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Lithos 72 (2004) 19–44

Monzonitic series from the Variscan Tormes Dome

(Central Iberian Zone): petrogenetic evolution from

monzogabbro to granite magmas

Francisco-Javier Lopez-Moro, Miguel Lopez-Plaza*

Departamento de Geologıa, Area de Petrologıa y Geoquımica, Facultad de Ciencias, Universidad de Salamanca,

Plaza de Los Caıdos, s/n, 37008 Salamanca, Spain

Received 11 February 2003; accepted 4 August 2003

Abstract

In the Iberian Massif, rocks of the K-rich plutonic series are not abundant, but towards internal parts of the belt represented

by the Central Iberian Zone there are some sectors where shoshonitic plutonism occurs over broad areas. One of these areas is

the anatectic Tormes Dome, encompassing two similar studied plutons (Pereruela and Vitigudino). A monzonitic association

has been defined, ranging from monzogabbros to quartz monzonites or scarce monzogranites. Enrichment in LREE, P, Sr and

Ba, a high water content (up to 5.5%) and a high degree of oxidation (Ni–NiO buffer) are the main features of magmas parental

to these monzonitic rocks. These petrographic and geochemical features allow these granitoids to be ascribed to the shoshonitic

type (‘‘SH-type’’), rather than to the I-type. The minimum emplacement pressure range is 410–230 MPa, whereas the estimated

solidus temperature range is 940–765 jC; i.e., above water-saturated solidus. O, Sr, and Nd isotopes point to open-system

processes. Apatite cathodoluminescence suggests that magma mixing was unlikely to have occurred for the most enriched

rocks. Assimilation/fractional crystallisation (AFC) modelling was performed for both plutons, permitting assimilation/

crystallization rates to be estimated between 0.16 and 0.25. Different contaminants have been inferred: a metapelite at upper

crustal level for the Vitigudino Pluton and a granulitic orthogneiss for the Pereruela Pluton. A liquid line of descent, linking

monzogabbroic members to quartz monzonites/monzogranites, can be reconstructed. Some loss of water and oxygen can be

inferred, although water remained in the system below the solidus, giving rise to auto-metamorphism at ca. 500 jC. On the

other hand, AFC processes are unlikely to have been the main factor in controlling the characteristic enrichment of the

monzonitic series, since the least contaminated samples are the most LILE- and LREE-enriched for both plutons. Thus, a

source-controlled chemical signature can be inferred. The experimental data indicate that a hybrid protholith (a mixture of

peridotite, amphibolite and metapelite) can account for the enrichment and hydrated nature of the least evolved monzogabbroic

magmas. Late-orogenic slab break-off and subsequent post-collisional extensional events appear to be a suitable scenario to

provide the heat that triggered partial melting close to the crust/mantle boundary.

D 2003 Elsevier B.V. All rights reserved.

Keywords: Monzonitic series; LILE-enrichment; Crystallization conditions; AFC modelling; Tormes Dome; Iberian Massif

1. Introduction

0024-4937/$ - see front matter D 2003 Elsevier B.V. All rights reserved.

doi:10.1016/j.lithos.2003.08.002

* Corresponding author. Fax: +34-23-294514.

E-mail address: milplaz@usal.es (M. Lopez-Plaza).

Variscan shoshonitic plutonic rocks are rather

scarce and most of them appear to be dispersed along

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4420

the inner parts of the belt (Barriere, 1977; Pagel and

Leterrier, 1980; Orsini, 1980; Gil-Ibarguchi, 1982;

Michon, 1987; Sabatier, 1991; Dias and Leterrier,

1994; Holub et al., 1997). The areas from the Tormes

Dome (Central Iberian Zone) studied to date are

characterised by a widespread shoshonitic plutonism,

encompassing the monzonitic series on which this

work focuses.

Shoshonitic magmas are assumed to represent a

transitional stage between calc-alkaline and alkaline

magmas during late- and post-collisional evolution

(Liegeois et al., 1998), but their marked chemical

variability is poorly understood in terms of geo-

dynamic setting and the true source characteristics.

The first aim of this study is therefore to assess the

role of fractional crystallization, assimilation and

mixing processes, allowing the composition of the

parental magmas to be estimated. Additionally,

mineral chemistry provides an opportunity to assess

the intensive variables of these magmas during

emplacement. Finally, a basic-acid evolution is

proposed on the basis of extensive petrogenetic

modelling.

Several models have already been proposed to

explain the characteristic LILE- and LREE-enrichment

of the shoshonitic rocks. This enrichment can be

attributed to the involvement of an anomalous, amphi-

bole- or biotite-bearing lithospheric mantle domain,

enriched in incompatible elements due to mantle meta-

somatism or the recycling of a metasedimentary com-

ponent in (paleo-?) subduction zones (e.g. Tatsumi and

Eggins, 1995; Feldstein and Lange, 1999; Janousek et

al., 2000). Based on experimental results, a hybrid

(peridotite, amphibolite and metapelite) source has

been proposed to explain the origin of the monzonitic

rocks of the NW Iberian Massif as well as the connec-

tion with associated granodiorites and monzogranites

(Castro, 2001; Lopez et al., 2001).

2. Geological setting

The anatectic Tormes Dome is located in the

innermost segment of the Variscan Iberian belt (the

Central Iberian Zone) (Fig. 1). This hinterland is

mainly composed of allochthonous polymetamorphic

terrains and abundant plutonic rocks that intruded into

Proterozoic metapelites and Cadomian/Ordovician

orthogneisses. Variscan plutonism in the northwestern

Iberian Massif is mainly represented by linear belts

made up of anatectic leucogranites and biotite gran-

itoids (Capdevila et al., 1973; Martınez et al., 1990)

(Fig. 1). The latter group—the so-called ‘‘older gran-

odiorites’’ (Capdevila et al., 1973)—in some places

also includes K-rich basic and intermediate rocks, as

well as monzogranites, defining monzonitic series

(Martınez, 1974; Garcıa de los Rıos, 1981; Gil-Ibar-

guchi, 1982; Bea et al., 1987).

These K-rich plutonic rocks occur in central sectors

of the Central Iberian Zone (inset of Fig. 1) (see

compilation of Bea et al., 1987; Lopez et al., 2001;

this work). From a geochemical point of view, the

K2O/Na2O ratio increases broadly from external

zones, such as the West Asturian–Leonese Zone with

K2O/Na2Oc1, to internal ones; namely, the Central

Iberian Zone, with K2O/Na2O1 (Bea et al., 1987).

In the Tormes Dome, two biotite granitoid belts

occur: the Ifanes–Sayago Belt and the Vitigudino

Belt (Lopez-Plaza et al., 1999; Ferreira et al., 2000),

each of them including one of the plutons studied

(Pereruela and Vitigudino, respectively) (Fig. 1).

Both biotite granitoid belts border a central leucog-

ranite belt, where peraluminous leucogranites are

associated with augen gneisses (Gonzalo et al.,

1994), including small lenticular bodies of granulitic

gneisses. Field observations indicate that the relation-

ships between leucogranites and biotite granitoids are

different. On the one hand, sheet-like intrusions of

coarse-grained leucogranites are concordant with

biotite granitoid intrusions and, on the other, later

fine-grained leucogranites show a cross-cutting rela-

tionship (Lopez-Plaza, 1982), as seen in the Vitigu-

dino Belt (Fig. 1).

Following the classification of Lameyre (1980),

biotite granitoid belts from the Tormes Dome include

the following three associations (Lopez-Plaza et al.,

1999):

(a) A tonalitic–dioritic association of high-K calc-

alkaline affinity. This consists of numerous small

sill-like bodies, either intruded into metapelites

and augen gneisses, or bordering the Ifanes–

Sayago Belt (Fig. 1) (Lopez-Plaza and Gonzalo,

1993).

(b) A granodioritic association, with vaugnerites

(monzodiorites and quartz monzodiorites), as well

Fig. 1. Simplified geological map of the Tormes Dome (Lopez-Plaza et al., 1999) and sample locations. Inset includes occurrences of K-rich

plutonic rocks from NW of Iberian Massif.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–44 21

as granodiorites and monzogranites. The vaugner-

itic rocks of this association occur as stocks either

enclosed by biotite granitoids or intruded into

augen gneisses (Fig. 1).

(c) A monzonitic association (MATD) of shosho-

nitic affinity, with monzodiorites (monzogab-

bros), monzonites, quartz monzonites and

monzogranites. Melanosyenitic rocks (ultrapo-

tassic) are also found locally (Lopez-Plaza et al.,

1999). MATD rocks occur as heterogeneous

outcrop-scale bodies consisting of basic and

intermediate rocks, porphyritic quartz monzon-

ites and limited monzogranites. In the case of the

Pereruela Pluton, this rock association is enclosed

by biotite (Fmuscovite) granites (Ifanes–Sayago

Belt), whereas in the Vitigudino Belt MATD

rocks mainly have metapelitic country rock,

metapelitic xenoliths appearing in the intermedi-

ate rocks. The relationship between basic and

felsic members of MATD rocks is equivocal.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4422

Although monzonitic rocks seem to be slightly

earlier than the porphyritic quartz monzonites

because the latter occur as dyke-like intrusions,

locally developed hybrid facies suggest they are

roughly coeval.

An extensional event (D2) is assumed to have

occurred in the Tormes Dome after a previous com-

pressional thickening regime (Escuder et al., 1994).

The extension was followed by a later compressional

event (D3), probably related to strike–slip shear zones

(Lopez-Plaza, 1982).

The MATD rocks underwent both extensional and

later compressional events (D2 and D3). Thus, low-

angle normal shear zones appear to be related to the

emplacement of porphyritic quartz monzonites, show-

ing a pervasive planar subhorizontal fabric that was

gently folded by the third Variscan phase (Lopez-

Plaza, 1982). K-feldspar megacrysts of quartz mon-

zonites were deflected around microgranular enclaves,

and at the same time the deformation related to

Fig. 2. R1–R2 diagram by de la Roche et al. (1980) plotting samples of th

filled circles: Pereruela Pluton. Large circles: amphibole-bearing rocks;

Monzogabbro, (3) Monzonite, (4) Quartz monzonite, (5) Monzogranite an

strike–slip shear zones resulted in a subvertical foli-

ation at brittle–ductile transition, indicating a super-

imposed solid-state fabric (Lopez-Plaza, 1982; Lopez-

Moro, 2000).

