Origin and petrogenetic implications of tourmaline-rich rocks in …hera.ugr.es/doi/14976985.pdf ·...

Origin and petrogenetic implications of tourmaline-rich rocks in

the Sierra Nevada (Betic Cordillera, southeastern Spain)

J. Torres-Ruiza,*, A. Pesquerab, P.P. Gil-Crespob, N. Velillaa

aDepartamento de Mineralogıa y Petrologıa, Facultad de Ciencias, Universidad de Granada,

Campus Fuentenueva, s/n, E-18002 Granada, SpainbDepartamento de Mineralogıa y Petrologıa, Universidad del Paıs Vasco 644, 48080 Bilbao, Spain

Received 13 September 2001; accepted 11 September 2002

Abstract

Tourmaline-rich rocks (up to 60% tourmaline) associated with low–medium grade metamorphic assemblages occur in the

Sierra Nevada area (Betic Cordillera, southeastern Spain). Tourmaline appears in a variety of forms: (1) stratiform tourmalinites;

(2) quartz– tourmaline nodules; (3) porphyroclasts in felsic orthogneisses; and (4) disseminations in psammopelitic meta-

sediments and gneisses. Tourmaline within these lithologic groups exhibits textural and chemical variations that reflect complex

premetamorphic growth under open-system conditions, and subsequent changes due to Alpine regional metamorphism.

Microprobe analyses of the tourmalines reveal a wide compositional variation between schorl and dravite end members with

variable contents of X-site vacancies (av. 0.084–0.225 apfu), Ca (av. 0.095–0.269 apfu), and excess of Al (up to 6.588 apfu)

compared with the theoretical value of 6 in ideal schorl and dravite. The amount of Ca may be significant in porphyroclasts from

the gneisses. Fe/(Fe +Mg) ratios for tourmalines in tourmalinites, metasediments, and gneisses range from 0.34 to 0.95, 0.16 to

0.92, and 0.28 to 0.97, respectively. Na/(Na +Ca) ratios are also variable, mostly ranging from 0.5 to 0.9. Many of the tourmalines

have complex chemical and colour zoning patterns, including significant fluctuations in Al, Fe, Mg, Na, Ca, Ti, and F.

Based on petrographic and chemical data, three generations of tourmaline have been established. The first generation

corresponds to magmatic–postmagmatic tourmaline that is represented by tourmaline porphyroclasts within the orthogneisses.

The second generation of tourmaline formed during tourmalinization of psammopelitic rocks giving rise to tourmalinites. The

third generation of tourmaline is represented by cellular textures, pale reaction rims and overgrowths developed during the

Alpine regional metamorphic overprint.

D 2002 Elsevier Science B.V. All rights reserved.

Keywords: Tourmalinites; Mineral chemistry; Metasomatism; Metamorphism; Sierra Nevada; Spain

1. Introduction

It is clear from many studies that boron can be

mobilized by metamorphic processes (Engel and

Engel, 1960; Harder, 1975a; Moran et al., 1992;

Bebout et al., 1993; Henry and Dutrow, 1996; Leeman

and Sisson, 1996; Sperlich et al., 1996), but the boron

budget during metamorphism is still a matter of debate.

In sedimentary and metamorphic rocks, boron abun-

dances are determined by interactions among solu-

tions, phyllosilicates, and tourmaline (Shaw, 1996).

0009-2541/02/$ - see front matter D 2002 Elsevier Science B.V. All rights reserved.

doi:10.1016/S0009-2541(02)00357-1

* Corresponding author. Fax: +34-958-243368.

E-mail address: [email protected] (J. Torres-Ruiz).

www.elsevier.com/locate/chemgeo

Chemical Geology 197 (2003) 55–86

The behaviour and geochemical distribution of boron

hinge on several factors such as protolith nature,

thermal history, activity of fluids, crystal chemical

relations, and stability of B-bearing minerals (Leeman

and Sisson, 1996). Much of the original boron in a

protolith may be released by dewatering and dehydra-

tion reactions with increasing metamorphism. In other

cases, however, significant boron enrichment is present

in some high-grade metamorphic rocks, raising ques-

tions about boron sources and retention of boron under

high-grade conditions. Undoubtedly, the degree of

boron retention will depend on availability of boron-

bearing minerals and their respective stability fields.

Tourmaline can play an important role because it is the

most widespread boron mineral and has a large P–T

stability. The nature of the fluids, nevertheless, may

have a critical effect on the tourmaline stability (Mor-

gan and London, 1989). It has long been recognized

that tourmaline may provide valuable information

during petrologic studies (Henry and Guidotti, 1985).

Tourmaline is an ideal monitor of the local boron flux

in metamorphic environments and can serve as a sink

or source of boron (Henry and Dutrow, 1996).

Tourmaline-rich rocks associated with leucocratic

gneisses and garnet-bearing pelitic schists occur in

the Mulhacen Nappe Complex of the Sierra Nevada

(Betic Cordillera, Spain; Fig. 1). They form part of a

complex geological setting that involves polyphase

deformation and low to medium-grade regional meta-

morphism during the Alpine orogeny. Although the

occurrence of tourmaline-rich rocks in this area has

been noted previously (e.g., Puga, 1971; Nieto, 1996),

these studies generally lack a comprehensive model to

explain their chemical variations, textures and the

significance. Detailed field and petrographic observa-

tions suggest a complex sequence of events with early

boron metasomatism and multistage tourmaline for-

mation. Tourmaline is able to preserve primary char-

Fig. 1. Simplified geologic map of the Sierra Nevada area showing the tourmaline-rich rocks occurrence (modified from IGME, 1980), and

generalized lithostratigraphy column of Nevado–Filabride Complex (modified from Gomez-Pugnaire et al., 2000), with detail of the

tourmaline-rich rock outcrops.

J. Torres-Ruiz et al. / Chemical Geology 197 (2003) 55–8656

acteristics through several geological processes due to

its mechanical and chemical stability. However,

despite its refractory character, petrographic data and

microprobe data indicate that deformation and

regional metamorphism caused significant effects on

tourmaline in Sierra Nevada. Destabilization pro-

cesses of tourmaline may have important consequen-

ces on the redistribution and boron content during

prograde metamorphism and anatexis.

In this paper, we present geological, mineralogical

and geochemical data on tourmaline-rich rocks from

the Sierra Nevada with the intention of establishing:

(i) the chronology of tourmaline formation processes;

(ii) the repercussion of the different mineralogenetic

processes (in relation with the igneous activity, meta-

somatism and regional metamorphism) on its textural

and chemical characteristics; (iii) the nature of the

tourmalinites protholith; and (iv) the possible source

of boron. The study of tourmaline may give some

insights into the geology of a region, and in turn may

provide a valuable information on the processes that

control the recycling of boron in the crust.

2. Geological setting

The Betic Cordillera in southeast Spain (Fig. 1)

represents the westernmost mountain lineament of the

Mediterranean Alpine chains, and its physiography is

the result of coeval contractional and extensional

processes. It has long been divided into an unmeta-

morphosed external domain and a dominantly meta-

morphosed internal domain (Egeler and Simon, 1969).

The internal domain, termed the Betic zone or Alboran

Domain, includes three tectonically superimposed

complexes: the Nevado–Filabride, Alpujarride, and

Malaguide complexes. The uppermost Nevado–

Filabride Complex contains a series of stacked thrusts

sheets including a Precambrian to Paleozoic basement

of graphite-bearing metapelite with intercalations of

graphite-bearing quartzite, and an overlying Permo-

Triassic sequence of metapelite, quartzite and marble

with local discontinuous sheets of orthogneiss in the

upper part of the complex. Small and discontinuous

intrusive bodies and subordinate lava flows of mafic

and tectonically interspersed serpentinized ultramafic

igneous rocks also occur in the Permo-Triassic cover

sequence.

The tectono-stratigraphic arrangement of the Ne-

vado–Filabride Complex is controversial. In most of

the regions two groups of tectonic units are distin-

guished: (i) the Veleta nappe (lower); and (ii) the

Mulhacen nappe (upper) (Puga et al., 1974; Dıaz de

Federico, 1980). In turn, the Mulhacen complex has

been subdivided into three tectonic units (Puga et al.,

1974; Dıaz de Federico, 1980). Despite the differ-

entiation in tectonic units, some workers claim that

the lithostratigraphic column of the Mulhacen nappe

can be reconstructed (De Jong and Bakker, 1991;

Jabaloy, 1993; Gomez-Pugnaire et al., 2000). Accord-

ing to Gomez-Pugnaire et al. (2000), the Mulhacen

nappe includes a basement of graphite-rich micaschist

and quartzite and a cover composed of three forma-

tions (Fig. 1): (1) Tahal schist formation, comprising

alternating quartzite and pale micaschist with garnet,

chloritoid, and/or amphibole, (2) metaevaporite for-

mation, consisting of discontinuous levels of scapolite-

bearing marble and metapelite; (3) marble and calc-

schist formation, composed of calcitic and dolomitic

marble with intercalations of calc-schist, micaschist

and quartz–feldspar gneiss. Locally, the basement and

cover sequences are separated by discontinuous and

scarce outcrops of metaconglomerate.

Within the Mulhacen nappe, two types of orthog-

neiss have been distinguished (Nieto, 1996): (1)

metagranites and orthogneisses of plutonic character

that occur intercalated between basement lithologies

of the Mulhacen nappe; (2) stratiform orthogneisses of

probably volcanic character associated with the

Permo-Triassic cover of the Mulhacen nappe. The

first type has been considered tectonically to be

syncollisional granites, representing granitic melts

generated by partial melting of Precambrian metase-

dimentary rocks recycled during the Hercynian orog-

eny. In contrast, the second type of orthogneiss

appears to be consistent with an intraplate environ-

ment (Nieto, 1996).

The Nevado–Filabride rocks record the effects of a

polyphase Alpine metamorphism, but pre-Alpine me-

tamorphic effects are also recognized in the basement.

In the Mulhacen Complex, the Alpine metamor-

phism evolves from high pressure and low-temper-

ature conditions, producing eclogites in metabasites

and blueschist-facies assemblages in metapelites, to

intermediate temperature and pressure conditions with

development of almandine-amphibolite facies assemb-

J. Torres-Ruiz et al. / Chemical Geology 197 (2003) 55–86 57

lages that were subsequently retrograded to the greens-

chist facies (Vissers, 1981; Gomez-Pugnaire and Fer-

nandez Soler, 1987; Bakker et al., 1989; Puga et al.,

2002). Although both phases developed a foliation, the

early foliation is significantly obliterated by the main

foliation that was generated during the almandine-

amphibolite facies event. Nevertheless, in many areas,

the main foliation is masked by further later deforma-

tion. These later deformational events are believed to

have formed during decompression and exhumation of

the Nevado–Filabride Complex. The extensional

deformation, subsequent to the structures related to

the thrusting, produced significant shear zones that are

accompanied by mylonitic fabrics in the greenschist

facies rocks. Later, brittle deformation gave to cata-

clasites, faults, and tectonic breccias (Galindo-Zaldı-

var et al., 1989; Martınez-Martınez and Azanon,

1997).

3. Field relations and petrography of

tourmaline-rich rocks

Tourmaline is widespread in leucocratic gneisses

and interlayered schists of the Permo-Triassic cover

from Mulhacen nappe. Gneisses appear as layers of

variable thickness ranging from several centimeters to

20 m (Fig. 1). The thickness of schistous layers varies

from a few millimeters to several decimeters.

