibles with positive anomalies of theheat-producing elements (HPE: K, U, and Th) andPb and with negative anomalies of Sr, Eu, Nb, andTi (fig. 8). Chondrite-normalized REE patterns (fig.11) have ∼80–150, decrease with increasing atomicnumber, have a moderate negative Eu anomaly( ), and are nearly flat from Er to Lu,∗Eu /Eu ∼ 0.4–0.8with . Orthogneisses have the compo-Lu ∼ 5–10N

sition of a slightly peraluminous granodiorite, with, wt %,ASI ∼ 1.05–1.1 CaO ∼ 1–3 Na O ∼ 3.5–4.22

wt %, and wt % (table 3e). Their traceK O ∼ 3–42

element contents are similar to those of migma-tites, although with somewhat higher Ba, Zr, andLREE and lower Li, Rb, Cs, Zn, and V. The 87Sr/86Srof metapelites calculated at the time of anatexispeak, at ∼310 Ma, are in the range of 0.7113–0.7195,with Sr in the range of 102–218, and310 Ma 310 Ma« «Nd between 25 and 28 (fig. 11; table 9e). In thecase of orthogneisses, these parameters are 87Sr/86Srat 310 Ma ∼0.7075–0.7091, , and310 Ma« Sr ∼ 48–72

Nd between 22.2 and 23.2 (fig. 12).310 Ma«Mafic precursors define a calc-alkaline associa-

tion with SiO2 in the range of 45–65 wt %, MgOin the range 25–30 wt %, , and lowNa O ≥ K O2 2

Journal of Geology V A R I S C A N B A T H O L I T H S I N I B E R I A 407

Figure 4. Exotic REE minerals in cortlandites and gabbros. BSE microphotograph of apatite crystals (light gray) witha plethora of minute inclusions of monazite, xenotime, thorite, and uraninite (white) in a cortlandite. Crystals ofapatite may be either xenomorphic or idiomorphic, suggesting a complex history of dissolution precipitation (lightgray).

TiO2 (table 4e; fig. 9). Their primitive mantle-nor-malized trace element patterns (fig. 8) are enrichedin incompatible trace elements and have strongpositive anomalies of Th, U, and Pb and negativeanomalies of Nb, Ti, and Eu. They are, therefore,remarkably parallel to the average composition ofPena Negra migmatites (see table 3e), except in theabsence of the negative Sr anomaly (fig. 8). The Nd/Th ratio clusters around 3.5 (fig. 10), a value ab-normally low for mantle-derived rocks but nearlyidentical to that of monazite-bearing metapelites(e.g., Bea and Montero 1999). Chondrite-normalizedREE patterns (fig. 11) have a negative slope( ) and, frequently, a small negativeLa /Lu ∼ 6–10N N

Eu anomaly ( ). Initial 87Sr/86Sr iso-∗Eu /Eu ∼ 0.6–0.9topes are close to 0.7058 and have Sr near310 Ma«∼25 and Nd between 21.0 and 21.7 (fig. 12).310 Ma«

Camptonitic lamprophyres form a Ti-rich maficalkaline association, with (table 5e;K O 1 Na O2 2

fig. 9). Camptonites have wt %, withSiO ∼ 442

wt % and wt % (table 5e). TheirTiO ∼ 4 MgO ∼ 82

primitive mantle-normalized trace element pat-terns (fig. 8) reveal a strong enrichment in the mostincompatible elements, although, in contrast to themafic precursors, they have negative Th, U, and Pbanomalies but moderate positive Sr, Zr, and Ti

anomalies. Chondrite-normalized REE patterns(fig. 11) show a strong enrichment in LREE, with

and and a small but con-La ∼ 150–300 Lu ∼ 8–16N N

stant positive Eu anomaly ( ).∗Eu /Eu ∼ 1.04–1.24The isotopic composition (fig. 12; table 10e) is moreprimitive than in the case of the mafic precursors,with initial 87Sr/86Sr close to 0.7049, 285 Ma« Sr ∼

, and Nd between 20.1 and 21.0.285 Ma10–12 «HO granodiorites and monzogranites are mod-

erately silicic and strongly peraluminous, with SiO2

commonly in the range of 65–69 wt %, CaO ∼ 1–wt %, wt %, wt %, and3 Na O ∼ 2.5–4 K O ∼ 3–52 2

(table 6e). Their primitive mantle-ASI ∼ 1.1–1.4normalized trace element patterns (fig. 8) are nearlyidentical to those of Pena Negra migmatites, exceptfor the higher Rb, Th, and U contents. Their chon-drite-normalized REE patterns (fig. 11) are also verysimilar to those of Pena Negra rocks, but for theuniformity of the Eu anomaly ( )∗Eu /Eu ∼ 0.5–0.6and slightly higher contents of the heaviest heavyrare earth elements. The 87Sr/86Sr calculated at 310Ma is in the range of 0.7084–0.7123, with 310 Ma«

and Nd between 23.2 and 25.8,310 MaSr ∼ 60–116 «values halfway between those of Pena Negra dia-texitic migmatites and orthogneisses (fig. 12).