Despite the lack of geochronological data for the

area, an age of 320F 5 Ma (207Pb/235U on monazite)

(Ferreira et al., 2000) has been obtained in Portugal

for the Ifanes Granite of the granodioritic association.

This age is within the presumed time span of the

extensional event D2 (340–320 Ma) (Escuder et al.,

1994; Fernandez-Suarez et al., 2000).

3. Rock types

The R1–R2 diagram by de la Roche et al. (1980)

was used for classification purposes (Fig. 2). Some of

the mafic rocks are plotted onto the undersaturated

field, and they are classified as syenogabbros and

monzogabbros. Nevertheless, according to I.U.G.S.

recommendations (Le Maitre, 1989), the former term

e plutons studied. Open circles: samples from the Vitigudino Pluton;

small circles: amphibole-free rocks. Fields: (1) Syenogabbro, (2)

d (6) Syenogranite.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–44 23

should be avoided and in the present work the term

monzogabbro has been used for both groups of

subsaturated rocks. The rest of samples plot within

the fields of monzonites, quartz monzonites and

monzogranites. Both plutons studied display similar

trends, except that no monzogranites have been found

in the Vitigudino Pluton.

Broadly speaking, two large groups of MATD

rocks can be established according to their modal

compositions.

3.1. Amphibole-bearing rocks

These range from monzogabbro to monzonite (Fig.

2; Table 1). Medium/fine-grained equigranular tex-

tures are common, although plagioclase and amphi-

bole phenocrysts in a fine-grained groundmass can be

sometimes found.

These contain plagioclase, amphibole (magnesio-

hornblende), K-feldspar and sometimes quartz as

major minerals, as well as diopsidic clinopyroxene

(En38Fs15Wo48), apatite, titanite, allanite and magne-

tite as accessory minerals (Table 1). The amphibole

content (3–27% modal) increases with the colour

index. Amphibole crystals occur either as clots or

Table 1

Mineral assemblages and modal compositions (vol.%) of the plutons stud

Pluton/

sample

Rock

type (1)

Rock

type (2)

Q

(%)

Pl

(%)

Kfs

(%)

Cpx

(%)

H

(

Pereruela

(1) Per B Mzd Mzgb 0.21 39.65 12.29 relict

(2) Per 9 Mzd n.d. 0.80 36.10 16.62 relict

(3) Arc 2 Mz n.d. 1.58 32.65 17.34 relict

(4) POR 15 Qmz Mz 5.19 37.88 24.40 absent

(5) POR 14 Qmz Qmz 10.2 35.6 27.2 absent a

(6) POR 14B Mzg Mzg 23.0 31.6 29.7 absent a

Vitigudino

(7) POR 106 Mzd Mzgb 0.50 40.79 16.41 relict

(8) E.76 Mzd n.d. 0.30 23.77 10.47 5.12

(9) E.84 Mz n.d. 0.17 21.94 22.88 relict

(10) D.60 Mz n.d. 0.62 31.48 32.81 relict

(11) E.3 Mz n.d. 0.97 34.92 36.42 relict

(12) POR 105 Qmz Qmz 14.8 35.5 28.9 absent a

(13) POR 109 Qmz Qmz 19.23 24.73 44.77 absent a

Rock type (1): according to Streckeisen’s (1973) classification; Rock type

syenogabbro; Mzd: monzodiorite; Mz: monzonite; Qmz: quartz-monzonite;

other accessory minerals (magnetite, zircon and xenotime); n.d.: not determ

12 and 13, Lopez-Moro, (2000); 8, 9, 10 and 11, Martınez (1974).

isolated euhedral crystals in which relict clinopyrox-

ene crystals have been found, supporting a secondary

origin for the amphibole.

Biotite crystals (15–42% modal) range from

idiomorphic to xenomorphic. They are hosted by

zoned plagioclase and in turn enclose apatite and

partially include allanite grains. The Fe* ratio [Fe2 +/

(Mg + Fe2 +)] in biotite hardly varies (0.42–0.43) and

the Ba content is not very high (BaO wt.%: 0.13–

0.16) (Table 2). Zoned plagioclase (22–41% modal)

ranges from An54–35 in monzogabbros to An37–24 in

monzonites, although exceptionally oscillatory zon-

ing also occurs. K-feldspar (10–36% modal) occurs

as anhedral or rounded crystals (ocelli and orbicules)

(Lopez-Moro, 2000), the latter probably autoliths, as

suggested by Eklund et al. (1998) for Svecofennian

rocks. The anhedral K-feldspar shows monoclinic

symmetry with 2Vx angle values < 60j; it is rich in

orthoclase and celsian components (Or85–93, Cn1–3)

(Lopez-Moro, 2000).

3.2. Amphibole-free rocks

In this group, biotite and magnetite are the only

mafic minerals. Compositionally, they range from

ied

bl

%)

Bt

(%)

Mus

(%)

Ap

(%)

Aln

(%)

Spn

(%)

Acc.

(%)

Colour

index

9.70 36.14 absent 1.71 absent absent 0.34 46.14

3.25 41.73 absent 0.54 n.d. n.d. 1.04 45.94

20.77 26.37 absent 1.37 absent 0.03 0.60 47.06

6.04 22.85 absent 1.78 0.85 0.85 0.15 30.75

bsent 10.1 absent 1.2 absent absent n.d. 25.80

bsent 3.0 absent 0.5 absent absent n.d. 15.20

18.10 21.48 absent n.d. absent n.d. 2.71 42.30

26.98 28.28 absent n.d. n.d. n.d. 5.04 65.42

16.14 26.47 absent n.d. n.d. n.d. 12.38 54.99

14.37 17.26 absent n.d. n.d. n.d. 3.43 35.06

8.50 15.15 absent n.d. n.d. n.d. 3.80 27.45

bsent 19.5 0.6 0.2 0.1 0.2 n.d. 20.00

bsent 4.36 4.19 0.61 absent absent 0.09 6.47

(2): according to classification of de la Roche et al. (1980); Sygb:

Mzg: monzogranite. Mineral abbreviations after Kretz (1983). Acc.:

ined. Data source: 1, 2 and 3, Gomez-Hernandez (1993); 4, 5, 6, 7,

Table 2

Biotite composition of mineral separates

Vitigudino Pluton Pereruela Pluton

Sample

rock

POR

106B

Mzgb

POR

109

Qmz

POR

15

Mz

POR

14

Qmz

POR

14B

Mzgr

SiO2 37.34 34.47 36.80 36.75 34.52

TiO2 2.49 3.12 2.87 3.20 3.57

Al2O3 15.71 18.34 14.57 15.59 17.21

Cr2O3 0.02 0.01 0.02 0.02 0.02

Fe2O3 1.86 2.08 2.45 4.02 2.77

FeO 15.38 17.66 15.05 16.05 16.36

MnO 0.21 0.24 0.25 0.26 0.29

MgO 13.33 9.29 12.69 10.61 10.89

CaO 0.60 0.47 0.05 0.02 0.27

Na2O 0.16 0.21 0.12 0.08 0.22

K2O 9.78 8.84 9.50 9.70 9.55

BaO 0.16 0.03 0.13 0.06 0.05

Rb2O 0.08 0.07 0.06 0.08 0.06

Li2O 0.10 0.12 0.06 0.07 0.08

V2O5 0.04 0.02 0.03 0.03 0.02

ZnO 0.04 b.d. 0.04 0.03 0.04

CuO 0.01 b.d. b.d. b.d. b.d.

H2O n.d. 4.14 3.68 3.45 3.84

Total 97.31 99.11 98.37 100.00 99.75

Si 5.53 5.25 5.63 5.58 5.25

AlIV 2.47 2.75 2.37 2.42 2.75

AlVI 0.27 0.54 0.25 0.38 0.33

Ti 0.28 0.36 0.33 0.37 0.41

Cr 0.00 0.00 0.00 0.00 0.00

Fe3 + 0.21 0.24 0.28 0.46 0.32

Fe2 + 1.90 2.25 1.92 2.04 2.08

Mn 0.03 0.03 0.03 0.03 0.04

Mg 2.94 2.11 2.89 2.40 2.47

V 0.00 0.00 0.00 0.00 0.00

Zn 0.00 0.00 0.00 0.00 0.00

Cu 0.00 0.00 0.00 0.00 0.00

Li 0.06 0.07 0.04 0.04 0.05

vi site 5.69 5.59 5.76 5.73 5.70

Ca 0.10 0.08 0.01 0.00 0.04

Na 0.05 0.06 0.04 0.02 0.06

K 1.85 1.72 1.85 1.88 1.85

Ba 0.01 0.00 0.01 0.00 0.00

Rb 0.01 0.01 0.01 0.01 0.01

A site 2.00 1.86 1.91 1.92 1.97

OH� n.d. 4.20 3.75 3.50 3.90

Fe* 0.42 0.54 0.43 0.51 0.49

Rock abbreviations as in Table 1; cations per formula unit based on

24 O atoms; FeO by wet chemistry and H2O by manometry;

Fe*= Fe2 +/(Fe2 + +Mg); n.d.: not determined; b.d.: below detection

limit.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4424

quartz monzonite to monzogranite (Table 1; Fig. 2).