Tourmaline occurs as: (1) very fine to fine-grained

crystals in stratiform tourmalinites (Fig. 2A,D); (2)

porphyroclasts (normally < 1 cm), particularly in

gneisses, accompanied by feldspar porphyroclasts

and quartz–feldspar lenses that are wrapped by my-

lonitic foliation (Fig. 2B,C); (3) generally submilli-

metric crystals disseminated throughout different

intercalated levels of orthogneiss and schist (Fig.

2C); (4) oval nodules (up to 25 cm in length)

composed of quartz–tourmaline, which appear within

the schists and gneisses; (5) clasts in quartz veins that

were derived from the host rocks by mechanical

remobilization (Fig. 2D). Stratiform tourmalinites

are a few millimeters to 50 cm thick and in outcrops

may be traced along strike for tens of meters. They

form well-defined bands, parallel to the mylonitic

foliation, that alternate with tourmaline-bearing

schists and gneisses. S–C structures, shear bands,

transposition, lenticular and tadpole forms are char-

acteristic features resulting from the mylonitization

(Fig. 2D).

The petrographic characteristics of tourmaline vary

depending on the host rocks in which they occur.

Three groups are distinguished (Table 1).

3.1. Tourmalinites

Tourmaline and quartz are the dominant minerals,

with the former constituting more than 20% to 60%

by volume. Minor amounts of muscovite are common.

Accessory minerals include biotite, epidote, titanite,

and garnet. Tourmaline and quartz are fine to very

fine-grained ( < 250 Am in length), and in thin sections

cut normal to the lineation they display a granoblastic

texture with interlobate to polygonal grain boundaries.

Tourmalinites show an oriented fabric defined by the

elongate shape of most the tourmaline grains, subpar-

allel to the compositional layering. Some tourmalinite

bands include alternating tourmaline-rich laminae

with very thin quartz-rich laminae, the latter with

< 60-Am crystals and with a strongly oriented fabric.

Tourmaline crystals may be transected by quartz-rich

microfractures. Tourmaline porphyroclasts and spher-

ical to oval spots of tourmalineF epidote ( < 500 Am)

may occur in some tourmalinite laminae.

In thin section, tourmaline crystals display varia-

ble colours: pale to very dark green, bluish green,

blue, orange and yellowish brown; but green colours

predominate. Many tourmalines show distinct zones

of dark green interiors and pale rims. The interior

zones commonly include very fine-scale growth

lamellae that typically are truncated by pale rims or

by microfractures or the matrix (Fig. 3A,B,D). Sector

zoning is observed in some crystals (Fig. 3C). Typ-

ically, the interior zones are embayed by pale dis-

cordant rims like those reported by Slack and Coad

(1989). One of the more characteristic features is the

development of patchy zoning and cellular morphol-

ogies, which may be rimmed by noncellular tourma-

line (Fig. 3D,E,F). Some crystals display interior

zones composed of two or more tiny grains that are

overgrown by pale outer zones. This may account for

coalescence of small grains, or crystallization centers,

mantled by outer zones. Inclusions are not very

abundant; quartz is the most common, forming irreg-

ularly shaped inclusions in the interior zones of the

crystals.


Fig. 2. (A) Stratiform tourmalinites (dark-coloured lithologies) alternating with gneisses (light-grey) and metasediments. (B) Mylonitic gneiss

with quartz– feldspar layers (light) and tourmaline porphyropclasts (black), in a section normal to the foliation. (C) Hand specimen showing an

alternation of tourmaline-bearing psammopelitic and gneissic layers, in a section normal to the foliation. Feldspar and quartz lenses with

tourmaline porphyroclasts define the mylonitic foliation that is deflected around the porphyroclasts. (D) Fine-grained layered mylonite with

tourmalinite lenses and ‘‘tourmalinite fish’’ in a section normal to the foliation. The layering is disturbed by asymmetric isoclinal folds. Note the

low-angle vein cutting the mylonitic foliation.


3.2. Metasediments

These rocks contain variable concentrations of

tourmaline, normally less than 9% by volume. Coex-

isting major minerals are garnet, mica (muscovi-

teF biotite), and quartz. Accessory minerals include

titanite, rutile, and zircon. Tourmaline occurs nor-

mally as: (i) fine to very fine-grained ( < 200 Am in

length), subhedral to anhedral crystals disseminated

within a foliated quartz-micaceous matrix; and (ii)

minor fine-grained porphyroclasts ( < 800 Am) wrap-

ped by the mylonitic fabric. Polycrystalline porphyr-

oclasts of tourmaline and quartz ( < 5 mm) are locally

observed. In these porphyroclasts, tourmaline grains

display fine-scale growth lamellae, which are concen-

trically developed about the c-axis. However, like the

tourmaline porphyroclasts that commonly were bro-

ken during mylonitization with fragments more or less

separated, the growth lamellae may appear abruptly

truncated. Subhedral to subrounded porphyroclasts of

garnet ( < 1 mm) are abundant, and similar to the

tourmaline porphyroclasts, they are also wrapped by

the foliation. Strain shadows and strain caps enriched

in quartz and micas, respectively, are commonly

associated with the porphyroclasts.

Fine-grained disseminated tourmaline crystals,

which are oriented in the foliation, commonly display

a discontinuous optical zoning characterized by dark

green cores and pale green rims. Green rims on pale

cores can also be observed. Both rims and cores show

a variable habit, from euhedral to anhedral shapes.

Some tourmaline crystals exhibit a distinct optical

zoning that involves Cr-bearing orange cores and

green to pale green overgrowths (Fig. 3G). The over-

growths exhibit petrographic characteristics similar to

those described by Henry and Dutrow (1996) and

Sperlich et al. (1996). The inner overgrowths are

separated from the core by sharp boundaries with a

discontinuous change in colour. Outer overgrowths

are more developed than the inner overgrowths, and

are separated from the latter by diffuse boundaries.

Where the inner overgrowths are pale green, the outer

overgrowths are deep green, and vice versa. Quartz

and opaque inclusions are common throughout the

overgrowths, mainly in the inner overgrowth.

3.3. Gneisses

In gneisses, tourmaline constitutes up to 10% by

volume. It occurs as disseminated subhedral to anhe-

dral clasts ( < 2 mm) in the matrix, and porphyroclasts

(normally < 1 cm) coexisting with porphyroclastic

feldspar (perthitic K-feldspar and albite), quartzFmuscovite. Epidote and Al-rich titanite are relatively

Table 1

Main characteristics of Sierra Nevada tourmaline

Tourmaline

types

Host rock Mineral

associationaGrain

size

Colourb Habit and texture Optical zoning

I tourmalinites Qtz, Ms, Bt,

Ep, Grt, Ttn

< 250 Am mainly pale

to dark green,

orange, brown

granoblastic texture, anhedral

to subhedral crystals, cellular

morphologies, reaction rims

and overgrowths

patchy zoning,

delicate growth

lamellae, sector zoning

II metasediments Qtz, Ms,

Bt, Grt, Ttn,

Rt, Zrn

< 800 Am pale to dark

green, orange

anhedral to subhedral crystals,

porphyroclasts and oriented

crystals within the foliated matrix,

distinct polar morphology,

pull-apart fractures

discontinous green

and orange cores

and pale green rims,

fine-growth lamellae

III gneisses Qtz, Kfs, Ab,

Ep, Ttn, Ap

< 1 cm green,

bluish green,

blue and

yellowish

brown

anhedral to subhedral crystals,

porphyroclasts and oriented

crystals within the foliated matrix,

distinct polar morphology, cellular

morphology, reaction rims and

overgrowths, pull-apart fractures

complex zoning

pattern, patchy zoning,

delicate growth zones,

discordant overgrowths

Qtz, quartz; Ms, muscovite; Bt, biotitte; Grt, garnet; Kfd, K-feldspar; Pl, plagioclasse; Ep, epidote; Ttn, titanite; Rt, rutile; Zrn, zircon.a Bolded minerals represent the modally dominant minerals in the samples.b Colour of the maximum absorption.


Fig. 3. Backscattered electron images of representative tourmaline from Sierra Nevada. The images have variable contrasts depending on the

associated minerals, so that contrasts do not exactly correspond to with compositional differences. BSE images of tourmaline in A–F are from

tourmalinites, G is from a schist, and H–L are from gneisses. (A) c-Axis section of tourmaline displaying fine-scale zonation in the interior

zones, and which is partially truncated by the rim. (B) Tourmaline showing fine-scale zonation that is replaced by patchy zoning in the interior

of the crystal. (C) Sector zoning in tourmaline with a concentric fine zonation. (D) Relics of complexely zoned tourmaline where the fine-scale

zonation has been embayed and truncated by pale rims. (E) Cellular textures in tourmaline grains that are rimmed by noncellular tourmaline. (F)

Detail of cellular morphology in a tourmaline crystal. (G) Tourmaline from metasediments displaying a Cr-bearing core surrounded by Cr-free

rim zones that do not show compositional polarity. (H) Tourmaline porphyroclast exhibing an asymmetric overgrowth of metamorphic

tourmaline and irregular patchs of the same composition as the overgrowth. (I) Prismatic tourmaline overgrowth on a porphyroclast displaying

fine-scale chemical zonation. (J) Irregular overgrowth of metamorphic tourmaline that truncate the finely zoned microstructure of the

porphyroclast. (K) Relict of abruptly, multiple zoned tourmaline surrounded by metamorphic tourmaline (right). Note the cellular texture on the

upper part of the associated porphyroclast. (L) Tourmaline overgrowth on a porphyroclast corroded with cellular texture.


abundant in some layers, and apatite appears as an

accessory. Tourmaline shows characteristic green,

bluish green, blue and yellowish brown colours. The

tourmaline commonly displays fine-scale growth

lamellae and irregular zones with complex colour-

zoning patterns (Fig. 3H–L). Some crystals exhibit

fine cellular pale overgrowths containing inclusions of

quartz. Generally, the overgrowths show an asymmet-

ric growth pattern with a different growth ratio and

acuteness of faces on opposite sides (Fig. 3H). This

growth pattern is similar to that developed by nucle-

ation on detrital tourmaline during diagenesis and

low-grade metamorphism (Henry and Dutrow, 1996;

Sperlich et al., 1996). The overgrowths may develop

on apparently homogeneous porphyroclasts (Fig. 3H)

and also on porphyroclasts exhibiting delicate growth

zones (Fig. 3I,J,K). Angular porphyroclasts of tour-

maline with fine-scale zonation commonly show

lamellae abruptly truncated by the overgrowths (Fig.

3J,K). As Tracy (1982) described in garnets, the break

in zoning between clast and overgrowth seems like a

stratigraphic ‘‘unconformity’’. Tourmaline porphyro-

clasts with growth zones embayed by pale tourmaline

are relatively common. Embayment may be so perva-

sive that, like tourmaline in the tourmalinites, cellular

morphologies are developed (Fig. 3K,L).