AL granodiorites and monzogranites are slightly

408 F . B E A E T A L .

Figure 5. Exotic REE minerals in diorites and quartz diorites. BSE microphotograph of an As-LREE mineral (whiteidiomorphic rim), probably gasparite, developed over a crystal of arsenopyrite (white-gray allotriomorphic core). Themiddle zone is composed of REE-Fe sulfides and oxides, probably hydrated. Diorites and quartz diorites do not containinherited monazite or thorite, having instead new uncommon REE minerals generated by reaction with the melt ofprevious REE phases or directly precipitated from the melt.

more silicic and less peraluminous than their HOequivalents. SiO2 is in the range of 66–73 wt %,

wt %, wt %,CaO ∼ 0.8–1.5 Na O ∼ 2.8–3.52

wt %, and (table 7e).K O ∼ 3.6–5.2 ASI ∼ 1.10–1.52

Their trace element composition bears many sim-ilarities with that of HO granitoids. Compared withPena Negra migmatites, they are enriched in U andTh but depleted in the most compatible trace ele-ments (fig. 8). Their chondrite-normalized REE pat-terns (fig. 11) are practically indistinguishable fromthose of HO granitoids. The isotopic composition(fig. 12; table 10e) is notably uniform: 87Sr/86Sr cal-culated at 310 Ma is in the range of 0.7073–0.7083,with and Nd between 22.1310 Ma 310 Ma« Sr ∼ 45–60 «and 24.1. These values are slightly, but signifi-cantly, more primitive than those of HO rocks butstill inside the range of Pena Negra orthogneissesand migmatites.

PL granites are more silicic than HO granitoidsbut equally peraluminous. SiO2 is in the range of69–70 wt %, wt %, wt %,CaO ∼ 2–3 Na O ∼ 3–42

wt %, and (table 8e).K O ∼ 3.5–5 ASI ∼ 1.05–1.152

Their trace element composition is also similar tothat of Pena Negra migmatites, except for the no-torious enrichment in Th and U (fig. 8) and the

more intense negative Eu anomaly ( ∗Eu /Eu ∼; fig. 9). The 87Sr/86Sr calculated at 310 Ma0.35–0.55

is in the of range 0.7093–0.7111, with 310 Ma« Sr ∼and Nd between 23.8 and 26.7 (fig.310 Ma70–60 «

12).Leucogranites have the greatest variations in age

and composition. Initially, we grouped them in twocategories. The first is composed of silicic graniteswith SiO2 in the range of 71–74 wt %, K O 12

, , , and trace ele-Na O CaO ∼ 0.6–1 ASI ∼ 1.2–1.42

ment contents similar to the most leucocratic mig-matites (e.g., table 9e, analyses 1 and 2), which usu-ally appear as independent intrusive bodies eitherin migmatites or in HO granodiorites. The othergroup has SiO2 in the range of 72–76 wt %;

; very low Ti, Fe, and Mg; an extremeNa O 1 K O2 2

depletion in most trace elements (e.g., table 9e,analyses 9 and 10); and nearly flat REE patterns atchondritic levels, with a precipitous negative Euanomaly and a secondary negative Nd anomaly (fig.11). According to their field relationships, theyprobably consist of leucocratic differentiates fromthe other granitoids. Leucogranites from the firstgroup were the first recorded magmatic products,and their production continued as long as crustal

Journal of Geology V A R I S C A N B A T H O L I T H S I N I B E R I A 409

Figure 6. BSE microphotographs of exotic REE minerals in diorites and quartz diorites. Top, Aggregate of melanoceritecrystals (white) partially transformed to allanite (gray). The melanocerite crystal contains minute idiomorphic in-clusions of amphibole, biotite, and plagioclase (see detail in bottom panel), indicating that it precipitated from themelt.