They are porphyritic granitoids with euhedral crys-

tals of K-feldspar enclosed within a medium/coarse-

grained groundmass. Plagioclase, quartz and biotite

are another major minerals and, exceptionally, white

mica (Vitigudino Pluton). Accessory minerals in-

clude apatite, zircon, xenotime, allanite and magne-

tite. Chlorite, titanite, epidote and sericite are

common secondary phases related to deuteric pro-

cesses. The plagioclase (25–36% modal) is found

in the groundmass or is hosted by K-feldspar

megacrysts. It is zoned (An30–17), although subhe-

dral albitized crystals appear to be related to perthi-

tization processes (Schermerhorn, 1956). The K-

feldspar (27–45% modal) occurs as megacrysts,

but scarce interstitial crystals can also be found in

the groundmass. Its structural state varies from

orthoclase to intermediate microcline (Lopez-Moro

et al., 1998). The biotite (3–20% modal) is more

abundant in the least evolved members. Biotite

from the felsic facies of the Vitigudino Pluton is

richer in Fe (Fe* = 0.54) than that from the Perer-

uela Pluton (Fe* = 0.49) (Table 2) despite the spe-

cific occurrence of more felsic granitoid rocks

(monzogranites) in the latter body. In both plutons,

biotite occurs as subidiomorphic crystals, either in

the groundmass or hosted by plagioclase and K-

feldspar, although it never appears in plagioclase

cores, suggesting an overlapping crystallization in-

terval of biotite and plagioclase, except for the

earliest plagioclase crystals.

4. Accessory minerals

One of the most striking features in the MATD

rocks is the abundance of apatite, allanite and, to a

lesser extent, titanite in monzonites (allanite: 0.9%

modal; apatite: 1.8% modal; Table 1). In contrast to

the amphibole-bearing rocks that do not contain

zircon, the amphibole-free ones have on average four

zircon inclusions per biotite crystal. This value is not

higher than the amounts of the granodioritic and

tonalitic–dioritic associations (average five zircon

inclusions), nor of those of the anatectic granites from

the Tormes Dome (average six zircon inclusions)

(Domınguez-Vadillo, 1991; this work).

The study of apatite by cathodoluminescence pro-

vides some clues to its origin. Apatite in the mon-

zonites of the MATD always occurs as zoned crystals,

with a large blue core mantled by a yellowish rim.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–44 25

Apatite grains for the porphyritic quartz monzonite

and monzogranites show small pale blue core and a

thick yellowish rim, although in most cases the entire

crystal is yellow in colour.

According to Mariano and Ring (1975), Mn2 +-rich

apatites are yellowish, whereas those enriched in Eu2 +,

Sm3 + and Dy3 + show violet–blue colours. The tran-

sition in apatite grains from blue to yellow, as well as

zoned cores from deep blue to light blue, as is the case

in the most evolved rocks, would be possible as long as

REE fractionation processes were involved (Wenzel

and Ramseyer, 1992). Thus, an early fractionation

stage could be represented by the apatite cores.

Bearing in mind that the apatites from the mon-

zonites do not exhibit yellow cores, a significant

Fig. 3. Harker diagrams. Shaded field: vaugnerites from the Tormes Dome

are from Peccerillo and Taylor (1976). Symbols as in Fig. 2.

involvement of restitic crustal material as well as of

refractory residua from an assimilated upper/middle

crust can be ruled out.

5. Geochemistry

Harker diagrams for the MATD rocks (Fig. 3)

show shoshonitic affinity on the one hand and, on

the other, an almost general decrease in the content

of most elements, except partly for Na2O and P2O5.

K2O and P2O5 contents of MATD rocks are higher

than in other shoshonitic intermediate rocks from

the Tormes Dome (vaugnerites). It is worth noting

that samples from both plutons (Pereruela and

(Lopez-Moro, 2000). Separating lines in K2O versus SiO2 diagram

Table 3

Major (wt.%), trace (ppm) and isotopic composition of representative samples

Sample Pereruela Pluton Vitigudino Pluton Wallrocks

PER-B

(1) Mzgb

POR-15

Mz

PER-A

(1) Mz

POR-14

Qmz

POR-111

Mzgr

POR-14B

Mzgr

9143

(2)

Mzgb

POR-106

Mzgb

9141

(2)

Mzgb

POR-105

Qmz

POR-109

Qmz

9144

(2) Qmz

POR-86

Ggo

DT-70

Mtp

SiO2 54.22 57.95 58.16 63.85 69.72 70.68 52.03 52.78 54.15 62.54 66.75 67.30 58.87 59.18

TiO2 0.49 0.87 0.40 0.60 0.23 0.18 0.63 1.18 0.79 1.01 0.34 0.34 0.56 0.94

Al2O3 18.18 17.51 16.15 16.94 15.19 15.97 15.28 17.05 18.07 16.64 17.09 16.2 18.76 19.23

Fe2O3 n.d. 0.51 n.d. 0.06 0.20 0.04 n.d. 0.68 n.d. 0.49 0.19 n.d. n.d. n.d.

FeO n.d. 4.60 n.d. 3.50 1.12 1.00 n.d. 4.85 n.d. 2.80 1.27 n.d. n.d. n.d.

Fe2O3t 4.91 5.17 4.88 3.57 1.34 1.04 6.52 5.61 7.47 3.35 1.48 2.3 6.10 7.75

MnO 0.06 0.06 0.05 0.03 0.02 0.02 0.11 0.11 0.13 0.06 0.02 0.03 0.04 0.09

MgO 4.58 3.45 3.77 1.86 0.73 0.57 7.95 5.46 3.61 2.21 0.78 1.09 1.70 2.76

CaO 5.45 4.25 4.33 2.46 1.33 1.33 7.11 6.54 5.57 2.71 1.35 1.39 2.21 0.62

Na2O 3.39 3.26 3.64 3.17 2.87 3.12 2.86 2.80 3.61 3.26 3.15 3.28 5.99 1.23

K2O 5.22 5.50 5.29 6.06 7.06 6.48 5.18 5.20 4.69 5.62 7.57 6.64 2.84 4.35

P2O5 0.88 0.90 0.98 0.52 0.22 0.21 0.65 0.62 0.85 0.55 0.34 0.30 0.27 0.17

L.O.I. 1.09 0.10 1.79 0.50 0.97 0.55 1.5 1.76 0.85 1.28 0.82 0.80 1.33 3.38

Total 98.47 98.96 99.45 99.55 99.66 100.15 99.82 99.03 99.79 99.17 99.67 99.93 98.67 99.69

K2O/Na2O 1.54 1.69 1.45 1.91 2.46 2.08 1.81 1.86 1.30 1.72 2.40 2.02 0.47 3.54

Fe* n.d. 0.57 n.d. 0.65 0.61 0.64 n.d 0.47 n.d 0.56 0.62 n.d n.d. n.d.

A/CNK 0.82 0.92 0.78 1.04 1.03 1.10 0.66 0.77 0.85 1.02 1.08 1.07 1.11 2.45

F n.d. n.d. n.d. n.d. n.d. n.d. 1532 n.d. 2185 n.d. n.d. 1004 n.d. n.d.

Rb 150 174.8 160 232.7 176.0 180.0 160 207 146 231.0 254.0 267 190 248

Cs 7 5.8 6 7.5 n.d. 4.0 n.d. 13.5 n.d. n.d. 5.2 n.d. 17 n.d.

Sr n.d. 1596 n.d. 697 632 670 1443 1116 1853 613 366 330 2120 97.9

Ba 2500 2776 2100 1623 1782 1957 4127 2905 2551 1560 1111 853 220 514

Pb 50 59.1 46 70.8 n.d. 84.8 87 46.8 89 n.d. 80.9 74 13.4 n.d.

Sc 17 n.d. 13 n.d. n.d. n.d. n.d. 20.1 n.d. 17.7 n.d. n.d. n.d. 10.1

V n.d. 102.0 n.d. 55.4 n.d. 15.8 153 152.0 125 58.4 19.9 29 57.2 113

Cr 99 116 68 127 n.d. 12 n.d. 128 114 n.d. 12 15 38.5 95.8

Co n.d. 16.0 n.d. 8.5 n.d. 91.7 n.d. 33.8 n.d. n.d. 58.1 n.d. 39.3 62.5

Ni n.d. 50.5 n.d. 23.4 n.d. 8.6 n.d. 77.8 26 n.d. 6.2 b.d.l. 20 43.2

Y n.d. 27.5 n.d. 16.1 n.d. 5.1 25 23.5 15 n.d. 9.5 5 38.3 31.2

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Zr n.d. 380 n.d. 282 n.d. 112 201 241 227 n.d. 132 166 200 205

Nb 18 28.1 22 12.4 n.d. 4.6 15 10.0 11 n.d. 6.2 11 13.1 18.2

Hf 7.80 8.4 7.30 7.3 n.d. 2.8 n.d. 6.0 n.d. n.d. 3.4 n.d. 5.41 n.d.

Ta n.d. 1.4 n.d. 1.0 n.d. 1.6 n.d. 0.9 n.d. n.d. 1.0 n.d. 1.95 n.d.

Th 28 42.9 32 22.4 n.d. 10.5 n.d. 46.0 n.d. n.d. 19.8 n.d. 14.9 n.d.

U 3.9 6.6 4.30 8.5 n.d. 1.9 n.d. 8.6 n.d. n.d. 4.6 n.d. 7.49 n.d.