Lenses of quartz, feldspar, mica, and epidote in

places define a mylonitic foliation that wraps around

tourmaline and feldspar porphyroclasts. Isolated

quartz ribbons may also appear around the porphyr-

oclasts. Augen of feldspar surrounded by alternating

thin laminae of fine-grained feldspar, quartz, and

epidote occur in some layers. Tourmaline porphyro-

clasts are commonly accompanied by quartz–feldspar

strain shadows, normally asymmetric, and micaceous

strain caps. Mica may also be concentrated in the

edges of the feldspar porphyroclasts (‘‘quarter mats’’,

Hanmer and Passchier, 1991), probably due to pref-

erential dissolution of quartz by pressure solution at

high-stress sites adjacent to the porphyroclasts. Micro-

structures in tourmaline and feldspar porphyroclasts

reflect a different response to mylonitization. Whereas

tourmaline porphyroclasts are commonly transected

by microfractures and microfaults that promoted dis-

placement of fragments and pull-apart microstruc-

tures, feldspars show a more ductile behaviour and

recrystallization effects in microfractures and along

clast margins. However, the fracturing in K-feldspars

appears to have played a significant role in the de-

formation. Porphyroclasts of feldspar can be trans-

ected by microfaults that reduce the feldspar clast size

and cause relative displacement of the fragments.

Fine-grained feldsparF quartz in narrow recrystal-

lized bands are developed along these fracture zones.

In contrast, fractures in plagioclase are scarce; instead,

bent twins, undulatory extinction, kinking, and sub-

grains are relatively common. The presence of patch

and chessboard albite in some feldspar porphyroclasts

is indicative of albitization processes, and that the

mylonitization involved some compositional changes.

Likewise, feldspar porphyroclasts exhibiting narrow,

partial to complete mantles of fine-grained feldspar

with a sharp core–mantle boundary are also observed.

The mantles may be linked to tails that extend on both

sides of the porphyroclasts parallel to the foliated

matrix. More complex intermediate microstructures

between strain shadows and mantled porphyroclasts

can be also found. More complex intermediate micro-

structures between strain shadows and mantled por-

phyroclasts can be also found. Micas exhibiting a

lozenge shape (mica-fish) are relatively common in

these rocks, generally accompanied of trails of tiny

mica fragments that show stair-step textures.

4. Analytical methods

Chemical analyses were performed on 32 repre-

sentative rock samples, and 12 tourmaline separates

from gneisses, metasediments, and tourmalinites.

Tourmaline separates were obtained using an isody-

namic magnetic separator, and later treatment with

cold HF, which dissolved quartz and most other

minerals.

Major elements and Zr were analyzed at the

Basque Country University by X-ray fluorescence

(XRF), using a Philips PW1710 instrument. Trace

elements, except Zr, were determined at the Granada

University (CIC) by ICP-mass spectrometry (ICP-

MS), after HNO3 + HF digestion of 0.1000 g of

sample powder in a teflon-lined vessel at 180jC and

200 p.s.i. for 30 min, evaporation to dryness, and

subsequent dissolution in 100 ml of 4 vol.% HNO3.

Precision was better than F 5% for analyte concen-

trations of 10 ppm. Calibration was done with the

international standards PM-S and WS-E (Govindaraju


et al., 1994) dissolved in the same way (see Table 5).

Samples for Sr isotope analyses (seven tourmaline

separates) were digested in the same way using ultra-

clean reagents and analyzed by thermal ionization

mass spectrometry (TIMS) in a Finnigan Mat 262

spectrometer after chromatographic separation with

ion-exchange resins. Normalization and Blank values

were 86Sr/88Sr = 0.1194 and 0.6 nanograms, respec-

tively. The external precision (2r), estimated by

analyzing 10 replicates of standard WS-E (Govindar-

aju et al., 1994), was better than 0.003% for 87Sr/86Sr.

The 87Rb/86Sr were determined directly by ICP-MS,

following the method developed by Montero and Bea

(1988), with a precision better than 1.2% (2r). The87Sr/86Sr measured value in laboratory for NBS 987

international standard was 0.710250F 0.0000044 for

89 cases.

Mineral compositions were determined with a

Cameca SX50 electron microprobe at the University

of Granada (CIC), equipped with four wavelength-

dispersive spectrometers, using both natural and syn-

thetic standards: natural fluorite (F), natural sanidine

(K), synthetic MnTiO3 (Ti, Mn), natural diopside

(Ca), synthetic Fe2O3 (Fe), natural albite (Na), natural

periclase (Mg), synthetic SiO2 (Si), synthetic Cr2O3

(Cr), and synthetic Al2O3 (Al). An accelerating volt-

age of 20 kV, with a beam current of 30 nA and a

beam diameter of about 2 Am, was used to analyze

tourmaline and associated minerals. Counting times

on peaks were twice those of backgrounds, with 15 s

for Na and K; 20 s for Ti and Ca; 25 s for Fe, Si and

Al; and 30 s for Mg. Data were reduced using the

procedure of Pouchou and Pichoir (1985). The sample

data base includes over 1250 analyses of all tourma-

line types in gneisses, schists, and tourmalinites, and

over 300 analyses of associated minerals (micas,

garnet, titanite, epidote, and feldspars).

5. Whole-rock chemistry

Major and trace element data for representative

samples of the three lithologic groups that contain

tourmaline in Sierra Nevada are given in Tables 2 and

3. The analytical data are represented on a series of

variation diagrams with various element oxides and

trace elements plotted against Al2O3, because alumi-

nium is a major component and it is usually immobile

during metamorphic-hydrothermal processes (Fig. 4).

5.1. Major elements

The gneisses show a restricted range of composi-

tions that fall in a well-defined area on TiO2, MgO,

Fe2O3 vs. Al2O3 diagrams (Fig. 4). They are charac-

terized by high silica (SiO2>72.46 wt.%), peralumi-

nous character (aluminium saturation index ASI>1.5),

and low contents of TiO2 ( < 0.31 wt.%), Fe2O3

Table 2

Representative major element analyses (wt.%) of gneisses, metasediments, and tourmalinites from the Sierra Nevada

Sample Gneisses Metasediments Tourmalinites

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16

SiO2 73.38 72.46 74.56 75.18 73.09 69.70 63.85 63.01 59.35 71.46 71.04 72.21 74.23 60.77 73.50 68.18

TiO2 0.22 0.31 0.17 0.17 0.18 0.98 1.04 1.06 1.05 0.91 0.88 0.89 0.84 1.04 0.84 0.93

Al2O3 13.84 13.53 13.84 13.32 13.59 14.25 18.25 20.60 20.31 13.90 12.75 12.23 12.63 19.10 11.47 14.80

B2O3 0.21 0.99 0.76 0.26 0.21 0.13 0.05 0.16 0.08 0.14 3.72 3.17 3.07 5.10 2.75 2.48

Fe2O3(t) 2.48 2.09 2.53 1.79 2.12 5.43 7.69 4.79 8.70 5.01 7.52 7.15 5.34 8.21 5.46 6.18

MnO 0.01 0.01 0.01 0.01 0.01 0.04 0.05 0.01 0.09 0.02 0.02 0.03 0.02 0.03 0.03 0.04

MgO 0.45 0.41 0.49 0.44 0.53 0.92 1.29 1.22 1.48 0.98 1.19 1.10 1.15 1.80 0.90 1.01

CaO 0.59 1.34 0.39 0.32 0.30 0.50 0.75 0.38 1.08 0.38 0.72 0.77 0.65 0.69 1.01 0.70

Na2O 3.95 2.05 3.43 2.73 4.50 1.01 1.23 1.58 0.82 0.22 1.24 1.05 0.78 2.26 2.06 1.35

K2O 2.49 4.80 1.43 2.69 1.89 4.33 3.21 3.37 4.19 4.61 0.04 0.24 0.04 0.05 0.98 2.16

P2O5 0.07 0.10 0.08 0.15 0.06 0.77 0.15 0.09 0.17 0.05 0.07 0.08 0.05 0.13 0.08 0.19

F 0.15 0.09 0.08 0.16 0.08 0.10 0.09 0.07 0.11 0.13 0.14 0.15 0.15 0.15 0.13 0.13

OMF 0.06 0.04 0.04 0.07 0.03 0.04 0.04 0.03 0.05 0.05 0.06 0.06 0.06 0.06 0.05 0.05

LOI 1.96 1.09 1.70 1.81 1.84 2.11 2.51 3.14 2.68 2.62 0.82 1.13 0.85 1.24 1.20 1.87

Total 99.74 99.23 99.44 98.96 98.36 99.53 100.12 99.46 100.06 100.37 100.09 100.14 99.74 100.51 100.36 99.97


( < 2.60 wt.%), and MgO ( < 0.53 wt.%), with variable

amounts of Na2O (2.04–6.34 wt.%), K2O (0.34–4.80

wt.%), and B2O3 (0.09–1.08 wt.%). The ranges of

SiO2 (72.46–76.78 wt.%) and Al2O3 (12.40–13.84

wt.%), together with the fact that S(FeO +MgO) is

low relative to CaO, suggest an igneous protolith for

the gneissic rocks. On a plot of SiO2 vs. Na2O +K2O

(Le Maitre et al., 1989), the data are consistent with a

rhyolitic or granitic precursor for the gneisses from

Sierra Nevada. However, whether the gneisses have a

volcanic or plutonic origin is a matter of debate (Puga

and Fontbote, 1966; Puga, 1971; Nieto, 1996).

Unlike the gneisses, the schists are characterized

by variable contents of silica (SiO2 = 58.46–74.69

wt.%), with generally higher concentrations in

Al2O3 (up to 21.31 wt.%), Fe2O3 (up to 9.17 wt.%),

MgO (up to 1.55 wt.%), and TiO2 (0.84–1.07 wt.%).

B2O3 contents are generally below 0.30 wt.%. Some

Table 3

Representative trace elements analyses (ppm) of gneisses, metasediments and tourmalinites from the Sierra Nevada

Sample Gneisses Metasediments Tourmalinetes

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16

Li 102.5 6.0 36.1 94.2 18.9 212.5 193.9 73.0 124.1 49.9 131.3 123.1 80.0 166.0 88.6 83.9

Rb 330.5 20.3 85.8 297.8 207.9 285.6 277.3 191.4 198.8 336.5 9.3 15.9 5.3 5.1 45.8 164.7

Cs 12.8 1.3 6.2 8.2 7.9 49.5 64.9 13.3 13.9 41.2 3.0 1.1 0.3 0.5 11.8 18.6

Be 4.9 2.3 5.8 3.1 5.8 2.6 2.9 3.4 2.6 4.2 3.5 2.0 4.2 6.0 3.2 3.4

Sr 50.5 6.2 85.0 58.9 64.7 29.8 160.9 281.3 117.4 16.0 135.7 134.6 126.8 218.5 66.4 127.6

Ba 78.9 2.3 102.8 94.7 76.8 194.8 405.4 457.8 338.6 197.0 9.1 21.4 3.8 6.3 83.7 374.6

Zr 138 68 82 101 105 265 209 183 162 281 298 286 316 173 327 269.0

Hf 3.1 1.4 1.8 1.9 2.3 8.8 7.0 5.9 5.0 9.6 11.3 10.5 11.4 5.5 12.2 9.0

Nb 18.0 13.3 11.6 11.1 16.0 17.0 18.1 18.4 18.7 16.6 16.2 14.9 14.0 20.1 14.4 16.8

Ta 3.9 3.5 3.3 2.3 4.5 1.3 1.4 1.4 1.3 1.3 1.2 1.2 1.8 1.7 1.4 2.1

Th 24.5 4.2 32.8 8.6 37.3 5.7 13.5 14.5 10.1 6.2 9.9 8.9 10.5 20.5 8.6 8.8

U 13.5 3.2 9.6 2.1 8.9 2.9 4.4 4.3 2.6 7.2 2.7 2.5 4.7 3.1 3.6 4.2

V 14.3 8.9 16.0 18.7 18.9 116.6 151.4 155.9 160.7 109.3 120.1 103.6 85.2 163.1 85.2 113.7

Sc 6.0 6.4 18.4 18.6 18.9 12.6 16.2 16.0 14.2 10.0 10.9 11.1 21.7 28.7 8.0 10.9

Cr 5.5 3.2 36.2 44.1 38.4 74.6 92.5 101.8 100.8 68.4 78.5 69.4 104.0 141.4 57.7 74.2