anatexis took place. The oldest among them havevery high initial 87Sr/86Sr, in the range of0.7180–0.7225, and ; the young-344 Ma« Sr ∼ 190–300est have initial 87Sr/86Sr in the range of0.7125–0.7147, with and310 Ma« Sr ∼ 120–150

Nd between 24.9 and 29 (fig. 12).310 Ma«

Discussion: Petrogenetic Outline andGeodynamic Implications

Mafic Precursors. The inclusion of accessoryphases characteristic of metapelites and anatecticgranites in early-crystallized major minerals, the

Figure 7. Rb-Sr isochrons of selected plutons; the numbers indicate the location in figure 1

Journal of Geology V A R I S C A N B A T H O L I T H S I N I B E R I A 411

Figure 8. Trace element composition of Gredos rocks normalized to the composition of the primitive mantle (takenfrom Hofmann 1988). Solid lines and gray bands represent the average and range of variation, respectively, for eachrock group. The dashed line indicates the average composition of the Pena Negra crustal protolith (see table 3e). Notehow the mafic precursors have the same trace element anomalies as crustal materials. See “Chemical and IsotopicComposition.”

existence of Fe-Mg amphiboles mantling titanianpargasitic hornblende, the trace element patternswith “crustlike” anomalies and interelement ratios(figs. 8 and 10), and the Sr and Nd isotope com-position (table 10e; fig. 12) all indicate that maficprecursors likely represent mantle magmas hybrid-ized with crustal materials. The concentric zoningof the largest bodies of mafic precursors, from cor-

tlandites and gabbros in the core to diorites andquartz diorites in the rim, would at first seem tosuggest that hybridization of the mafic magma withcrustal materials occurred at the level of emplace-ment, that is, at 15–18 km, according to geobaro-metric estimations. Other evidence, however, in-dicates that hybridization occurred at a deeperlevel. First, in Sr and Nd isotope evolution dia-

412 F . B E A E T A L .

Figure 9. Major element relations of Gredos rocks. Dots, mafic precursors; crosses, camptonitic lamprophyres. Thecontoured field represents peraluminous granitoids. See “Chemical and Isotopic Composition.”

grams, mafic precursors fall on lines that representclosed-system fractionation of an already homo-geneous magma, instead of mixing lines with hostanatexites (fig. 13). Second, the accessory phases(monazite, xenotime, thorite, xenocrystic apatiteand zircon, arsenopyrite, wolframite, etc.) that re-veal assimilation of crustal rocks are abundant asinclusions “quenched” in early primocrysts of theinternal cortlandites and gabbros, equilibrated at adepth greater than the emplacement level but sel-dom found in the external diorites and quartz di-orites, where they were probably dissolved or trans-formed by reaction with the melt. Since this

situation is exactly opposite to what one wouldexpect from in situ mixing, we suggest that thecurrent petrographic variety of zoned bodies is dueto the magmatic differentiation of an already hy-bridized magma. Some hybridization may also haveoccurred locally at the intrusion level by minglingwith host anatexites (Morenoventas et al. 1995),but it does not seem to have played a significantrole in producing the petrographic variety and pe-culiar chemical and isotopic composition of themafic precursors.

The hybrid mafic magma may have been gen-erated either (1) by partial melting of a mixture of

Journal of Geology V A R I S C A N B A T H O L I T H S I N I B E R I A 413

Figure 10. Th versus Nd plot of Gredos rocks. Dots,mafic precursors; crosses, camptonitic lamprophyres.The contoured field represents migmatites and orthog-neisses from Pena Negra. Values of Nd/Th close to 2.6are characteristic of monazite-bearing crustal materials(e.g., Bea and Montero 1999). The lowest values of Th/Nd in mantle-derived rocks occur in alkaline rocks, suchas those of Kola, which have (F. Bea, P. Mon-Th/Nd ≤ 6tero, and J. Molina, unpub. data). The anomalously lowvalues recorded in mafic precursors are consistent withtheir hypothesized origin as hybrid rocks.