La 144 195.3 176 57.13 n.d. 25.29 n.d. 164.6 n.d. n.d. 50.18 n.d. 34.1 58.21

Ce 246 369.5 300 119.4 n.d. 49.06 n.d. 328.4 n.d. n.d. 106.8 n.d. 73.6 117.80

Pr n.d. 38.9 n.d. 12.58 n.d. 4.81 n.d. 35.01 n.d. n.d. 10.61 n.d. 7.84 13.46

Nd 104 142.1 121 47.02 n.d. 17.59 n.d. 133.4 n.d. n.d. 38.97 n.d. 30.3 50.46

Sm 15.00 18.52 16.00 8.24 n.d. 2.75 n.d. 18.57 n.d. n.d. 6.79 n.d. 7.11 9.50

Eu 3.80 4.4 3.60 2.0 n.d. 1.84 n.d. 3.97 n.d. n.d. 1.66 n.d. 1.24 1.42

Gd 6.20 11.21 8.20 5.26 n.d. 1.68 n.d. 10.67 n.d. n.d. 3.87 n.d. 5.92 7.58

Tb 1.00 1.4 1.10 0.71 n.d. 0.21 n.d. 1.33 n.d. n.d. 0.54 n.d. 1.03 1.10

Dy 4.90 6.11 5.50 3.42 n.d. 1.03 n.d. 5.63 n.d. n.d. 2.12 n.d. 6.10 5.81

Ho n.d. 1.02 n.d. 0.58 n.d. 0.18 n.d. 0.90 n.d. n.d. 0.34 n.d. 1.35 1.17

Er n.d. 2.67 n.d. 1.43 n.d. 0.45 n.d. 2.22 n.d. n.d. 0.77 n.d. 3.38 2.90

Tm n.d. 0.28 n.d. 0.22 n.d. 0.07 n.d. 0.31 n.d. n.d. 0.09 n.d. 0.61 0.43

Yb 2.00 2.21 2.00 1.32 n.d. 0.51 n.d. 1.81 n.d. n.d. 0.69 n.d. 3.62 2.80

Lu 0.3 0.30 0.27 0.21 n.d. 0.08 n.d. 0.26 n.d. n.d. 0.09 n.d. 0.53 0.41

SREE 527.2 794.0 633.7 259.5 n.d. 105.6 n.d. 707.1 n.d. n.d. 223.5 n.d. 176.82 273.05

(La/Sm)N 5.92 6.5 6.79 4.3 n.d. 5.7 n.d. 5.5 n.d. n.d. 4.6 n.d. 2.64 3.36

(La/Yb)N 48.15 59.1 58.84 28.9 n.d. 33.5 n.d. 60.8 n.d. n.d. 48.7 n.d. 6.77 14.92

(Gd/Yb)N 2.47 4.0 3.27 3.2 n.d. 2.7 n.d. 4.7 n.d. n.d. 4.5 n.d. 1.31 2.17

(Eu/Eu*) 1.21 0.9 0.97 0.9 n.d. 2.6 n.d. 0.9 n.d. n.d. 1.0 n.d. 0.59 0.51

d18O n.d. 8.4 n.d. 8.7 n.d. 9.0 n.d. 9.8 n.d. 11.1 11.0 n.d. 9.6 11.7

87Rb/86Sr n.d. 0.3169 n.d. 0.9664 n.d. 0.7763 n.d. 0.5370 n.d. n.d. 2.0113 n.d. 0.25000 3.908387Sr/86Sr n.d. 0.708523(6) n.d. 0.712251(6) n.d. 0.711331(6) n.d. 0.707188(6) n.d. n.d. 0.720911(6) n.d. 0.713795(6) 0.733010(7)87Sr/86Sr320 n.d. 0.707079 n.d. 0.707849 n.d. 0.707795 n.d. 0.707035 n.d. n.d. 0.711751 n.d. 0.712656 0.715210147Sm/144Nd n.d. 0.0788 n.d. 0.1059 n.d. 0.0945 n.d. 0.0841 n.d. n.d. 0.1053 n.d. 0.12516 0.1312143Nd/144Nd n.d. 0.512241(4) n.d. 0.512231(4) n.d. 0.512139(4) n.d. 0.512222(4) n.d. n.d. 0.512113(4) n.d. 0.511978(4) 0.511991(4)143Nd/144Nd320 n.d. 0.512076 n.d. 0.512010 n.d. 0.511941 n.d. 0.512045 n.d. n.d. 0.511892 n.d. 0.511716 0.511715

eNd320 n.d. � 2.9 n.d. � 4.2 n.d. � 5.5 n.d. � 3.5 n.d. n.d. � 6.5 n.d. � 9.9 � 9.9

Mzgb: monzogabbro; Mz: monzonite; Qmz: quartz monzonite; Mzgr: monzogranite; Ggo: garnet-bearing granulitic orthogneiss; Mtp: metapelite; Fe2O3t: total iron; L.O.I.: loss on ignition; n.d.: not determined; b.d.l.: below detection limit;

Fe*: FeO/(FeO +MgO); A/CNK: molecular Al2O3/(CaO+Na2O+K2O); REE-normalizing values from Nakamura (1974). Errors in brackets are 2 S.D.. (1) Data from Gomez-Hernandez (1993); (2) data from ITGE (2000).

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F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4428

Vitigudino) follow similar trends. The A/CNK in-

dex ranges from 0.66 to 1.08 for the Vitigudino

Pluton and from 0.78 to 1.10 for the Pereruela

Pluton (Table 3). The R1–R2 diagram (Fig. 2)

shows a typical curvilinear trend for both plutons

from undersaturated fields to quartz monzonites or

granites.

Fig. 4. Binary diagrams for trace el

The K2O/Na2O ratio is high, ranging between 1.45

and 2.46 for the Pereruela Pluton (average 1.86) and

between 1.30 and 2.40 for the Vitigudino Pluton

(average 1.85) (Table 3).

A significant feature of the MATD is the scarcity of

incompatible elements, although exceptionally Pb

correlates moderately well with silica (Fig. 4). Some

ements. Symbols as in Fig. 2.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–44 29

other elements, such as Rb, are also incompatible, but

poorly correlated with silica.

Among the compatible elements, the V vs. silica

diagram displays a good correlation (Fig. 4). The most

mafic rocks show similar concentrations of Cr, Ni and

V in both plutons (Table 3, Fig. 4). Cr and Ni contents

Fig. 5. (a) Chondrite-normalised multi-element patterns (chondrite values ta

(values from Nakamura, 1974). Pm: interpolated values. Shaded field: Viti

in the inset.

are lower than those proposed for primary melts of

mantle peridotites (Kuehner et al., 1981).

The switch from incompatible to compatible be-

haviour (see inflections in the diagrams of silica vs.

Nb, Y, and Yb, Fig. 4) is, however, common in the

MATD. Strong enrichments in Ba, Rb, Th, K and

ken from Thompson, 1982). (b) Chondrite-normalised REE patterns

gudino Pluton. Values of silica weight percent are shown in brackets

Fig. 6. Isotope diagrams. (a) eNd320 vs. 1/Nd showing the different

trends for the Pereruela and Vitigudino plutons; (b) d18O vs. SiO2

(wt.%) showing subparallel and decoupling trends for both plutons.

Fractional Crystallization (FC) trend after Matsuhisa (1979).

Symbols as in Fig. 2.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4430

LREE are typical mainly of the mafic rocks (Fig. 5a

and b). An apparent arc-affinity can be inferred from

the negative anomaly in Nb, Ta and Ti, together with

an enrichment in LILE and LREE, mainly in mon-

zonites and monzogabbros (LaN/YbN = 61–48). Sub-

parallel REE patterns are common for both plutons

(Fig. 5b), a slight negative Eu anomaly being ob-

served for a majority of the samples (Eu/Eu* = 0.9–

1.0), except for extreme members from the Pereruela

Pluton; namely, the monzogabbro and the monzog-

ranite, with slight and strong positive anomalies,

respectively (Fig. 5b; Table 3). The negative Sr

anomaly is also worth noting and probably indicates

prior plagioclase fractionation (Fig. 5a).

High Rb and Th contents give rise to an upward

convexity (Fig. 5a), except in the K-feldspar enriched

sample POR-14B. Rb and Th enrichment is common

in other K-rich rocks, such as vaugnerites, from the

Tormes Dome (Lopez-Plaza et al., 1999).

Regarding the isotopic data, monazite age (320

Ma) obtained for biotite granites from the Portuguese

sector of the Ifanes–Sayago Belt provides a reference

age in this work, which could be viewed as minimal

since the MATD rocks are assumed to be earlier than

the biotite granites.

eNd320 and (87Sr/86Sr)320 values (Table 3; Fig. 6a)

range from � 2.9 to � 5.5, and 0.707079–0.707849

for the Pereruela Pluton, and from � 3.5 to � 6.5 and

0.707035–0.711751 for the Vitigudino Pluton. These

data indicate slight differences for both plutons, main-

ly regarding the most mafic rocks, and point to large

variation in the Nd isotopic system, combined with

narrow ranges in Sr isotopes for the Pereruela Pluton.

The Vitigudino Pluton displays a relatively more

pronounced variation in Sr.

Concerning oxygen isotopes, MATD rocks show

lower contents in d18O for the Pereruela Pluton, with a

narrow range ( + 8.4xto + 9.0x), whereas samples

from the Vitigudino Pluton reflect a more pronounced

crustal character, having higher values and a broader

range ( + 9.8 to + 11.1) (Table 3). The d18O values for

rocks from the Pereruela Pluton correlate well with

silica (Fig. 6b); better than those from the Vitigudino

Pluton.

The overall geochemical features of MATD rocks

described above provide some insight into granitoid

typology. These features certainly do not match I-type

granites, but rather HiBaSr-type, according to trace

element geochemistry considered by Tarney and Jones

(1994). Nevertheless, the MATD rocks have higher

K2O/Na2O ratio than many HiBaSr-granitoids, such

as some Caledonian appinitic rocks (Fowler and

Henney, 1996) as well as the vaugneritic rocks from

the Tormes Dome (Lopez-Plaza et al., 1999).

By contrast, the MATD rocks correspond rather

well to the shoshonitic type (SH-type) established by

Jiang et al. (2002). Firstly, this is supported by the

fact that the MATD rocks have the same association

(monzogabbro/monzodiorite–monzonite–monzog-

ranite). Secondly, the following geochemical similar-

ities with the SH-type can be found (Table 4): (a) a

high K2O and K2O/Na2O ratio, as well as a high

P2O5 and low SiO2/P2O5 ratio, (b) a high LREE and

LREE/HREE ratio, (c) a high content of some LILE,

such as Sr and Ba, as well as F and (d) a relatively

high eNdt and a relatively wide range in (87Sr/86Sr)t.