Co 7.7 2.5 20.9 17.4 23.2 6.6 11.6 13.1 8.8 7.4 9.8 9.5 57.8 51.3 6.2 7.8

Ni 10.6 5.6 15.4 22.4 11.1 23.8 35.0 27.1 23.8 24.1 62.3 28.5 29.1 52.4 18.1 25.1

Cu 15.2 5.8 20.7 10.7 14.5 14.6 18.5 19.4 18.8 30.2 11.7 12.8 5.2 23.8 11.8 18.9

Zn 41.2 4.9 28.4 45.2 37.6 44.3 26.7 32.3 130.9 37.1 75.5 67.6 73.9 88.5 41.4 41.7

Ga 22.5 13.9 18.7 22.9 21.2 18.8 24.8 27.6 25.2 18.4 18.0 17.6 16.7 24.5 16.4 20.6

Mo 0.5 0.3 5.5 4.6 5.7 0.4 0.5 0.3 0.7 0.3 0.9 0.5 9.9 8.9 0.3 0.2

Sn 36.6 12.5 13.7 31.9 16.3 18.8 12.1 14.2 11.8 19.5 6.1 9.6 12.4 9.9 9.1 6.2

Tl 0.8 0.1 0.4 1.0 0.5 2.0 1.9 0.9 1.2 1.8 0.0 0.1 0.0 0.1 0.5 0.8

Pb 4.4 0.6 3.4 10.4 5.8 1.6 1.4 2.2 47.1 2.3 3.2 3.1 1.4 2.1 3.1 1.2

Y 21.35 0.46 16.92 8.13 31.74 15.43 38.12 18.21 24.42 15.08 14.31 18.52 11.94 13.70 20.56 14.7

La 24.35 1.76 10.84 7.46 34.76 14.22 41.82 60.24 35.84 19.01 19.33 21.26 6.10 15.05 15.73 15.7

Ce 54.49 8.07 21.67 16.25 75.56 34.01 85.35 117.00 69.42 93.38 38.25 47.07 14.11 35.15 40.26 35.9

Pr 5.49 0.46 2.82 2.00 8.84 3.61 9.34 13.64 8.80 5.37 3.78 5.24 1.63 4.06 4.26 4.3

Nd 18.66 1.58 9.83 7.35 30.54 13.36 36.10 48.75 32.59 19.55 13.52 19.06 5.67 15.85 16.58 16.9

Sm 3.90 0.28 2.36 1.74 7.00 2.73 7.01 9.28 6.46 4.07 2.75 3.79 1.44 3.23 3.55 3.7

Eu 0.42 0.06 0.47 0.26 0.75 0.56 1.36 1.91 1.40 0.75 0.73 0.82 0.43 0.67 0.67 0.6

Gd 3.24 0.32 2.33 1.80 6.70 2.65 6.18 7.61 5.72 3.59 2.26 3.48 1.57 2.83 4.09 3.4

Tb 0.57 0.06 0.44 0.30 1.06 0.47 1.04 1.00 0.89 0.62 0.38 0.57 0.31 0.47 0.67 0.6

Dy 3.66 0.40 2.95 1.75 5.86 3.32 6.68 4.81 5.31 3.90 2.66 3.74 2.20 2.76 4.20 3.0

Ho 0.78 0.09 0.63 0.30 1.21 0.75 1.42 0.77 1.03 0.73 0.56 0.79 0.52 0.52 0.88 0.6

Er 2.24 0.28 1.91 0.78 3.30 2.22 3.97 1.84 2.90 2.06 1.69 2.22 1.54 1.41 2.48 1.7

Tm 0.34 0.05 0.34 0.11 0.54 0.35 0.59 0.27 0.46 0.31 0.25 0.35 0.24 0.19 0.41 0.3

Yb 2.23 0.33 2.27 0.64 3.54 2.28 3.60 1.65 2.94 1.92 1.60 2.30 1.49 1.07 2.65 1.7

Lu 0.29 0.04 0.30 0.11 0.49 0.32 0.51 0.26 0.42 0.28 0.24 0.30 0.20 0.15 0.39 0.2


of the schists show chemical similarities to pelites

(relatively low SiO2, high Al2O3, and MgO>CaO),

whereas some others have higher SiO2, lower Al2O3,

and relatively low FeO +MgO, indicating a psammitic

parentage. Puga (1971) and Nieto (1996) suggested

that these rocks derive from tuffite protoliths. How-

Fig. 4. Major and trace element variations vs. Al2O3 for gneisses, metasediments and tourmalinites from Sierra Nevada, and plots of La vs. Th

and Yb vs. Y for gneisses, metasediments and tourmalinites from Sierra Nevada.


ever, the wide variation of Al2O3 (between 11.67%

and 21.31%), low Na2O, high Ti/Nb (272–406), and

the Al2O3–TiO2 trend support the interpretation that

the schists have a significant sedimentary component.

Chemical compositions of the tourmalinites are

comparable to those of the schists: variable amounts

of SiO2 (57.10–75.2 wt.%), and Al2O3 (10.58–21.12

wt.%), with smaller variations in Fe2O3 (5.34–8.59

wt.%), MgO (0.9–1.69), and Na2O (0.71–2.26

wt.%). Such major element variations account for

the modal volume and composition of the tourmaline

and coexisting minerals. Most of tourmalinites, how-

ever, show a depletion in K2O ( < 0.25 wt.%). On

plots of TiO2 vs. TiO2/Zr (Fig. 5), the schists and

tourmalinites fall along linear arrays that deviate

clearly from trends for the gneisses, suggesting a

psammopelitic nature for precursors of these tourma-

linites. In clastic sediments, TiO2/Zr correlates well

with the Al2O3 content and the range of TiO2/Zr

variations has been proposed as a chemical indicator

of sorting efficiency (Garcia et al., 1991). A TiO2/Zr

(� 100) ratio of 0.4 appears to show the approximate

limit between psammitic and pelitic materials in the

studied area, value that is consistent with the pub-

lished data on other shale and sandstone series (Garcia

et al., 1991).

Some major element oxides, particularly K2O, for

schists and tourmalinites display significant scatter on

variation diagrams (Fig. 4). Fe2O3 and MgO show

broad correlations with Al2O3. Covariation of TiO2

and Al2O3 indicates that these elements behaved

coherently during tourmalinization and regional meta-

morphism (Fig. 4). However, the Y-intercept (c 0.6)

suggests an original enrichment of Ti-minerals in the

sedimentary precursor, and accounts for the presence

of titanite in the metasediments.

5.2. Trace elements

Tourmalinites show a comparable pattern to the

metasediments except for Rb, Cs, Ba, and Cu that are

lower in the tourmalinites. Trace elements in the

metasediments and tourmalinites such as Ba, V, Cr,

Ga, Nb, and Zr vs. Al2O3 plots show a good corre-

lation that reflect, apart from the chemical consistency

between metasediments and tourmalinites, a similar

behaviour during hydrothermal and metamorphic pro-

cesses (Fig. 4). The other trace elements exhibit

complex variations. Concentrations of high field

strength elements (HFSE), except for Zr that show a

negative correlation with Al, appear to be controlled

by phyllosilicates and other neoformed minerals (gar-

net, titanite, epidote), taking into account the positive

correlation between HFSE and Al. Negative correla-

tion between Al2O3 and Zr, and positive correlation

between SiO2 and Zr (not shown), reflect preferential

occurrence of zircon in quartz-rich sedimentary pro-

toliths relative to some of the tourmalinites. On Zr

and V vs. Al plots, the gneisses are clearly discrimi-

nated (Fig. 4). Y and Nb contents in gneisses are

characteristic of collisional and volcanic arc environ-

ments, based on the fields defined by Pearce et al.

(1984).

Notwithstanding the chemical affinity described

above, the relative depletion of Rb, Ba, and Cs in the

tourmalinites compared to the metasediments reflects

the scarcity of micas in the former. Amounts of Li in

the metasediments (avg. 106 ppm) and tourmalinites

(avg. 113 ppm) are relatively high compared with the

gneisses (avg. 52 ppm). Concentrations of Be in the

tourmalinites (avg. 3.3 ppm) are slightly lower than

those in the metasediments (avg. 4.2 ppm) and

gneisses (avg. 4.6 ppm). Sn contents in the tourmalin-

ites (avg. 11.0 ppm) are also lower than those in the

metasediments (avg. 16.4 ppm) and gneisses (avg.

28.4 ppm).

Fig. 5. Variation of TiO2 against 100 TiO2/Zr. The vertical line at

100 TiO2/Zr = 0.4 represents the approximate limit between

psammitic and pelitic materials.


Fig. 6. Chondrite-normalized plots of REE abundances for gneisses, metasediments and tourmalinites, and Chondrite-normalized plots of REE

abundances in selected tourmalines from gneisses, metasediments and tourmalinites. Chondrite data from Boynton (1984).


5.3. Rare earth elements

Chondrite-normalized plots of REE abundances for

gneisses, metasediments, and tourmalinites reveal the

following characteristics (Fig. 6):

(1) Gneisses have a wide range of total REE (13.8–

180 ppm) and fractionated patterns with variable

enrichment of LREE relative to HREE (LaN/

LuN = 2.00–8.63), and large negative Eu anoma-

lies (Eu/Eu*= 0.31–0.62). Two of the samples

show relative depletion of REE and lower LaN/

LuN ratios (2.00 and 4.25), accompanied by

depletion in Y (Table 3, Fig. 6).

(2) Metasediments display REE patterns more homo-

geneous than those the gneisses, with total REE

abundances ranging from 80.8 to 269 ppm.

Shapes of the REE patterns are generally similar

for psammitic and pelitic metasediments, with

LaN/LuN ratios and Eu/Eu* values mostly rang-

ing from 4.64 to 8.78 and 0.45 to 0.70, res-

pectively. Excepting the three samples that show a

greater fractionation (LaN/LuN = 12.59 to 24.35),

the patterns are comparable to Post-Archean

Australian Average Shale (PAAS), thus suggest-

ing little REE mobility during regional meta-

morphism.

(3) Most of tourmalinites show variable abundances

of REE (SREE= 46.6–183.6 ppm), with fractio-

nated REE patterns (LaN/LuN = 3.17–9.32) and

variable negative Eu anomalies (Eu/Eu*= 0.51–

0.86). Relative to these, one sample has the

highest REE content (SREE= 330 ppm) with a

more fractionated pattern (LaN/LuN = 14.91) and a

pronounced negative Eu anomaly (Eu/Eu*= 0.31).

Two other samples show the lowest REE contents

(SREE= 28.3–38.7 ppm) with small negative Eu

anomalies (Eu/Eu*= 0.76–0.93).