peridotites plus biotite-bearing crustal materialsjust below the crust-mantle interface or (2) by thehybridization in the lower crust of an uncontami-nated mantle magma from an unknown depth.Given that (i) the amount of magma produced wasvery small, as indicated by available geophysicaland geological evidence; (ii) the melting event waslimited to the uppermost mantle, because it did notinvolve the metasome layer that was later thesource of alkaline magmas (see “Camptonites”);(iii) the hydrous component of biotite may havecaused local partial melting of mantle rocks at arelatively low temperature (Hirose and Kawamoto1995; Hirose 1997); and (iv) all mafic precursors,even the most magnesian, have a noticeable crustalcomponent (a situation difficult to envisage if hy-bridization was something accidental that hap-pened to magmas during their emplacement), welean toward the idea of melting of a hybrid mantle-crust source. Simple calculations with the expres-sion derived by Winther (1995) for estimating the

composition of partial melts reveal that melts witha major element composition close to the averagemafic precursors could be generated at12007–12507C and 10–14 kbar from an 8 : 2 mixtureof fertile lherzolite and metapelite. This proportionmay also explain the Sr and Nd isotope character-istics of mafic precursors, if we assume an isotopiccomposition for each component similar to the av-erage Pena Negra materials and the subcontinentalmantle of Europe during the Variscan orogeny, re-spectively (e.g., Voshage et al. 1990; Becker 1996;Stille and Schaltegger 1996; Lu et al. 1997, etc.).

The conditions for a suitable hybrid mantle-crustsource could have been achieved in two mutuallycompatible scenarios: a tectonic melange zone atthe base of the subducted crustal slab or in smallimbricated wedges of lower crust and upper mantle,caused by compressional low-angle faults. In eitherof these situations, the subsequent increase in heat,due to either frictional heating or the shifting ofisotherms to a new equilibrium position, couldhave caused first the breakdown of biotite and thenthe wet melting of the lherzolite. This mechanismseems capable of producing unconnected smallpods of volatile-rich mafic magmas that, if trans-ported to the middle crust along fault zones, wouldhave caused the characteristic lineaments of plu-tons parallel to the main Variscan structures. Itshould be emphasized that, given their small vol-ume, the amount of heat these melts advected fromthe mantle could not have contributed significantlyto the copious anatexis that occurred after theiremplacement.

Peraluminous Granitoids. Field relationships,major and trace element geochemistry, and isotopegeology indicate that the central Iberian granitesare nearly pure crustal magmas. On the «t Nd ver-sus «t Sr diagram (fig. 12), they plot between PenaNegra orthogneisses and metasediments, and there-fore no contribution from mantle components isevident, except, perhaps, in those few zones wheresome mixing with mafic precursors has locally oc-curred. According to the isotopic composition, HOand AL granodiorites likely derived from a sourcerich in orthogneisses or metagraywackes, whereasPL and FA may derive solely from metapelites.

Geobarometric studies in the central anatecticcomplexes indicate that the depth of melting at thepeak of the anatexis was ∼15–18 km. Given thatHO, and some FA and PL plutons, are transitionalto migmatites and have nearly the same mineral-ogy, we may assume that they were also producedat the same depth, which may therefore be used asa rough estimate of the upper boundary of the an-atexis zone. To estimate the lower boundary, we

Figure 11. Chondrite-normalized REE patterns of Gredos rocks (chondritic values from McDonough and Sun 1995).Note that, whereas the mafic precursors have negative Eu anomalies and “metapelite-like” patterns, camptoniteshave positive Eu anomalies and patterns such as those common in alkaline rocks.

Journal of Geology V A R I S C A N B A T H O L I T H S I N I B E R I A 415

Figure 12. «t Nd versus «t Sr plot of Gredos rocks. Note how the range of peraluminous granitoids is exactly thesame as that of Pena Negra materials, so that mantle components cannot be identified. Note also how the maficprecursors have a less primitive isotope composition than camptonites.

must bear in mind that, with very few exceptions,Gredos granites neither contain relics of a restiticassemblage equilibrated inside the stability field ofgarnet nor does their chemistry suggest that theymay have left a garnet-bearing residuum. We there-fore assume that the source layer of granites wasinside the stability field of cordierite, at a pressureslightly lower than that of the garnet-in univariantand, hence, that the garnet-bearing felsic xenolithsfound in the camptonites had been scavenged frombelow the zone of granite production. If so, the pres-sure at which the xenolith mineral assemblageswere equilibrated (∼5.5–7.5 kbar) should also markthe maximum depth at which granites may haveoriginated, near ∼23 km. This reasoning constrainsthe productive zone to a layer about 7 km thicklocated from ∼15 km to ∼22 km in depth.