Table 4

Comparison of average compositions of I- and S-type granitoids in the world from Whalen et al. (1987) and Barbarin (1999), shoshonitic

granitoids from Jiang et al. (2002) (SH-1) and shoshonitic granitoids from the Tormes Dome (SH-2)

Genetic type S-type I-type SH-1 SH-2

Number of

samples

577 991 21 10a

SiO2 70.27 69.17 63.12 62.53

TiO2 0.48 0.43 0.57 0.53

Al2O3 14.10 14.33 15.37 16.79

Fe2O3 0.56 1.04 2.12 0.15

FeO 2.87 2.29 2.11 1.43

MnO 0.06 0.07 0.09 0.05

MgO 1.42 1.42 1.57 2.27

CaO 2.03 3.20 4.19 3.02

Na2O 2.43 3.13 3.65 3.28

K2O 3.96 3.40 5.48 6.01

P2O5 0.15 0.11 0.28 0.58

Na2O+K2O 6.37 6.53 9.14 9.29

K2O/Na2O 1.64 1.09 1.50 1.86

FeO/MgO 2.38 2.27 2.56 1.52

SiO2/P2O5 468 628 225 108

A/CNK 1.18 0.98 0.78 0.97

Fb (ppm) 895 491 1825 1573

Rb (ppm) 217 151 234 197

Ba (ppm) 468 538 2756 1881

Sr (ppm) 120 247 1020 635

Th (ppm) 18 18 53.4 22

Zr (ppm) 165 151 257 162

Ce (ppm) 58 66 145 198

LREE/

HREE

– – 7.13 29.19

eNdt � 4 to � 17 � 4 to � 9 + 1.4 to � 7.3 � 2.9 to � 6.5

(87Sr/86Sr)t 0.706–0.760 0.706–0.712 0.709–0.712 0.707–0.712

d18O (x) + 10 to + 14 + 5 to + 10 + 11.6 to + 11.9 + 8.4 to + 11.1

a The most basic rocks are not included for comparative purposes.b Data from Jiang et al. (2002) except samples from the Tormes Dome (SH-2). eNdt, (

87Sr/86Sr)t and d18O for S- and I-types from Barbarin

(1999).

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–44 31

Finally, MATD rocks show mineralogical similarities,

such as a high magnesium number in biotite, ranging

from 0.58 to 0.46 (Table 2), similar to the SH-type

(0.65–0.51).

6. Discussion

6.1. Crystallization conditions

6.1.1. Temperature

6.1.1.1. Amphibole-bearing rocks. Liquidus temper-

atures for amphibole-bearing rocks were estimated

(Table 5) on the basis of plagioclase and whole-rock

composition, using the geothermometers of Kudo and

Weill (1970) and Mathez (1973). The methodology of

Cotkin and Medaris (1993) was also used, considering

anorthite contents in plagioclase cores, as long as these

did not represent reabsorbed crystals. The temperatures

obtained are consistent, being 1092–1065 jC for

monzogabbros and 970–937 jC for monzonites, i.e.,

comparable to those of other monzonitic series (Duch-

esne et al., 1998). Using the apatite saturation model of

Harrison and Watson (1984), the temperatures for

some monzogabbroic rocks (PER-B and 9141) are

similar to those calculated with the plagioclase-melt

geothermometer (Table 5). Nevertheless, the least

Table 5

Liquidus and solidus temperatures and estimated pressures

Sample Vitigudino Pluton Pereruela Pluton

9143

Mzgb

POR-106

Mzgb

9141

Mzgb

POR-105

Qmz

POR-109

Qmz

9144

Qmz

PER-B

Mzgb

POR-15

Mz

PER-A

Mz

POR-14

Qmz

POR-14B

Mzgr

Temperature (C)

(L) Pl-melt

(F 60 jC)1087–1058 1092–1065 1063–1051 902–882 856–815 880–853 1060–1048 970–937 951–926 875–864 864–805

(L) Apatite

saturation

(F 25 jC)

a.m.u. a.m.u. a.m.u. 1100 1025 1000 1020 1100 1130 1050 980

(L) Zircon

saturation

(F 25 jC)

a.m.u. a.m.u. a.m.u. – 800 830 a.m.u. a.m.u. a.m.u. 865 790

(S) Fe3 + in Bt – 880 – – 800 – – 840 – 765 800

(S) Fe3 + in Bt

and bulk rock

(F 7 jC)

– 940 – – >1000 – – 890 – >1000 >1000

(S) Hbl–Pl

(F 40 jC)– 604–600 – – – – – 657–648 – – –

(S) Ab–Kfs

(F 40 jC)– 580–534 – – – – – 537–450 – 600–502 535–488

Pressure (MPa)

Al in Hbl

(F 60 MPa)

– 130 – – – – – 250–210 – – –

Bt (ASM)

(F 100 MPa)

– – – – 300 – – 410–345 – 230 300

Geothermometer applications: Pl-melt, after Kudo and Weill (1970) and Mathez (1973); apatite saturation, after Harrison and Watson (1984);

zircon saturation, after Watson and Harrison (1983); Hbl–Pl, after Holland and Blundy (1994); Fe3 + in biotite, after Wones and Eugster (1965);

Fe3 + in biotite and bulk rock, after Burkhard (1991); Ab–Kfs, after Fuhrman and Lindsley (1988); (L): liquidus-estimated temperature; (S):

solidus-estimated temperature; figures in brackets: errors of calibrations; a.m.u.: accesory-mineral undersaturated.

Geobarometric applications: Al in Hbl, after Anderson and Smith (1995); Bt (ASM) according to the fH2O of Wones (1972) by considering the

solidus-estimated temperature from Fe3 + in biotite and Fe3 + in biotite and bulk rock, as well as an oxidation state in the Ni–NiO buffer.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4432

differentiated rocks from the Vitigudino Pluton (POR-

106 and 9943) give temperatures of around 950 jC,lower than those obtained by plagioclase thermometry

(around 1100 jC), indicating that the former may be

underestimated. The opposite appears to be the case for

the monzonites from the Pereruela Pluton, whose

estimations made with the apatite solubility model

are higher than 1100 jC, as compared to less than

1000 jC based on plagioclase thermometry. The above

reasoning, combined with the silica vs. P2O5 diagram

(Fig. 3), allows us to pinpoint the silica content in

which apatite saturation was reached, inferring this to

be around 55% for both plutons.

Phosphorus enrichment in monzonitic rocks may

be an effect of the following processes: (a) the

presence of inherited apatite, (b) magma mixing or

assimilation processes with a P-enriched end-member,

(c) magma peraluminosity, and (d) apatite accumula-

tion. Inherited apatite is unlikely to occur if the

underestimated temperatures obtained with the apatite

geothermometer for the most mafic samples are taken

into account. Bearing in mind the cathodolumines-

cence observations, it may be inferred that apatite

grains from the most P-enriched samples do not seem

to display any indications of crustal-affinity (Lopez-

Moro, 2000) and, consequently, mixing or assimila-

tion processes cannot be relevant. Neither does the

enrichment in P appear to be linked to increasing

peraluminosity since the most enriched samples are

not peraluminous (Table 3). Therefore, apatite accu-

mulation is the most likely process to be invoked in

order to explain P-enrichment along the same line of

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–44 33

argument inferred by Eklund et al. (1998) for other

shoshonitic rocks.

On the other hand, clinopyroxene thermometry

(Lindsley, 1983) resulted in underestimated values

( < 600 jC, not included in Table 5). The solidus

temperature derived from the biotite ranges between

940 and 880 jC for the monzogabbro and 890 to

840 jC for the monzonite (Table 5). Temperature

estimates by means of Ab–Kfs and Hbl–Pl geo-

thermometer pairs are always below 657 jC (Table

5), suggesting a fluid interaction under subsolidus

conditions. Nevertheless, taking into account the

isotopic signature of this fluid phase (dDH2O be-

tween � 50 and � 47) (Lopez-Moro, 2000), any

involvement of metamorphic waters should be ruled

out. This suggests that the temperatures obtained

with Kfs–Pl and Hbl–Pl geothermometer pairs

may be simply related to a cooling involving late-

post magmatic fluids.

6.1.1.2. Amphibole-free rocks. For zircon-saturated

rocks, the method of Watson and Harrison (1983)

was used. Liquidus temperatures afforded values

ranging from 865 to 790 jC (Table 5). This range

is not very different from that obtained using the

plagioclase-melt thermometer (875–815 jC for zir-

con-saturated rocks plotted on the descending trend

of silica vs. Zr, Fig. 4), suggesting a lack of inherited

zircon. The apatite saturation model of Harrison and

Watson (1984) was applied but the results obtained

seem to be excessively high (above 1100 jC, Table5). These high temperatures are unlikely to be a

consequence of the magma peraluminosity since the

most peraluminous rocks (Table 3) show the lowest

temperatures.

Solidus temperatures were assessed using the com-

positions of late crystallization phases. Thus, the Ab–

Kfs thermometric pair was applied, also yielding too

low values (between 600 and 502 jC) (Table 5), eventhough the K-feldspar crystals have been heated to

remove unmixing effects. As in the case of the

amphibole-bearing rocks, these results indicate a fluid

interaction under subsolidus conditions, defining an

evolution in the K-feldspar in accordance with an

incoherent solvus (Lopez-Moro et al., 1998). Never-

theless, based on the biotite compositions, the solidus

temperature was estimated between 800 and 765 jC(Table 5). These results are consistent with a possible

solidus not saturated in H2O (Piwinskii and Wyllie,

1968).

6.1.2. Pressure

The lack of thermal aureoles in the country rock

and of suitable mineral pairs makes any assessment of

emplacement pressures difficult. Estimations were

therefore carried out based on amphibole and biotite

(Table 5). The results obtained from the biotite are

underestimated, since water saturation conditions

(PLfPH2O) were assumed, a situation that does

not seem evident here. Starting out from the anni-

te + O2 = magnetite + K-feldspar + H2O equilibrium

(1), and assuming the calibration of the equilibrium

after Wones (1972), PH2O conditions of 410–345

MPa for the monzonite, and 230 MPa for the quartz

monzonite were estimated for the Pereruela Pluton.

Moreover, the estimates based on Al in amphibole are

also underestimated inasmuch as the amphiboles of

the MATD do not show the Al-tschermak but the

pargasite –hastingsite substitution (Lopez-Moro,

2000). On considering the temperatures obtained with

the hornblende-plagioclase thermometer, pressure

conditions after Al in hornblende have been assessed

to be 250–210 MPa for the Pereruela Pluton, and 130

MPa for the Vitigudino Pluton (Table 5), being lower

than those estimated with biotite.