Overall, REE patterns of the tourmalinites are

similar to those of the metasediments that, together

with the behaviour of HFSE, are consistent with a

psammopelitic parentage for these tourmalinites. Plots

of REE vs. Al2O3 lack well-defined trends (not

shown), suggesting some mobility of REE during

hydrothermal alteration and regional metamorphism

and/or the influence of accessory minerals on REE

systematics. The good correlation observed between

Y and Yb (Fig. 4) indicates a coherent behaviour of Y

relative to HREE during hydrothermal and metamor-

phic processes. On La–Th and Ce–Th plots (Fig. 4),

metasediments and gneisses define a La- and Ce-rich

trend and a Th-rich trend, respectively, suggesting a

different distribution of accessory minerals. In this

way, the good correlation between Nd and Th for the

metasediments, with Nd/Th ratios close to 2.6,

appears to be characteristic of monazite-bearing crus-

tal materials (Bea and Montero, 1999). Lower Nd/Th

ratios in the gneisses could reflect the presence of

other accessory minerals such as thorite. The Nd/Th

values in the tourmalinites are of the same order as

those in the metasediments.

6. Tourmaline chemistry

Tourmaline has a complex chemistry and its basic

formula can been written as XY3Z6(T6O18)(BO3)

V3W (Hawthorne and Henry, 1999), where X =Na,

Ca, K, vacancy; Y= Fe2 +, Mg, Mn2 +, Li, Al, Cr3 +,

V3 +, Fe3 +, (Ti4 +); Z =Mg, Al, Fe3 +, V3 +, Cr3 +; T = Si,

Al, (B); B =B, (vacancy), V=OH, O; W=OH, F, O.

As tourmalines from Sierra Nevada have insignif-

icant amounts of Li, the structural formulae were

calculated using the cation normalization procedure

Y +Z+T= 15. This assumes no vacancies on the Y,

Z and T sites and is not a function to the oxidation

states of Fe or the amount of OH on the W and V

sites. To a first approximation, the sum of the cation

charges in the X +Y+Z+T sites (>49) does not

appear to indicate the presence of Fe3 + in tourma-

lines of Sierra Nevada.

The tourmaline-bearing samples are grouped here

into three categories: (1) tourmalinites, (2) metasedi-

ments, and (3) gneisses. The last category can be

subdivided into mica-free gneisses (TFG) and mica-

bearing gneisses (TFMG). Statistical data from

microprobe analyses of tourmaline from each category

are summarized in Table 4. As a whole, results for the

Sierra Nevada tourmalines correspond in composi-

tional space to Li-poor aluminous tourmaline. ICP

data and charge-balance constraints are consistent

with negligible Li. Most of the tourmalines belong

to the alkali group, which have vacancies in the X site,

up to 0.55 apfu (Fig. 7). Variations in their chemical

compositions are represented here on the Al–Fe(tot)–


Mg and Ca–Fe(tot)–Mg ternary diagrams of Henry

and Guidotti (1985) (Fig. 7).

6.1. Tourmalinites

Most tourmalines from the tourmalinites fall with-

in compositional fields that correspond to tourmalines

from clastic metasediments and Li-poor granitoids

(Fig. 7). The porphyroclasts and cores of the tourma-

lines show a wider compositional variation than the

rims. Microprobe data indicate that much of the

chemical variability is in the atomic ratios of Na/

(Na +Ca) and Fe/(Fe +Mg) (Fig. 8). These ratios

range from 0.5 to 0.93 and 0.32 to 0.97, respectively,

and in general lack a systematic. The rims have a

more restricted composition with Na/(Na +Ca) and

Fe/(Fe +Mg) falling between 0.80–0.91 and 0.30–

0.55, respectively (Fig. 8). Ferromagnesian contents

are significantly less than 3 apfu for all tourmaline

grains, reflecting the presence of Al in the Y site (Fig.

8). A substantial portion of the Al is located in the Y

site via the proton- and alkali-deficient vectors (Fig.

8). Data that fall to the right of the proton-loss

substitution trend may indicate the presence of tetra-

hedral Al through the tschermak substitution. Ti

contents are low and variable, less than 0.3 apfu,

showing a broad positive correlation with Fe/

(Fe +Mg) ratios (Fig. 8). The amount of Ca is

variable ( < 0.44 and < 0.18 apfu for cores and rims,

respectively, Fig. 8), and the exchange vectors:

CaMgNa� 1Al� 1, CaMg2(OH)Na� 1Al� 2O� 1, and

CaONa� 1 (OH)� 1, seem to be possible to explain

Ca incorporation in the tourmaline cores (Fig. 9).

6.2. Metasediments

Tourmalines from the metasediments plot mostly in

the field defined for tourmalines from clastic meta-

sediments, with Mg contents higher than in the

tourmalinites (Fig. 7). A plot of Na/(Na +Ca) vs. Fe/

(Fe +Mg) reveals a positive correlation for the tour-

maline cores, with a wide variation in Fe/(Fe +Mg)

ratios (Fig. 8). In contrast, the rims have a constant

Na/(Na +Ca) ratio and variable Fe/(Fe +Mg) values

(Fig. 8). Contents of (Fe +Mg) are less than the

theoretical 3.0 apfu of schorl–dravite, mainly toward

Table 4

Average compositions of tourmalines

Gneisses Not-tourmalinized metasediments Tourmalinites

Clasts Rims (TFG) Rims

(TFMG)

Cores Cr-bearing

Cores

Rims Cores Rims

Si 6.052 (0.053) 6.080 (0.057) 6.064 (0.064) 5.979 (0.066) 5.842 (0.119) 5.994 (0.051) 5.937 (0.059) 6.002 (0.038)

Al 5.774 (0.284) 6.413 (0.138) 5.964 (0.096) 6.291 (0.177) 6.588 (0.184) 6.016 (0.156) 6.178 (0.201) 6.148 (0.084)

Cr 0.001 (0.001) 0.001 (0.001) 0.001 (0.001) 0.005 (0.004) 0.075 (0.052) 0.004 (0.003) 0.003 (0.004) 0.002 (0.001)

Ti 0.144 (0.070) 0.083 (0.017) 0.098 (0.016) 0.081 (0.028) 0.135 (0.021) 0.137 (0.077) 0.218 (0.064) 0.102 (0.014)

Fe 1.895 (0.595) 1.944 (0.139) 1.733 (0.283) 1.027 (0.381) 1.474 (0.094) 1.435 (0.361) 1.813 (0.323) 1.210 (0.148)

Mg 1.121 (0.667) 0.455 (0.198) 1.135 (0.247) 1.614 (0.491) 0.844 (0.142) 1.413 (0.319) 0.849 (0.407) 1.535 (0.200)

Mn 0.014 (0.011) 0.026 (0.014) 0.005 (0.003) 0.002 (0.004) 0.002 (0.002) 0.001 (0.001) 0.002 (0.002) 0.001 (0.001)

Ca 0.269 (0.165) 0.120 (0.024) 0.144 (0.034) 0.135 (0.119) 0.150 (0.038) 0.095 (0.069) 0.236 (0.090) 0.108 (0.021)

Na 0.634 (0.157) 0.719 (0.053) 0.754 (0.052) 0.645 (0.091) 0.610 (0.050) 0.809 (0.062) 0.530 (0.092) 0.744 (0.055)

F 0.428 (0.125) 0.296 (0.045) 0.299 (0.044) 0.268 (0.067) 0.453 (0.174) 0.254 (0.057) 0.170 (0.058) 0.220 (0.026)

X-vac 0.084 (0.066) 0.146 (0.051) 0.087 (0.040) 0.211 (0.088) 0.255 (0.047) 0.084 (0.079) 0.223 (0.070) 0.139 (0.050)

vac/(vac +Na) 0.117 (0.078) 0.169 (0.057) 0.104 (0.048) 0.241 (0.085) 0.269 (0.054) 0.085 (0.142) 0.296 (0.082) 0.157 (0.058)

Mg/(Mg+Fe) 0.366 (0.205) 0.187 (0.070) 0.397 (0.092) 0.604 (0.161) 0.374 (0.046) 0.498 (0.118) 0.314 (0.140) 0.558 (0.065)

xs charge 0.339 (0.243) 0.712 (0.162) 0.347 (0.158) 0.331 (0.066) 0.542 (0.121) 0.293 (0.234) 0.504 (0.102) 0.139 (0.050)

W(OH) 0.233 (0.228) � 0.008 (0.160) 0.354 (0.166) 0.397 (0.084) 0.005 (0.124) 0.453 (0.240) 0.325 (0.111) 0.453 (0.052)

R+ +R2 + 4.215 (0.413) 3.397 (0.169) 3.930 (0.121) 3.568 (0.353) 3.286 (0.154) 3.860 (0.191) 3.677 (0.282) 3.715 (0.120)

R3 + 5.966 (0.263) 6.523 (0.147) 6.095 (0.099) 6.399 (0.168) 6.768 (0.194) 6.199 (0.117) 6.469 (0.194) 6.284 (0.081)

Na/(Na +Ca) 0.705 (0.174) 0.857 (0.029) 0.839 (0.039) 0.836 (0.135) 0.803 (0.049) 0.897 (0.054) 0.693 (0.105) 0.873 (0.026)

No. 243 63 78 67 60 71 95 26

Numbers in parentheses are standard deviation.


Fe-rich compositions that suggest an enrichment in Al

as the Fe content increases (Figs. 7 and 8). The Cr-

bearing orange cores fall within the region correspond-

ing to Li-poor granitoids (Fig. 7); they contain small

amounts of Cr (up to 0.06 apfu) and have a very

restricted composition with lower (Fe +Mg) contents

reflecting their more aluminous character (Fig. 7).

Many tourmaline compositions plot on the right of

the proton-loss substitution that may account for the

presence of tetrahedral Al (Fig. 9). Except for the

orange cores, tourmalines show an increase of Ti with

increasing Fe/(Fe +Mg) (Fig. 8). On plots of Ca +Mg*

vs. Na +Al*, tourmalines exhibit linear trends, except-

ing the Cr-bearing cores, which are consistent with the


CaONa� 1(OH)� 1 substitutions (Fig. 9). However,

on plots of Na and X-site vacancies vs. Ca, the

influence of the alkali deficiency vector is noteworthy

(Fig. 9).

6.3. Gneisses

Many tourmalines from gneisses plot in the com-

positional fields of clastic metasediments and Li-poor

granitoids (Fig. 7). Significantly, and in contrast to

tourmaline from the mica-free gneisses, tourmaline

Fig. 7. Ternary (Ca-X-site vacancy-Na +K) diagram (after Hawthorne and Henry, 1999), and ternary Al–Fe–Mg and Ca–Fe–Mg plots (after

Henry and Guidotti, 1985) showing electron microprobe data for different textural types of tourmaline in gneisses, metasediments and

tourmalinites from Sierra Nevada (Spain). TFG= tourmaline and feldspar-bearing gneisses. TFMG= tourmaline, feldspar, and mica-bearing

gneisses.


Fig. 8. Plots of Na/(Na +Ca) vs. Fe/(Fe +Mg), X-site vacancy vs. Altotal, Fe vs. Mg, and Ti vs. Fe/(Fe +Mg) for tourmalines from Sierra Nevada.

Symbols as in Fig. 7. TFG and TFMG abbreviations as in Fig. 7.


Fig. 9. Plots of Na +Al* vs. Ca +Mg*, Na vs. Ca, and X-site vacancies vs. Ca for tourmalines from Sierra Nevada. Symbols as in Fig. 7.

Al*=Al3 + + Fe3 + + 2Ti–Li; Mg*=Mg+Fe2 + +Mn+ 2Li–Ti (Henry and Dutrow, 1990). R+ +R2 +=(Na+ + 2Ca2 + +K+ + Fe2 + +Mg2 + +Mn2 +).