The first anatectic granites are represented by afew small bodies of leucogranites dated at ∼344 Ma,in many localities associated with strong shearzones, that were produced from protoliths with aheat production of about 4–5 mW m23 (see next par-agraph). Simple thermal modeling (Bea and Pereira1990) indicated that, with such an elevated heatproduction, the protolith may reach temperaturesclose to 7007C in a few million years if buried to adepth of ∼20 km. Local accumulation of volatiles

and deformation in shear zones would probablyhave increased the amount of melt and enhancedits segregation, thus producing the older leuco-granites. Crustal melting at the batholithic scale,however, did not start until extension was activeand the previously thickened crust had alreadybeen thinned by ∼7–8 km, as inferred from the ev-olution of the metamorphic gradient (Franco Gon-zalez and Sanchez Garcıa 1987). Generalized ana-texis started at 330–325 Ma, reached its maximumintensity at 315–310 Ma, and halted at 295–290 Ma,when extension ceased and the last compressionalphase began.

Despite the close relationship between extensionand granite generation, it is doubtful whether theincrease in heat necessary for melting was causedsolely by the rise in the subcrustal heat flux relatedto crustal thinning, because the metamorphism ofneighboring zones with presumably the same geo-dynamic evolution reached only the greenschist orlower amphibolite facies. The anomalously highconcentrations of HPE in all the materials from theAvila batholith, including migmatites and relatedmetasediments, indicate that radiogenic heat mayhave played an important role as well. The averagecurrent heat production of the Pena Negra mig-matites, which probably represent the protolith of

416 F . B E A E T A L .

Figure 13. Sr and Nd isotope evolution diagrams ofsamples from a zoned body of mafic precursors and hostanatexites, granites, and migmatites. Dots, mafic pre-cursors; crosses, host rocks. Note how mafic precursorsplot into well-defined lines corresponding to 340 Ma,which in no way may be considered mixing lines withhost rocks.

most granitoids, is ∼2.5 mW m23, much higher thanthe average crustal heat production (0.74–0.94 mWm23; Chapman and Furlong 1992). Granites are stillmore enriched in HPE, so that the current heat pro-duction of HO and AL granodiorites is ∼3 mW m23

and that of PL and FA granites is still more elevated,in the range of 4–5 mW m23. These values contrastwith those of the central Iberia metasediments spa-tially unrelated to granites, which rarely have aheat production higher than ∼1.5 mW m23 (esti-mated from 123 samples), and with those of thecrustal unit beneath the batholith, ∼0.75 mW m23

(estimated from granulitic xenoliths). We suggest,therefore, that granite production in central Iberiaoccurred during postthickening extension but onlyin those crustal layers that had a fertile composi-

tion and minimum average heat production (cal-culated at 340 Ma) of ∼2.5–3 mW m23. Recent nu-merical modeling of the thermal structure ofcollisional orogens has shown the feasibility of thismechanism (e.g., Zen 1995; Huerta et al. 1998).

Camptonites. The mineralogy, chemistry, andisotopic composition of camptonites indicate theyderived from a metasomatically enriched mantlesource with a negligible crustal component. Theelevated LaN/LuN ratios suggest they were gener-ated within the garnet stability field, consistentwith a minimum depth for the melting locus ofabout ∼60 km, inferred from the equilibration pres-sure of early phenocrysts. The maximum meltingdepth must have been about 75–80 km, marked bythe highest pressure at which a near-solidus liquidcomposition of normal silicate magma with dis-solved CO2, such as that of camptonites, is stablein the peridotite-H2O-CO2 system (White and Wyl-lie 1992). We therefore suggest that the source ofcamptonite magmas was the zone between ∼60 kmand ∼85 km in depth, corresponding to the uppermetasome layer (Haggerty 1995) of the stable lith-ospheric subcontinental mantle, slowly formed bythe action of the vapor phase released by ascendingmagmas that crystallize as they approach the sub-horizontal ledge found in the peridotite-H2O-CO2

solidus at 20–22 kbar (Wyllie 1987; White and Wyl-lie 1992). The topology of the solidus also meansthe metasome layer is a potentially productivesource of magmas, especially in a relatively hotmantle with a temperature near 11007C at ∼70 kmdepth. Given that the slope of the solidus becomesincreasingly positive from the ledge to shallowerdepths as the molar fraction of CO2 in the vaporphase increases (Wyllie 1987, fig. 9), the adiabaticdecompression of a metasomatized peridotitewould easily bring it across the solidus and producevolatile-rich alkaline mafic melts, such as thosefrom which the Gredos camptonites crystallized.Since the camptonite age indicates that melting inthe metasome layer occurred at the end of the late-orogenic extension, we suggest that the release ofpressure in the upper mantle, concomitant withcrustal thinning, was the main factor that causedthe decompressional melting of the metasomelayer. The close spatial and temporal relationshipsbetween peraluminous granites and camptoniticlamprophyres in central Iberia, hence, does not im-ply any direct genetic link between them, reflectinginstead that they were originated as a consequenceof the same geodynamic process.