6.1.3. fO2

The presence of magnetite as a primary phase,

together with the abundance of titanite (Wones,

1989), in part due to a process of uralitization,

indicates oxidizing conditions in both magmatic and

late- to post-magmatic stages. Oxygen fugacity was

estimated on the basis of the biotite composition

(method of Wones and Eugster, 1965), as well as by

combining the biotite composition and the whole-rock

Fe+ 3/Fe+ 2 ratio (Burkhard, 1991) (Fig. 7), giving a

range between 10� 13.2 and 10� 15 MPa. Using both

methods, similar results were obtained for the amphi-

bole-bearing rocks, the values being slightly above the

Ni–NiO buffer (Fig. 7). Nevertheless, Burkhard’s

method for amphibole-free rocks results in unrealistic

estimates (temperatures above 1000 jC). The method

of Wones and Eugster (1965) suggests a decrease in

fO2 with evolution, in a similar way to trend II

established by those authors, characteristic of magmas

with decreasing H2O contents.

Fig. 7. Log fO2 vs. T diagrams. (a) After Wones and Eugster (1965). (b) After Burkhard (1991). Symbols as in Fig. 2.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4434

6.1.4. Water content

In contrast to tholeiitic and low-K magmas, calc-

alkaline and shoshonitic magmas are believed to be

enriched in volatile components (Moore et al.,

1998).

Starting from the empirical models based on the

H2O(vapour) = H2O(melt) equilibrium (Moore et al.,

1998), the water content was determined for these

monzonitic magmas. Taking into account the PH2O

calculated from equilibrium (1), and assuming a

maximum pressure of 300 MPa, the following results

were obtained: 5.4% H2O (F 0.5) for the monzonitic

rocks and 5.0% H2O (F 0.5) for the quartz-monzon-

ites. Similar values were estimated using the exper-

imental criteria of Naney (1983) (minimum

conditions of 5.5% H2O for the amphibole-bearing

rocks and a maximum of 5% H2O for the biotite

rocks).

6.2. Petrogenetic modelling

AFC seems to have been the most relevant process,

as shown by the following criteria.

6.2.1. Irregular linear trends

Geochemical evolutionary trends with inflections

in slope may reflect saturation in a mineral phase,

resulting in striking shifts in the residual liquid com-

position. In addition, irregular linear trends indicate a

continual crystallization process, ruling out a major

role for magma mixing.

For instance, an initial increase in Na2O, owing to

the participation of clinopyroxene and biotite as

crystallizing phases, can be observed (Fig. 3). This

is followed by a descending trend once clinopyroxene

has reached the solidus and at the same time the

importance of plagioclase fractionation increased.

F.-J. Lopez-Moro, M. Lopez-Plaz

The behaviour of P represents further evidence of this,

also showing an inflection point (Fig. 3) due to apatite

saturation, as already mentioned.

Concerning the trace elements, zigzag-shaped

trends for Nb, Y and Yb can be explained in terms

of titanite, apatite and zircon saturation, respectively

(Fig. 4).

6.2.2. Corroded plagioclase cores

The anorthite content of corroded plagioclase cores

decreases with increasing whole-rock silica for both

plutons (Fig. 8), which suggests that magma mixing

processes were unlikely to have occurred; as other-

wise the anorthite content of the cores would not have

undergone significant changes. Instead, fractional

crystallization, with the accumulation of plagioclase

cores on the walls of the magma chamber, can be

assumed. Plagioclase cores and residual liquid could

have been joined by convective flow and slumping.

This process seems to have been continuous, in such a

way that differentiation of the residual liquid corre-

lated with the changing composition of the plagioclase

cores. A steeper trend for the Vitigudino Pluton,

starting with more Ca-enriched cores, suggests some-

what different conditions in the magma chamber; e.g.,

water content (Loomis, 1982).

Fig. 8. An (%) in plagioclase vs. whole rock SiO2 (wt.%). Open

symbols: samples from the Vitigudino Pluton; filled symbols:

samples from the Pereruela Pluton. Large circles: corroded cores of

plagioclases of amphibole-bearing rocks; small circles: corroded

cores of amphibole-free rocks. Ellipses: outer zones of plagioclase

crystals.

6.2.3. Modelling of the liquid line of descent

Variation in the Nd, Sr and O isotope data indicates

an open-system evolution for the whole monzonitic

series (Table 3, Fig. 6a and b). The d18O vs. silica

diagram shows two contrasting trends for the two

plutons studied, both having a steeper slope than

fractional crystallization trends (Matsuhisa, 1979)

(Fig. 6b). Similarly, distinct evolution of both masses

is apparent in the eNd320 vs. 1/Nd diagram, with an

almost linear trend being observed for samples from

the Pereruela Pluton (Fig. 6a). This linearity suggests

binary mixing or AFC processes in which the bulk

distribution coefficient of Nd and the rate of assimi-

lation/crystallisation were constant (Albarede, 1995;

Janousek et al., 2000). Ruling out a closed-system

evolution, assimilation-fractional crystallization mod-

elling was carried out for both plutons in order to

explain such contrasting evolutionary trends.

The lower oxygen isotope signature obtained for

the Pereruela Pluton seems to indicate a contaminant

with a lower-d18O signature than that of the Vitigu-

dino Pluton. Among outcropping wallrocks, two

probable candidates were considered: a relatively

low d18O garnet-bearing granulitic orthogneiss

(d18O= + 9.6) for the Pereruela Pluton and a high

d18O metapelitic rock (d18O= + 11.7) for the Vitigu-

dino Pluton; the latter is supported by the appearance

of small metapelitic xenoliths in the intermediate

rocks from this pluton.

Liquid evolution during crystallization and assim-

ilation was modelled using major-element least-

squares modelling on the basis of observed mineral

compositions (Table 6) and whole-rock chemistry of

the respective contaminants (Table 3). The general

mixing equation of Bryan et al. (1969) was used and

implemented using the MacGPP package (Geist et al.,

1989). MnO was omitted as well as P2O5, in stage 1

for the Pereruela Pluton, since apatite is not a frac-

tionating phase. The total iron content was used

(FeOt) instead of the separate FeO and Fe2O3.

Since no appropriate cumulates were found, the

sum of squared residuals (R2) was used to check the

validity of the model with R2V 1.5, the maximum

value for a valid fractional crystallization modelling

(Luhr and Carmichael, 1980). The amount of frac-

tionating zircon was calculated by mass balance

modelling according to the equations of Evans and

Hanson (1993).

a / Lithos 72 (2004) 19–44 35

Table 6

Mineral compositions used in petrogenetic modelling

Stage Mineral SiO2 TiO2 Al2O3 FeOt MnO MgO CaO Na2O K2O P2O5

Pereruela Pluton

1 Cpx 53.48 0.05 0.61 6.97 0.27 14.07 24.14 0.40 0.00 –

Bt 40.45 1.76 15.08 14.85 0.20 17.17 0.00 0.13 10.37 –

Pl 54.79 0.00 28.74 0.11 0.00 0.00 10.76 5.48 0.12 –

Kfs 63.90 0.00 21.56 0.02 0.00 0.00 0.00 1.01 13.51 –

2 Cpx 53.73 0.12 0.64 5.66 0.24 14.87 24.49 0.25 0.01 0.00

Bt 36.13 3.32 16.51 18.99 0.29 14.03 0.16 0.14 10.43 0.00

Spn 31.11 40.53 0.00 0.14 0.05 0.00 27.79 0.38 0.00 0.00

Pl 61.41 0.00 23.86 0.66 0.00 0.12 5.24 8.52 0.20 0.00

Kfs 63.90 0.00 21.56 0.02 0.00 0.00 0.00 1.01 13.51 0.00

Ap 0.00 0.00 0.00 0.22 1.59 0.56 54.78 0.00 0.00 42.84

3 Bt 38.69 3.50 17.27 19.50 0.41 10.36 0.04 0.04 10.2 0.00

Pl 61.95 0.00 23.41 0.04 0.00 0.00 5.60 8.78 0.21 0.00

Ap 0.00 0.00 0.00 0.22 1.59 0.56 54.78 0.00 0.00 42.84

Vitigudino Pluton

1 Cpx 54.12 0.17 0.65 3.40 0.13 16.30 25.18 0.05 0.00 0.00

Bt 38.56 2.57 16.22 17.61 0.22 13.76 0.62 0.17 10.1 0.00

Pl 54.79 0.00 28.74 0.11 0.00 0.00 10.76 5.48 0.12 0.00

Spn 31.11 40.53 0.00 0.14 0.05 0.00 27.79 0.38 0.00 0.00

Ap 0.00 0.00 0.00 0.22 1.59 0.56 54.78 0.00 0.00 42.84

2 Bt 36.13 3.32 16.51 18.99 0.29 14.03 0.16 0.14 10.43 0.00

Pl 54.79 0.00 28.74 0.11 0.00 0.00 10.76 5.48 0.12 0.00

Ap 0.00 0.00 0.00 0.22 1.59 0.56 54.78 0.00 0.00 42.84

c 1 + 2 Cpx 54.12 0.17 0.65 3.40 0.13 16.30 25.18 0.05 0.00 0.00

Bt 40.45 1.76 15.08 14.85 0.20 17.17 0.00 0.13 10.37 0.00

Pl 54.79 0.00 28.74 0.11 0.00 0.00 10.76 5.48 0.12 0.00

Spn 31.11 40.53 0.00 0.14 0.05 0.00 27.79 0.38 0.00 0.00

Mag 0.00 0.29 0.23 99.49 0.00 0.00 0.00 0.00 0.00 0.00

Ap 0.00 0.00 0.00 0.22 1.59 0.56 54.78 0.00 0.00 42.84

FeOt = total iron. Mineral abbreviations after Kretz (1983). Analysis recalculated to 100%. P2O5 content is not considered in stage 1 for the

Pereruela Pluton since apatite is not the fractionating phase.

Data after Lopez-Moro (2000).

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4436

Three major steps were considered for the Perer-

uela Pluton, corresponding to the main rock types, and

the same were considered for the Vitigudino Pluton,

except that here there are no equivalent monzogranites

(Table 7).