R3 +=(Al3 + + 4/3Ti4 +). TFG and TFMG abbreviations as in Fig. 7.


compositions from mica-bearing gneisses define a

trend where the overgrowths are enriched in Mg

relative to the porphyroclasts (Fig. 10). Some tour-

maline porphyroclasts fall into the calcic group (Fig.

7) and plot below the schorl –dravite line with

(Fe +Mg)>3 (Fig. 8), suggesting the presence of a

significant uvite component in these tourmalines. A

plot of Na(Na +Ca) vs. Fe/(Fe +Mg) ratios for por-

phyroclasts shows a broad positive correlation in the

range of 0.36–0.90 and 0.28–0.95, respectively,

whereas the Na(Na +Ca) ratios for overgrowths are

within a narrow range, from 0.75 to 0.90 (Fig. 8). Data

for Ti vs. Fe/(Fe +Mg) also show a broad positive

correlation for cores, whereas Ti values remain nearly

constant for the overgrowths (Fig. 8). On Ca +Mg*

vs. Na +Al*, Na vs. Ca and X-vacancies vs. Ca

plots, data for the tourmaline porphyroclasts fall

approximately on linear arrays, suggesting that the

most consistent mechanisms for the incorporation of

Ca into the tourmalines involve the substitutions


CaONa� 1OH� 1 (Fig. 9). The overgrowths have

high Na and low Ca contents, with the Ca contents

remaining constant despite variations of Na and X-

site vacancies (Fig. 9).

6.3.1. Chemical zoning

Tourmalines from the Sierra Nevada generally

display optical zoning patterns in which significant

changes in composition define three main types of

chemical zoning related to the lithology:

(1) A discontinuous core-to-rim zoning wherein cores

are sharply overgrown by chemically distinct

tourmaline. This type of zonation is typical of

tourmalines from metasediments where the over-

Fig. 10. Ternary Al–Fe–Mg and Ca–Fe–Mg plots (after Henry and Guidotti, 1985) for tourmalines from mica-free gneisses and mica-bearing

gneisses. Note that compositional trend from clast to rim is different depending on the presence or absence of mica.


growths are believed to have formed during

metamorphism (Fig. 3G). Some cores are rela-

tively rich in Cr (up to 1.84 wt.% Cr2O3). Near the

boundary with the optical discontinuity is a

significant compositional gradient involving two

main trends: (i) one in which Al, Mg, and Ca

decrease abruptly, whereas Fe, Na, and Ti increase

(Figs. 11A and 3G). Si shows very small

variations and F varies in an irregular manner.

This zoning pattern is presumably a consequence

of a combination of FeMg� 1, NaFeAl� 1()� 1,

TiFeAl� 2 and Na(OH)Ca� 1O� 1 substitutions.

(ii) Another trend in which Al, Fe, Ca, Ti, and F

decrease, whereas Mg, Na, and Si increase (Fig.

11B). This trend probably accounts for substi-

tutions such as SiMgAl� 2, Si2MgAl� 2Ti� 1,

MgFe� 1, NaMgAl� 1()� 1, (OH)F� 1, Mg(OH)

Al� 1O� 1 and Na(OH)Ca� 1O� 1. Despite asym-

metric overgrowths in some crystals, there is no

evidence of a compositional polarity such as

expected from the metamorphic conditions in this

region. At medium grade of metamorphism,

compositional polarity disappears in metamorphic

tourmaline (Sperlich et al., 1996; Henry and

Dutrow, 1996).

(2) Tourmaline porphyroclasts displaying an abrupt

multiple, oscillatory zoning, which are rimmed by

continuous or discontinuous overgrowths (typical

of gneisses). The oscillatory zoning patterns show

no systematic variation in composition across the

crystals. Most of the analyzed porphyroclasts from

TFG at the clast-overgrowth boundary have a

significant compositional discontinuity such that

Al, Fe, and Na increase, whereas in the over-

growths, Mg, Ca, Ti, and F decrease (Figs. 11C

and 3J). The compositional changes through the

Fig. 11. Representative chemical zoning patterns for tourmaline from Sierra Nevada (Spain). (A and B) Compositional profiles for tourmaline

associated to metasediments (sections parallel to the c-axis). Note the lack of compositional polarity. (C and D) Microprobe traverses across

tourmaline porphyroclast-overgrowth from mica-free and mica-bearing gneisses, respectively. (E and F) Compositional zoning patterns for

tourmalines from tourmalinites. All the elements show more or less pronounced fluctuations.


clast-overgrowth interface may be the result of

a combination of exchange vectors such as

MgFe� 1, NaAlCa� 1Mg� 1, Na(OH)Ca� 1O� 1,

and OHF� 1. In contrast, porphyroclasts from

TFMG are characterized by an increase of Al, Na

and Mg, and a decrease of Fe, Ca, Ti, and F in the

overgrowths (Figs. 11D and 3H), which reflect the

substitutions MgFe� 1, ( )AlNa� 1Fe� 1, CaONa� 1

Table 5

Trace elements, REE and strontium isotope analyses (ppm) of toumaline separates

Sample Gneisses Metasediments Tourmalinites Standardsa

1 3 7 8 9 11 12 13 14 15 16 PM-S WS-E

Li 530.0 565.9 30.1 36.7 53.2 261.8 291.8 204.1 308.5 269.8 286.6 7.58 13.49

Rb 4.1 1.6 13.4 2.6 5.3 0.1 6.3 10.2 1.1 1.1 5.5 1.08 25.75

Cs 2.2 1.6 2.3 16.5 1.9 1.0 1.3 2.0 0.3 0.0 0.1 0.48 0.51

Be 2.6 3.8 1.1 0.9 0.5 5.1 3.6 3.4 4.5 4.6 4.5 0.45 1.10

Sr 591.8 573.1 123.7 208.6 329.5 324.5 365.3 283.6 337.2 266.3 290.0 290.07 415.92

Ba 11.9 14.6 23.0 20.3 22.0 2.8 9.8 8.3 1.8 6.7 1.9 153.44 336.42

Zr 9.1 5.6 13.3 17.4 15.1 14.6 16.8 7.9 11.9 8.2 7.0 36.96 196.57

Hf 0.5 0.3 1.8 1.6 1.5 3.2 2.8 2.9 2.8 2.8 1.9 1.03 5.22

Nb 1.6 2.8 7.9 2.5 10.9 4.0 4.9 2.2 3.2 2.4 1.9 2.26 17.54

Ta 1.2 0.8 0.8 0.3 0.8 0.6 0.5 0.2 0.4 0.7 0.6 0.22 1.17

Th 7.74 8.12 4.93 11.87 17.62 10.40 16.52 2.47 3.57 5.81 5.80 0.06 2.97

U 1.4 1.9 0.8 1.1 2.9 2.0 2.5 0.4 0.4 0.7 0.7 0.03 0.59

V 177.4 150.9 2054.2 457.1 292.4 256.7 324.0 342.2 310.0 291.5 234.6 189.79 337.51

Sc 67 46 88 25 3 25 25 26 21 27 17 34.36 29.28

Cr 26.5 48.2 1154.5 294.8 307.1 161.7 205.4 183.5 169.0 176.6 134.5 310.47 101.03

Co 7.4 5.4 34.6 18.6 38.2 20.2 11.6 10.8 9.5 13.3 17.3 47.88 43.70

Ni 5.7 11.6 105.9 93.5 225.1 76.0 54.4 81.3 58.7 55.0 64.4 115.50 52.74

Cu 5.5 4.8 0.0 4.6 6.3 7.4 5.7 6.7 6.8 4.5 8.4 57.84 67.56

Zn 323.6 340.7 228.2 116.2 590.1 197.9 160.3 106.8 120.2 168.5 165.0 59.25 125.34

Ga 83.9 70.8 43.6 34.4 28.9 44.9 44.6 40.1 42.2 47.8 43.7 15.75 22.87

Mo 0.3 0.2 0.2 0.5 0.2 0.0 0.1 0.1 0.1 0.2 0.1 1.83 3.66

Sn 189.8 74.9 0.0 12.6 0.0 48.2 59.7 34.9 4.3 6.5 27.5 3.03 18.52

TI 0.0 0.0 0.1 0.1 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.04 0.19

Pb 5.4 3.7 3.5 2.3 39.9 3.7 2.3 1.2 1.9 1.7 2.6 2.47 12.85

Y 14.6 17.4 8.9 15.8 4.8 14.2 36.9 12.3 2.9 11.6 2.7 11.48 30.86

La 19.95 18.89 23.72 132.83 62.72 35.74 88.30 10.03 5.16 21.75 9.50 2.81 27.01

Ce 38.35 36.95 58.96 280.24 119.72 83.89 199.36 21.00 12.38 40.38 23.56 6.86 60.06

Pr 3.86 4.07 5.70 32.11 15.88 11.29 22.89 2.08 1.14 4.11 2.36 1.09 7.68

Nd 12.16 14.36 21.17 118.61 58.56 46.80 83.42 6.93 4.04 14.47 8.31 5.76 33.35

Sm 2.56 3.30 3.88 22.15 10.78 10.23 16.67 1.22 0.77 2.64 1.71 1.75 8.67

Eu 0.43 0.47 0.94 4.50 2.16 1.78 2.33 0.52 0.29 1.10 0.58 1.08 2.25

Gd 2.30 3.15 2.87 13.97 6.41 7.09 11.09 1.34 0.65 1.81 1.23 2.15 6.85

Tb 0.40 0.55 0.41 1.48 0.56 0.94 1.61 0.30 0.11 0.26 0.15 0.37 1.08

Dy 2.47 3.08 2.23 5.86 1.94 4.20 8.00 2.36 0.77 1.39 0.89 2.12 5.89

Ho 0.48 0.58 0.42 0.76 0.23 0.57 1.40 0.60 0.16 0.30 0.16 0.41 1.19

Er 1.38 1.58 1.03 1.57 0.46 1.22 3.99 1.81 0.40 0.93 0.37 1.13 2.95

Tm 0.23 0.24 0.16 0.18 0.06 0.18 0.62 0.27 0.07 0.16 0.06 0.19 0.43

Yb 1.38 1.50 1.00 1.09 0.37 1.11 3.90 1.69 0.39 1.07 0.36 1.02 2.46

Lu 0.21 0.22 0.15 0.13 0.06 0.16 0.61 0.23 0.05 0.15 0.05 0.16 0.37

Rb (ppm) 4.10 1.59 2.62 5.24 0.12 1.08 1.13

Sr (ppm) 591.75 573.14 208.54 329.54 324.54 337.18 266.31

Rb87/Sr86 0.0207 0.0080 0.0364 0.0461 0.0010 0.0093 0.0115

Sr87/Sr86 0.7126 0.7120 0.7187 0.7284 0.7195 0.7168 0.7175

a Data for international standards (Govindaraju et al., 1994) analyzed in Granada laboratory.


(OH)� 1, and Al2Ti� 1Fe� 1. In this case, however,

compositional variations within the porphyroclasts

indicate that the uvite substitution has been

operative.

(3) Fine-scale oscillatory chemical zoning in subhe-

dral to euhedral tourmaline from tourmalinites that

commonly shows patchy zoning with in interior

growth zones (Fig. 3B,D,E). The oscillatory

zonation involves significant fluctuations in Al,

Fe, Mg, Na, Ca, Ti, and F (Figs. 11E and 3E), but

there is no systematic variation in the composi-

tional profiles of individual crystals. The discord-

ant reaction rims, pale green patchs and cells of

cellular textures, are characterized by an increase

of Mg and Na accompanied by the decrease of Al,

Fe, Ca and Ti (Figs. 11F and 3D).