This notion also finds support from the wide-spread occurrence in western and central Europe ofother camptonite dikes, notably similar to the ones

Journal of Geology V A R I S C A N B A T H O L I T H S I N I B E R I A 417

reported here, whose ages span from the Triassicto the Cretaceous (e.g., Ferreira and Macedo 1979;Dostal and Owen 1998, etc.). These dikes are ob-viously unrelated to the Variscan granite magma-tism but are associated with extensional processesof lithospheric magnitude.

Conclusions

The picture revealed so far is that magmatism inthe Avila batholith was triggered by compressionalthickening of the continental crust and reached itsmaximum intensity during the thinning of the pre-viously thickened crust, a process that seems tohave been extremely important in the entire west-ern European Variscan Belt (e.g., Escuder Viruete etal. 1994; Faure et al. 1997; Gardien et al. 1997; Aer-den 1998).

Early hybrid mafic magmas were produced dur-ing or immediately after the main compressivephase but before extension, by partial melting of amixture of mantle and biotite-bearing crustal rocksat or just below the crust-mantle interface. Thissituation could have occurred either in a melangeat the base of the subducted crustal slab below cen-tral Iberia during the main compressive stage or bythe imbrication of small wedges of mantle andcrustal materials due to low-angle faults. A mixtureof fertile peridotite and crustal materials in an8 : 2 proportion could explain the isotopic and el-emental peculiarities of the mafic precursors.These melts were emplaced in the middle crust assmall bodies aligned with Variscan structures, prob-ably along deep faults. They subsequently under-went fractional crystallization at ∼340 Ma. Theamount of heat advected to the middle crust bymafic magmas was negligible, insufficient to causeany significant melting of crustal rocks.

Local anatexis of the middle crust also occurredduring the main compressive facies. It was probablycaused by radioactive heating of fertile protolithswith an elevated heat production and enhanced bythe action of shear zones. It produced limited mig-matization and a few synkinematic leucogranitebodies dated at ∼344 Ma. The production of graniteson a batholithic scale, however, occurred from 330Ma to 295 Ma, once extension had initiated. Geo-barometric constraints indicate that the granite

source was placed between ∼15 km and ∼22 km indepth. Thermal conduction from below due to re-laxation after crustal thinning contributed to heat-ing the middle crust but, for the most part, theincrease in heat resulted from radiogenic heat pro-duced by a protolith with an anomalously high heatproduction of over 2.5–3 mW m23.

The thinning of the crust from 310 Ma to 285Ma produced a pressure release in the middle crustfrom 4–4.5 kbar to nearly subvolcanic conditions.This also affected the metasome layer of the uppermantle, which underwent decompressional melt-ing facilitated by the topology of the peridotite-H2O-CO2 solidus in the range of 22–15 kbars. Thereis no direct genetic link between camptonites andgranites, despite their close spatial and temporalassociation. They represent distinct melting eventsin those layers of the upper mantle and the middlecrust with a suitable composition, ultimatelycaused by crustal thinning and related upper-man-tle decompression during the postcollisional evo-lution of the Variscan belt in central Iberia.

The existence of a metasome layer capable of pro-ducing alkaline partial melts at the end of the Var-iscan orogeny implies that the lithospheric mantlebeneath central Iberia was relatively stable for along period before the production of camptonitemagmas; that is, it was not actively involved insubduction during the Variscan orogeny. This isconsistent with the hypothesis of an intracrustalsubduction of the Ossa Morena lower crust beneathcentral Iberia, supported by geophysical (Aller1996) and structural evidence (Azor et al. 1994; Ex-posito Ramos et al. 1998), and with the scarcity ofmantle-derived materials and mantle signatures inthe otherwise notably abundant granites that char-acterize the Central Iberian Zone.

A C K N O W L E D G M E N T S

This article has benefited from many discussionswith A. Azor, F. G. Lodeiro, and F. Simancas. Chris-tine Laurin improved the English. Revisions madeby James G. Brophy and three anonymous refereesare gratefully acknowledged. This research hasbeen financed by the Spanish DGICYT grant PB96-1266.

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