First, major-element based modelling was per-

formed in order to establish the cumulate minerals

involved in all the stages, as well as the residual liquid

(F), contaminant (Cont), and assimilation/crystalliza-

tion rate (r) (Table 7). Following this, trace-element

AFC modelling was carried out using equations of

DePaolo (1981) (Table 8) and isotopic data were also

included in order to check the validity of the AFC

process, shown graphically in Fig. 9. Trace-element

partition coefficients used for AFC modelling are

given in Table 9. Different partition coefficients for

the same element and mineral have been used in each

evolution stage, depending on the silica content of the

melt.

Concerning the results, it is worth emphasising

the low assimilation/crystallization rate. This ranges

between 0.16 and 0.20 for the Pereruela Pluton, and

it is almost constant for the Vitigudino Pluton (0.24–

0.25) (Table 7). The percentages of assimilated

material are similarly low, ranging from 7% to

12% for the Pereruela Pluton, and from 5% to 26%

for the Vitigudino Pluton. Such low rates make

assimilation plausible since at upper/middle crustal

levels higher rates would not be expected (DePaolo,

1981).

Table 7

Multi-stage major-element based least-squares modelling of liquid line of descent for the Pereruela Pluton and the Vitigudino Pluton

Stages Cumulate minerals (%) Cont R2 F f r

Cpx Pl Bt Fks Spn Ap Mag

Pereruela Pluton

(1) Mzgb (54.22)–Mz (58.16) 3.14 56.31 31.19 9.35 – – – 12.05 0.29 67 67 0.20

(2) Mz (58.16)–Qmz (63.85) 6.42 49.34 34.42 5.28 0.46 3.67 – 11.70 0.36 64 31 0.17

(3) Qmz (63.85)–Mzgr (70.68) – 48.64 47.71 – – 3.53 – 7.39 0.29 72 3 0.16

Vitigudino Pluton

(1) Mzgb (52.78)–Qmz (62.54) 11.74 33.59 37.84 – 0.47 1.28 – 25.97 0.17 57 57 0.25

(2) Qmz (62.54)–Qmz (67.30) – 48.63 46.00 – – 5.38 – 5.07 0.88 86 43 0.24

Mzgb (52.78)–Qmz (66.75)a 10.91 42.32 39.11 – 2.71 1.87 2.84 33.93 0.08 49 – 0.25

Cont: percentage of wallrock assimilation by using a garnet-bearing granulitic orthogneiss (sample POR 86) for Pereruela Pluton and a

metapelite (sample DT-70) for Vitigudino Pluton. R2: sum of the squared residuals. F: percentage of residual liquid. f: percentage of residual

liquid with respect to stage 1. r: assimilation/crystallization rate. Numbers in brackets are whole-rock silica contents (wt.%). Mineral

abbreviations after Kretz (1983); rock abbreviations as in Table 3.a This nearly-bulk evolution based modelling has been carried out in order to estimate the percentages of cumulate minerals and the

assimilation/crystallization rate used in the isotopic modelling.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–44 37

The calculated proportions of residual liquids sug-

gest a limited production of felsic magmas, particu-

larly for the Pereruela Pluton, where a final percentage

of residual liquid was estimated to be around 3%, in

agreement with the scarcity of monzogranites.

Table 8

Trace-element modelling for the Pereruela Pluton

Element Stage 1 Stage 2

Mzgb (54.22)–Mz (58.16) Mz (58.16)–

Measured AFC Measured

La 176 192 57.13

Ce 300 332 119.4

Sm 16.00 21.35 8.24

Nd 121 141 47.02

Eu 3.60 2.77 2.0

Yb 2.00 3.02 1.32

Lu 0.27 0.41 0.21

Rb 160 149 232.7

Sr n.d. n.d. 697

Ba 2100 2234 1623

Cr 68 62 127

Y n.d. n.d. 16.1

Nb 22 10 12.43

CC 0.9998

PSD 15.30

AFC: assimilation–crystallization according to equations of DePaolo (198

of fractionating allanite at stage 2 was estimated to be 0.32; whereas the

respectively. Silica weight percentages of modelled samples are shown

calculated values in the AFC process. PSD: average standard deviation of

The cumulate assemblages of major minerals are

clinopyroxene–plagioclase–biotite–Kfs in the two

first stages of the Pereruela Pluton and only plagio-

clase–biotite in the third stage, the latter being similar

to the latest stage of the Vitigudino Pluton. An early

Stage 3

Qmz (63.85) Qmz (63.85)–Mzgr (70.68)

AFC Measured AFC

60.38 25.29 16.51

117.57 49.06 41.40

8.19 2.75 2.6

44.78 17.59 14.58

1.9 1.84 1.01

1.40 0.51 0.51

0.17 0.08 0.07

193.4 180.0 170.2

716 670 679

1254 1957 2044

53 12 4

19.9 5.1 8.4

12.74 4.6 4.7

0.9913 0.9999

27.91 7.52

1) with garnet-bearing granulitic gneiss as contaminant. Percentages

fractionating values for zircon are 0.09 and 0.12 at stages 2 and 3,

in brackets. CC: Correlation coefficients between measured and

estimated compositions.

Fig. 9. AFC modelling for selected samples from the Pereruela Pluton (a) and Vitigudino Pluton (b), showing diagrams of combined Nd, O and

Sr isotopic ratios obtained for the elements indicated. In figure (a), calculations were obtained by modelling Sr, Y and Lu data, assuming

averaged r = 0.165 in stages 2 and 3, as well as a garnet-bearing granulitic orthogneiss (POR-86) from the Tormes Dome as a contaminant. By

contrast, in figure (b) calculations were based on Yb data, assuming r= 0.25 and a metapelite as contaminant (DT-70) from the Tormes Dome. In

both figures, error bars (2r) are shown for the three isotopic systems. Labelled ticks indicate decreasing melt fraction ( F). Averaged bulk

distribution coefficients (D) were estimated according to normative weight fraction of crystallizing mineral phases, calculated by major element

mass-balance modelling in stages 2 and 3 with the same contaminant. Symbols as in Fig. 2.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4438

crystallization of Kfs has been already reported for the

monzodiorites of appinitic suites (Weiss and Troll,

1989), and also for the Svecofennian monzonitic

series with similar silica ranges (Eklund et al., 1998).

With respect to accessory minerals, titanite appears

in stage 2 for the Pereruela Pluton, whereas apatite

does so in stages 2 and 3; i.e., after saturation has been

reached. Analogously, the AFC modelling for the

Table 9

Trace element partition coefficients used for AFC modelling

Element Aln Ap Pl Bt Cpx Fk Spn Zrn Mag

La 820 14.5; 46.1 0.32 0.32; 3.18 0.6; 1.1 0.07; 0.12 2; 4 16.9 0.05

Ce 635 21.1; 34.7;41.6 0.27 0.32; 2.80 0.51; 1.83 0.037; 0.11 53.3 16.75 0.05

Sm 205 46; 62.8 0.013; 0.13 0.058; 1.55 0.9; 5.23 0.018; 0.11 10; 21 4.94; 14.4 0.05

Eu 81; 111 27.3; 30.4 2; 2.11 0.24; 0.86 1.56; 4.10 1.13; 4 6.3 3.3; 16 0.05

Yb 30.8 23.9; 60 0.049 0.44; 0.53 1.58; 6.36 0.012; 0.015 11 527 0.25

Lu 7.7; 33 13.8; 20.2; 60 0.046 0.33; 0.613 1.54; 5.93 0.006; 0.015 6; 10 641 0.05

Rb – 0.01 0.041 2.2; 3.2 0.032 0.49; 0.34 – – 0.01

Sr – 5 2.84; 4; 4.4 0.12; 0.36 0.516 3.76; 3.87 0.001; 06 – 0.01

Ba – 0.1 0.36 0.57*; 5.36 0.001; 0.131 3.67*; 5.37; 5.9* 0.001 – 0.01

Cr 380 – 0.2 2.6*; 4.89*; 17 1.07 – – 189.5 10

Y – 40 0.13 0.03; 2 3.1 0.1 – – 2

Nb – 0.1 0.06 6.4 0.75; 0.8 – 6.3 – 2.5

Data from Rollinson (1993), Lopez-Ruiz and Cebria (1990) and references therein; data with asterisk from Lopez-Moro (2000); mineral

abbreviations after Kretz (1983).

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–44 39

Vitigudino Pluton indicates titanite and apatite in stage

1, and only apatite in stage 2. Other significant acces-

sory phases are allanite and zircon, which appear to

have a strong influence on LREE and HREE chemis-

try, respectively. Fig. 10 illustrates how the addition of

minor amounts of allanite and zircon can improve the

fit of the model, these mineral phases having partici-

pated after reaching their saturation. Concerning trace

Fig. 10. Comparative results of AFC modelling for the Pereruela

Pluton in stage 2 by using only major elements in one case,

involving apatite as fractionating phase, and major plus trace

elements in the other case. Note strong influence on LREE and

HREE owing to the fractionation of allanite (0.32%) and zircon

(0.09%), respectively.

elements, in general AFC modelling shows a good fit

between the measured and calculated values, mainly

for stage 3 (Table 8). Likewise, the validity of the

process is confirmed by the similar values of calculated

bulk partition coefficients for Nd at stages 2 and 3 (3.1

and 3.0, respectively), accounting for the almost rec-

tilinear trend in eNd320 vs. 1/Nd diagram (Fig. 6a).

Nevertheless, the discrepancy in Eu at this stage should

be noted; it can be explained as follows. Increasing

oxidation conditions from POR14 to POR 14B favour

a low Eu+ 2 partition coefficient. This fact together with

a strong plagioclase and apatite fractionation accounts

for a high Eu-positive anomaly for latest K-rich and

limited residual liquids (POR 14B). A slightly higher

magnesium number in POR14B, not only in whole

rock but also in biotite (Tables 2 and 3), is consistent

with the higher oxidation state.

AFC modelling using isotopic data broadly con-

firms the validity of the model for both plutons (Fig.