6.3.2. Trace elements

Trace elements and REE in tourmaline (Table 5)

were determined on tourmaline separates from gneisses

(two samples), non-tourmalinized metasediments (two

samples), and tourmalinites (seven samples). In gen-

eral, relative to host rocks, the tourmalines are depleted

in Rb, Cs, Ba, Nb, Ta, U and Th, as enriched in Sr, Sc,

V, Cr, Co, Ni, Zn, and Ga. Tourmalines from the

gneisses have higher values of Li (530–566 ppm)

and Sn (75–190 ppm) than those from the tourmalin-

ites (262–309 and 4.3–59.7 ppm, respectively). Tour-

malines from metasediments are characterized by

lower Li (30–53 ppm) and insignificant Sn (0–12.6

ppm). The lower Li contents in the latter are due to

preferential Li partitioning into other phases (e.g.,

micas) rather than in tourmaline (Dutrow et al., 1986).

Table 6

Average compositions of biotite (Bt) and muscovite (Ms)

Gneisses Metasediments Tourmalinites

Bt (n= 10) Ms (n= 32) Bt (n= 8) Ms (n= 21) Bt (n= 8) Ms (n= 5)

SiO2 35.96 (0.70) 48.25 (0.65) 35.98 (1.99) 48.48 (1.27) 34.84 (0.26) 47.91 (0.86)

TiO2 2.46 (0.47) 0.39 (0.15) 2.92 (1.28) 0.47 (0.13) 2.51 (0.18) 0.63 (0.17)

Al2O3 16.07 (1.05) 29.20 (1.23) 16.17 (0.89) 31.42 (2.56) 16.44 (0.39) 30.60 (0.57)

Cr2O3 0.03 (0.03) 0.01 (0.01) 0.02 (0.01) 0.02 (0.02) 0.03 (0.03) 0.02 (0.01)

FeO 25.95 (3.57) 4.720 (0.78) 21.41 (5.40) 3.16 (1.49) 27.61 (0.98) 4.57 (1.01)

MnO 0.11 (0.06) 0.02 (0.02) 0.03 (0.01) 0.01 (0.02) 0.09 (0.03) 0.03 (0.02)

MgO 5.14 (2.56) 1.52 (0.34) 8.39 (3.84) 1.77 (0.38) 4.10 (1.02) 1.29 (0.51)

CaO 0.02 (0.03) 0.03 (0.05) 0.06 (0.07) 0.01 (0.01) 0.00 (0.01) 0.01 (0.01)

Na2O 0.07 (0.02) 0.21 (0.19) 0.06 (0.01) 0.71 (0.48) 0.05 (0.01) 0.27 (0.07)

K2O 9.05 (0.24) 10.19 (0.58) 8.75 (0.56) 9.43 (1.03) 9.20 (0.07) 10.40 (0.37)

F 0.76 (0.19) 0.46 (0.23) 0.90 (0.23) 0.30 (0.15) 0.48 (0.03) 0.31 (0.12)

Cl 0.64 (0.28) 0.01 (0.01) 0.30 (0.05) 0.01 (0.01) 0.82 (0.06) 0.01 (0.00)

OMF 0.46 (0.09) 0.20 (0.10) 0.45 (0.10) 0.13 (0.06) 0.39 (0.02) 0.13 (0.05)

Total 95.80 (0.69) 94.83 (0.58) 94.54 (0.82) 95.65 (1.12) 95.81 (0.39) 95.91 (1.41)

Si 2.803 (0.023) 3.267 (0.037) 2.770 (0.065) 3.219 (0.073) 2.753 (0.008) 3.210 (0.009)

ALIV 1.197 (0.023) 0.733 (0.037) 1.230 (0.065) 0.781 (0.073) 1.247 (0.008) 0.788 (0.009)

AlVI 0.280 (0.071) 1.598 (0.064) 0.239 (0.092) 1.675 (0.098) 0.285 (0.039) 1.631 (0.018)

Ti 0.144 (0.028) 0.020 (0.008) 0.171 (0.079) 0.023 (0.007) 0.149 (0.011) 0.032 (0.009)

Cr 0.001 (0.001) 0.000 (0.000) 0.000 (0.000) 0.000 (0.000) 0.001 (0.001) 0.000 (0.000)

Fe2 + 1.696 (0.260) 0.268 (0.045) 1.389 (0.395) 0.177 (0.085) 1.825 (0.071) 0.256 (0.053)

Mn 0.007 (0.004) 0.001 (0.001) 0.002 (0.001) 0.001 (0.001) 0.006 (0.002) 0.002 (0.001)

Mg 0.593 (0.286) 0.154 (0.034) 0.954 (0.417) 0.175 (0.037) 0.483 (0.119) 0.130 (0.052)

Ca 0.002 (0.002) 0.002 (0.003) 0.005 (0.006) 0.000 (0.001) 0.008 (0.001) 0.000 (0.001)

Na 0.011 (0.003) 0.028 (0.024) 0.009 (0.002) 0.091 (0.061) 0.008 (0.002) 0.035 (0.010)

K 0.900 (0.033) 0.881 (0.053) 0.862 (0.080) 0.892 (0.097) 0.928 (0.012) 0.890 (0.045)

F 0.186 (0.044) 0.099 (0.050) 0.217 (0.049) 0.063 (0.032) 0.119 (0.007) 0.065 (0.024)

Cl 0.085 (0.039) 0.001 (0.001) 0.039 (0.006) 0.001 (0.001) 0.110 (0.008) 0.002 (0.000)

Mg/(Mg+ Fe) 0.258 (0.116) 0.365 (0.071) 0.406 (0.173) 0.521 (0.119) 0.208 (0.049) 0.337 (0.136)

Structural formula of micas on the basis of 11 O. Numbers in parentheses are standard deviation.


REE contents of tourmalines from the gneisses,

tourmaline-poor metasediments, and tourmalinites are

given in Table 5. Whereas REE patterns for the

tourmalines from gneisses mimic those of the host

rocks, the REE patterns for tourmalines from meta-

sediments and tourmalinites show appreciable differ-

ences relative to their host rocks. On the other hand,

tourmalines from the metasediments lack negative Eu

anomalies, and tourmalines from tourmalinites that

have lower REE contents display positive Eu anoma-

lies; this feature probably reflects a contribution of Eu

from the hydrothermal fluids (Fig. 6). Tourmaline

from metasediments and tourmalinites shows an

LREE enrichment and HREE depletion relative to

its host rock, whereas turmaline from gneisses only

shows an HREE depletion (Fig. 6).

6.3.3. Strontium isotopes

Results of Rb–Sr isotopic analyses of tourmaline

separates from gneisses (two samples), metasediments

(two samples), and tourmalinites (three samples) are

listed in Table 5. Taking into account the low Rb/Sr

ratios of tourmalines (0.001–0.04), their 87Sr/86Sr

ratios may be approximated by the initial ratios of

the samples. The initial 87Sr/86Sr ratios for tourma-

lines from the tourmalinites (87Sr/86Sr = 0.7167–

0.7195) and from metasediments (0.7184–0.7187)

are very similar, but more radiogenic than those for

tourmalines from the gneisses (87Sr/86Sr = 0.7119–

0.7125). These data seem to indicate that 87Sr/86Sr

isotope ratios in the tourmalines were acquired from

two different reservoirs, which may correspond to

protoliths of the gneisses and metasediments. Further-

more, such data support a common precursor for

metasediments and tourmalinites.

7. Composition of associated minerals

7.1. Micas

Micas are major components of the gneisses and

metasediments, but are subordinate in the tourmalin-

ites; muscovite predominates over biotite. Muscovite

is phengitic with a composition between 2.15VAl

(IV) +Al(VI)V 2.73 and 0.37V Fe +Mg+Ti + SiV0.86; muscovite from the metasediments has the

highest values of AlTOT (av. 2.456 apfu) and Mg/

(Mg + Fe) ratio (av. 0.521) (Table 6). Biotites have a

compositional range between 1.34VAlTOTV 1.60

and 2V Fe +Mg+Ti + Si� 3V 2.43. AlTOT is some-

what higher for biotites from tourmalinites (av. 1.532

apfu) than from metasediments (av. 1.469 apfu) and

gneisses (av. 1.477 apfu), whereas the Mg/(Mg + Fe)

ratio is lower for tourmalinite-hosted biotites (av.

0.208) than for those in metasediments (av. 0.406)

and gneisses (av. 0.258) (Table 6).

Correlations between Mg/Fe ratios of the micas and

reaction rims of the tourmalines are illustrated in Fig.

12. Calculated distribution coefficients (Kd) for bio-

tite–tourmaline and muscovite–tourmaline pairs are

f 0.53 and f 083, respectively. The array of data

suggests that the late tourmaline and micas probably

attained chemical equilibrium during regional meta-

morphism.

7.2. Garnet

Garnet commonly occurs as colourless, very fine to

fine-grained, subhedral to subrounded crystals. In

most of the rocks, it is almandine-rich with a restricted

range in spessartine component ( < 7 mol%) and a

wide range in grossularite component (up to 53.4

mol%), particularly in mica-bearing gneisses that have

Fig. 12. Plot of Mg/Fe ratios of tourmaline and micas in metase-

diments. Tourmaline data correspond exclusively to pale discordant

rims.


46.0–53.4 mol% grossularite. Garnet in the metasedi-

ments has more variable and higher Mg contents

(1.4–20.0 mol% pyrope) than in the gneisses (0.6–

0.9 mol% pyrope) and tourmalinites (2.7–3.7 mol%

pyrope) (Table 7). In the metasediments, garnet dis-

plays a slight variation in composition from cores to

rims, the variation generally recording an increase in

the Mg/Fe ratio and a decrease in Ca.

7.3. Epidote

Epidote appears mainly in the gneisses where it

forms colourless, subhedral to anhedral, very fine-

grained crystals, and spot-forming recrystallized

grains together with tourmaline. Epidote in mica-free

gneisses is clinozoisite-rich (av. 97.27 mol%) with

minor amounts of epidote (av. 2.67 mol%) and pie-

montite (av. 0.07 mol%) components. Epidote in

mica-bearing gneisses comprises clinozoisite (av.

82.07 mol%), epidote (av. 17.73 mol%), and piemon-

tite (av. 0.20 mol%). The compositional variation of

epidote in metasediments is similar to that of epidote

in the mica-bearing gneisses (Table 8). In the gneisses,

allanite cores rimmed by epidote may constitute an

accessory.