9), except for the Sr isotopic data in stage 3 for the

Pereruela Pluton. The high Sr content in sample POR-

14B and in the contaminant itself can account for this

decoupling, buffering the isotopic composition during

contamination. Also, the discrepancy can be explained

in terms of different rates of interdiffusion (Lesher,

1990). The best fits regarding the evolutionary trends

for the Pereruela Pluton are obtained with Y and Lu,

whereas if the fraction of residual liquid is considered

the best fits with respect to major-element based

modelling are obtained with Y in stage 3 and Lu in

stage 2. On the other hand, a reasonable fit is obtained

Fig. 11. (Fe +Mg)–Ca–K diagram plotting experimental melts

using a hybrid rock (mixture) as starting composition whose end-

member proportions are 8:1:1 (peridotite, pelite and basalt) (Lopez

et al., 2001). Temperature and pressure conditions are also shown.

Monzogabbros of MATD plot near experimental melts at 1200 jCand 140 MPa.

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4440

with Yb and all the isotopic systems considered for

Vitigudino Pluton.

6.3. Origin of source geochemistry and geotectonic

setting

The characteristic enrichment of the MATD may

have resulted from a previously enriched source as:

(a) The isotope data indicate that amphibole-bearing

rocks are the most enriched in Sr, Ba, P and REE

for both plutons studied, and at the same time they

are the least contaminated at upper/middle crustal

levels.

(b) The above argument can be extended to F and

H2O contents since these are also indicators of

source enrichment in agreement with data of

Eklund et al. (1998). However, carbonate meta-

somatism does not seem to have played a major

role owing to low F/H2O ratio and high SiO2

contents for subsequently generated melts repre-

sented by the monzogabbroic rocks of MATD.

(c) The high total content of REE with subparallel

patterns for the monzogabbroic rocks may be an

indicator of source enrichment (Evans and

Hanson, 1997).

Accepting the affinity of the MATD to the SH-type

granitoids, the involvement of metasediments in the

source could be invoked to account for the enriched

source (Jiang et al., 2002). Using a mixture of perido-

tites, amphibolitic rocks and metasediments, experi-

mental modelling has been performed to explain the

composition of some shoshonitic rocks from NW

Iberian Massif (Lopez et al., 2001). According to their

experimental results, the variable contribution of a

metasedimentary component may have caused a cer-

tain obliquity of REE patterns for primary melts (Lopez

et al., 2001). The least differentiated rocks from the

plutons studied here show small differences in REE

contents and patterns, suggesting similar hybrid pro-

tholiths. On the other hand, the experiments of Lopez et

al. (2001) allow experimental melts to be compared

with MATD rocks by plotting on the (Fe +Mg)–Ca–K

diagram (Fig. 11). The temperature derived from

experiments (1200 jC under high pressure conditions)

appears to be slightly higher than the liquidus temper-

ature estimated in this work (1090 jCF 60 jC under

low pressure). Likewise, the assumed metasedimentary

and amphibolitic components of this experiment may

account for the high-water contents estimated for

MATD magmas in this work (c 5.5%).

The occurrence of K-rich plutonic rocks in the

axial part of the NW Iberian Massif (inset of Fig. 1)—

i.e., nearby allochthonous complexes—is an observa-

tion that should not be overlooked. Allochthonous

complexes consist of a pile of units, some of them

showing ophiolitic affinities (Arenas et al., 1986), and

are hence possible candidates as hybrid protholihs of

K-rich plutonic rocks. Late-orogenic slab break-off

could have provided heat for partial melting of such

hybrid protholiths close to the crust/mantle boundary,

or even at crustal levels. Post-collisional extensional

events enhanced the rise of isotherms, further contrib-

uting to partial melting of such deep hybrid protho-

liths. South of the Tormes Dome—the Central System

batholith—there are no K-rich plutonic rocks (Fig. 1),

and late Variscan calc-alkaline granites and alkaline

magmas (camptonites) appear instead, suggesting a

different scenario, with the probable contribution of a

subcratonic mantle (Bea et al., 1999). In any case, this

type of compositional variability can also be ex-

plained in terms of mantle heterogeneity (see discus-

sion in Janousek et al., 2000).

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–44 41

7. Conclusions

K-rich plutonic series are well represented in the

Tormes Dome area by two plutons (the Pereruela and

Vitigudino plutons), both forming part of two biotite

granite belts. On the basis of its relatively extreme

K-enrichment, the monzonitic association is well

defined, with monzogabbros, monzonites, quartz

monzonites and scarce monzogranites. K and LREE

enrichment, and to a lesser extent in Sr, Ba, P and F,

is the main geochemical characteristic of MATD

rocks. Additionally, two other salient features of

MATD magmas are noteworthy: first, their strongly

hydrated nature and, second, their oxidized character,

close to the Ni–NiO buffer, showing a loss of water

from 5.5% to 5.0%. The overall petrographic and

geochemical characteristics allow MATD rocks to be

ascribed to the so-called SH-type granitoids (shosh-

onitic granitoids).

The minimum emplacement pressure is estimated

at 410–345 MPa for amphibole-bearing rocks. Their

solidus, not saturated in volatiles, was reached, at

temperatures between 940 and 840 jC, whereas the

solidus temperature for the less mafic members ranges

from 800 to 765 jC.The cathodoluminescence study seems to suggest a

lack of restitic apatites. It also indicates that mixing

process cannot account for the P enrichment in

monzonitic rocks.

Assimilation/fractional crystallisation (AFC) mod-

elling was performed, using major- and trace-ele-

ment as well as Sr, Nd and O isotope data, to

explain the evolution from monzogabbro to mon-

zogranitic members. Different contaminants appear

to have been involved: a granulitic gneiss for the

Pereruela Pluton and a metapelite for the Vitigudino

Pluton, although similar low assimilation/crystalliza-

tion rates are inferred for both of them (r = 0.16–

0.20 for Pereruela and r = 0.24–0.25 for Vitigudino).

The monzonitic liquid line of descent was addressed

starting from monzogabbroic magmas at a liquidus

temperature range of 1092–1048 jC to monzogra-

nitic magmas at 864–790 jC. Apatite, allanite and,

to a lesser extent, titanite, zircon and magnetite were

the fractionating phases controlling the trace-element

variations.

The least evolved members appear to have under-

gone a low degree of contamination at upper/middle

crustal levels, but they are the most enriched in Ba, Sr

and LREE, strongly suggesting a source contamina-

tion. The experiments are consistent with a hybrid

(involving peridotites, amphibolites and metasedi-

ments) protholith of these MATD rocks, suggesting

the recycling of subducted components.

Acknowledgements

The authors wish to thank Dr. Janousek for his

constructive review and suggestions; to Prof. A.

Castro for previous discussion. N. Skinner and J.J.

Lopez reviewed the English text. They are also

indebted to I. Armenteros for making a cathodolumi-

nescence microscope available, and to Prof. G. Moore

for software to estimate water contents in magmas.

Appendix A. Analytical techniques

Twelve fresh samples were selected for the analysis

of whole-rock, five of them for electron microprobe.

Minerals were analysed with a Camebax micro-

probe at the University of Oviedo, Spain. Operating

conditions were 15 kV and 20 nA, with a counting

time of 20 s. Major- (except for FeO, obtained by

titration) and trace-elements of biotite separates were

analysed by ICP mass spectroscopy, except Rb, which

was analysed by atomic absorption spectroscopy.

Whole-rock analyses were carried out at the ‘‘Ser-

vicio General de Analisis Quımico Aplicado de la

Universidad de Salamanca’’ for major elements; at the

‘‘ACTLABS’’ in Canada, and the ‘‘Service d’Analy-

ses des Roches du CNRS’’ in Nancy (France) for trace

elements including REE. Major element analyses

were carried out by inductively coupled plasma-atom-

ic emission spectroscopy (ICP-AES), whereas trace

elements and REE were determined by ICP-MS.

Five samples were also selected for the study of

Rb–Sr and Sm–Nd isotopes. They were analysed at

the ‘‘Laboratorio de Geocronologıa de la Universidad

Complutense de Madrid’’. Sr was run on Ta single

filaments, whereas Sm and Nd were run on Ta–Re–

Ta triple filaments. All of them were determined using

VG Sector 54 mass spectrometer thermal ionization,

with five Faraday cups in multidynamic mode, except

for Sm, which was measured in static mode. The Sr

F.-J. Lopez-Moro, M. Lopez-Plaza / Lithos 72 (2004) 19–4442

measurements were corrected for possible interference

by 87Rb and normalized to 88Sr/86Sr = 0.1194. The

measurements of Nd were corrected for interference

by 142Ce and 144Sm and normalized to 146Nd/144Nd =

0.7219. During this study, the NBS-987 standard gave

an average value (n = 12) for 87Sr/86Sr of 0.710271F0.00002, with an average value (n = 7) for 143Nd/144Nd of 0.511806F 0.000005, which is equivalent

to those obtained by this laboratory in a period of 24

months: 87Sr/86Sr = 0.710261F 0.00002 (2r, n = 112)and 143Nd/144Nd = 0.511809F 0.000008 (2r, n = 41).eNd values were calculated using CHUR parameters:143Nd/144Nd = 0.512638 and 147Sm/144Nd = 0.1967

(DePaolo and Wasserburg, 1976). Decay constants

used are: 1.42� 10� 11 year� 1 (87Rb) (Steiger and

Jager, 1977) and 6.54� 10� 12 year� 1 (147Sm) (Lug-

mair and Marti, 1978). The 2 S.D. error on eNdcalculations is F 0.4. The Rb, Sr, Sm and Nd con-

centrations used to age-correct the isotopic data were

obtained by ICP-MS.

Oxygen isotope analyses were performed at the

University of Salamanca (Servicio General de Anali-

sis de Isotopos Estables). Oxygen extraction for

isotopic analysis followed of Clayton and Mayeda

(1963), but employed a loading technique similar to

that described by Friedman and Gleason (1973) and

CIF3 as reagent (Borthwick and Harmon, 1982).

About 10 mg of finely powdered sample was loaded

into nickel vessels and reacted for 15 h at 690 jC. Theoxygen released was converted to CO2 by means of a

carbon-rod heated by a platinum wire. Isotope ratios

were determined on a VG SIRA-II mass spectrometer.

Results are reported in the usual notation, as d18Ovalues relative to the V-SMOW reference standard.

The typical overall reproducibility of duplicate runs

was within F 0.2x(1r).

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