7.4. Titanite

Titanite occurs as very fine-grained crystals with an

anhedral habit. In general, microprobe analyses indi-

cate that titanite contains an appreciable amount of Al

and F. The average contents of (Al + Fe3 +) in titanite

from the gneisses, metasediments, and tourmalinites

are 0.235, 0.162, and 0.179 apfu, respectively (Table

7). The F content is higher in titanites from gneisses

(av. 0.189 apfu) than metasediments (av. 0.118 apfu)

Table 7

Average compositions of garnet (Grt) and titanite (Ttn)

Gneisses Metasediments Tourmalinites

Grt (n= 14) Ttn (n= 8) Grt (n= 78) Ttn (n= 9) Grt (n= 8) Ttn (n= 6)

SiO2 37.62 (0.35) 31.03 (0.58) 36.44 (1.08) 29.58 (0.70) 36.72 (0.25) 30.02 (0.67)

TiO2 0.16 (0.07) 31.79 (2.62) 0.09 (0.08) 35.70 (0.69) 0.10 (0.04) 34.97 (2.79)

Al2O3 21.23 (0.56) 5.48 (1.75) 21.20 (0.39) 3.30 (0.38) 21.07 (0.07) 4.09 (1.57)

Cr2O3 0.02 (0.02) 0.02 (0.03) 0.02 (0.03) 0.01 (0.01) 0.02 (0.02) 0.00 (0.00)

Fe2O3 1.16 (0.49) 1.45 (1.06) 0.99 (0.29)

FeO 23.58 (5.38) 33.48 (1.88) 33.47 (1.73)

MnO 2.13 (1.10) 0.02 (0.03) 0.78 (0.62) 0.01 (0.01) 1.09 (0.87) 0.02 (0.02)

MgO 0.55 (0.67) 0.03 (0.03) 2.12 (1.27) 0.01 (0.01) 0.83 (0.09) 0.00 (0.00)

CaO 14.48 (4.97) 27.58 (0.49) 5.55 (2.27) 27.63 (0.30) 6.91 (0.94) 27.69 (0.75)

Na2O 0.03 (0.02) 0.04 (0.01) 0.03 (0.02)

K2O 0.09 (0.16) 0.04 (0.03) 0.02 (0.01)

F 1.86 (0.68) 1.14 (0.14) 1.32 (0.71)

OMF 0.79 (0.29) 0.48 (0.06) 0.56 (0.30)

Total 99.76 (0.83) 98.30 (0.62) 99.67 (1.30) 98.42 (1.46) 100.20 (0.24) 98.59 (0.83)

Si 2.992 (0.010) 0.997 (0.016) 2.949 (0.038) 0.963 (0.014) 2.969 (0.010) 0.970 (0.021)

Al 1.991 (0.063) 0.207 (0.066) 2.022 (0.028) 0.127 (0.013) 2.008 (0.011) 0.155 (0.059)

Ti 0.009 (0.004) 0.768 (0.063) 0.006 (0.005) 0.875 (0.029) 0.006 (0.002) 0.850 (0.070)

Cr 0.001 (0.001) 0.000 (0.000) 0.001 (0.002) 0.000 (0.000) 0.001 (0.000) 0.000 (0.000)

Fe3 + 0.028 (0.012) 0.035 (0.025) 0.024 (0.007)

Fe2 + 1.570 (0.368) 2.267 (0.138) 2.264 (0.123)

Mn 0.143 (0.073) 0.001 (0.001) 0.054 (0.043) 0.000 (0.000) 0.074 (0.060) 0.000 (0.000)

Mg 0.065 (0.080) 0.001 (0.001) 0.255 (0.149) 0.000 (0.000) 0.100 (0.012) 0.000 (0.000)

Ca 1.231 (0.418) 0.949 (0.020) 0.482 (0.197) 0.964 (0.022) 0.599 (0.080) 0.959 (0.026)

Na 0.002 (0.001) 0.002 (0.001) 0.002 (0.001)

K 0.004 (0.006) 0.002 (0.001) 0.001 (0.001)

F 0.189 (0.069) 0.118 (0.014) 0.135 (0.072)

Structural formula of garnet on the basis of 12 O. Structural formula of titanite calculated after Oberti et al. (1991).

Numbers in parenthesis are standard deviation.


and tourmalinites (av. 0.135 apfu). Al and F show a

good positive correlation. Although some authors have

suggested that pressure favours the substitution

AlFTi� 1O� 1 (Enami et al., 1993), Al-rich titanites

occur also in low-pressure terranes. Fluid composition

(H2O/HF) and bulk chemistry appear to be the impor-

tant factors in determining the formation of Al-rich

titanites (Gibert et al., 1990; Markl and Piazolo, 1999).

7.5. Feldspars

Feldspars are major components of the gneisses in

which K-feldspar shows very low contents of Na2O

(0.63–0.85) and albitic plagioclase has CaO between

0.1 and 0.5 wt.%.

8. Discussion

The origin of tourmalinites has been ascribed to one

of the following processes: (1) precipitation from

exhalative fluids or colloids (Ethier and Campbell,

1977; Slack, 1982; Slack et al., 1984; Plimer, 1988)

(endo-syngenetic model); (2) premetamorphic hydro-

thermal replacement of aluminous sediments or vol-

canics (Slack et al., 1993); (3) diagenesis and meta-

morphism of B-rich sediments or evaporites (Abraham

et al., 1972; Slack et al., 1984); (4) metasomatism by

B-rich fluids of magmatic or metamorphic affiliation

(Appleby and Williams, 1988; Steven and Moore,

1995; Torres-Ruiz et al., 1996). Other interpretations

(e.g., Slack and Coad, 1989; Pesquera and Velasco,

1997) involve a combination of processes that can be

grouped within these models.

The petrographic, mineralogical, and geochemical

data support a premetamorphic metasomatic model for

the Sierra Nevada tourmalinites. However the dis-

cordant reaction rims of the tourmaline are attributed

to metamorphic processes.

8.1. Textural relationships and compositions

Based on textural and compositional data, tourma-

linites and associated rocks show distinct generations

of tourmaline developed during pre-metamorphic and

metamorphic stages.

Tourmaline porphyroclasts in gneisses exhibiting

textural relationships similar to feldspar porphyro-

clasts suggest that tourmaline development began

prior to regional metamorphism. The presence, in

places, of micas, epidote and garnet embayed by

tourmaline is related to recrystallization processes of

preexisting tourmaline during regional metamor-

phism. In addition, the compositional range of the

metamorphic tourmaline is more restricted than that of

the premetamorphic tourmaline (Figs. 7–9).

Abrupt, multiple, fine-scale chemical zoning in the

tourmaline is probably a growth zoning developed

during premetamorphic hydrothermal processes, rather

than recrystallization of crystalloblasts in a boron-rich

sediment or a metasomatic effect induced by infiltra-

tion of fluids during regional metamorphism. How-

ever, textural relationships are not consistent with the

fact of that oscillatory zoning in tourmaline is a

metasomatic growth feature developed during meta-

morphism, as Yardley et al. (1991) claimed for some

metamorphic minerals. According to Taylor and Slack

(1984), formation of multiple growth zones in tourma-

line is indicative of a high fluid–rock environment. We

consider this zoning and overall chemistry in the Sierra

Table 8

Average compositions of epidote

Gneisses

(TFG), n= 8

Gneisses

(TFMG), n= 10

Metasediments,

n= 9

SiO2 39.76 (0.34) 38.48 (0.28) 38.21 (0.68)

TiO2 0.01 (0.01) 0.09 (0.03) 0.13 (0.05)

Al2O3 32.75 (0.51) 27.21 (1.02) 27.82 (1.74)

Cr2O3 0.00 (0.00) 0.01 (0.01) 0.07 (0.06)

Fe2O3 1.41 (0.32) 9.20 (1.34) 8.28 (2.38)

MnO 0.03 (0.02) 0.09 (0.04) 0.08 (0.04)

MgO 0.00 (0.00) 0.01 (0.01) 0.03 (0.01)

CaO 24.13 (0.27) 22.69 (0.44) 23.26 (0.35)

Total 98.10 (0.70) 97.77 (0.51) 97.87 (0.91)

Si 3.020 (0.010) 3.019 (0.009) 2.985 (0.023)

Al 2.933 (0.029) 2.516 (0.074) 2.561 (0.142)

Ti 0.001 (0.001) 0.005 (0.002) 0.007 (0.003)

Cr 0.000 (0.000) 0.000 (0.001) 0.004 (0.004)

Fe3 + 0.080 (0.019) 0.544 (0.083) 0.488 (0.146)

Mn 0.002 (0.002) 0.006 (0.002) 0.005 (0.002)

Mg 0.000 (0.000) 0.002 (0.001) 0.003 (0.001)

Ca 1.964 (0.024) 1.907 (0.020) 1.947 (0.011)

Clinozoisite 97.270 (0.60) 82.070 (2.66) 83.870 (4.78)

Epidote 2.670 (0.61) 17.730 (2.65) 15.960 (4.73)

Piemontite 0.070 (0.05) 0.200 (0.08) 0.170 (0.08)

Structural formula of epidote on the basis of eight cations and

charge balance.

Numbers in parenthesis are standard deviation.


Nevada tourmaline to reflect discontinuous variations

in the composition of constituents attached to the

growing crystal layers through reaction-transport feed-

back. The presence and relative abundance of tourma-

line porphyroclasts in gneisses displaying multiple

oscillatory zoning may represent previous tourmaline

that filled pockets in rhyolite or granite. Occurrences

of tourmaline nodular segregations in granitic rocks

can be explained as the result of: (1) postmagmatic

replacement related to hydrothermal alteration of crys-

tallized granite (Nemec, 1975; Rozendaal and Bruwer,

1995); (2) magmatic-hydrothermal processes related to

exsolution of aqueous, B-rich fluids from saturated

granitic magma (Sinclair and Richardson, 1992; Sam-

son and Sinclair, 1992). The oscillatory zonation can

be explained by repeated changes in crystal–fluid

equilibrium due to sudden changes of temperature,

pressure, H2O saturation, and fluid composition, or on

the basis of disequilibrium crystal growth. The multi-

ple chemical zoning in this case possibly reflects rapid

disequilibrium growth of the tourmaline. Disequili-

brium can lead to autocatalytic processes that increase

the potential for oscillatory behaviour during crystal

growth (Ortoleva et al., 1987). The relatively homoge-

neous porphyroclasts, nevertheless, could be inherited

magmatic tourmaline.

The cellular textures, patchy zoning, and pale dis-

cordant reaction rims are interpreted as metamorphic

features. Zoning irregularities in fine-grained tourma-

line from tourmalinites may be due to coalescence of a

number of small tourmaline crystals followed by

relatively smooth or discontinuous growth around

the coalesced center. These textural features may have

formed via fluid–solid reactions such as Slack and

Coad (1989) proposed for pale reaction rims in the

Kidd Creek massive sulfide deposit. In tourmalines

from Sierra Nevada, the regional metamorphism

caused significant effects on the tourmaline. Evidence

is from the cellularity of many tourmaline crystals,

abundance and volume of reaction rims, and partial

loss of crystal integrity (Fig. 3B–F). The breakdown

of tourmaline may reflect the activity of alkaline and

boron undersaturated fluids rather than variations in

P–T conditions during metamorphism (Weisbrod et

al., 1986; Morgan and London, 1989).

Abundance of cellular textures in the fine-grained

tourmalines relative to the coarse-grained tourmalines

denotes a greater effect of dissolution–diffusion pro-

cesses on the former. Such processes are related to

transmissivity of the rock to fluids. The leaching

characteristics are determined by several factors

including structural properties and the reactivity of

different sites near the mineral surface (Casey and

Bunker, 1990). It is obvious that a fine-grained rock

will have a higher density of reactivity sites than a

coarse-grained rock because its total surface area of

mineral grains is greater. On the other hand, most

diffusion effects will occur in the rim of the tourma-

line because this is where the strongest concentration

gradients occur, although for small crystals, diffusion

may penetrate to interior zones. This means that fine-

grained tourmaline will be more prone to dissolution–

diffusion effects, and consequently, to the develop-

ment of cellular textures. Since diffusion away from

the dissolution front can be inhibited by partial iso-

lation from the ambient fluid (Hibbard, 1995), the

dissolution process will proceed depending on the

ability of the fluids to remove reactive constituents

from the dissolution front. Dissolution of tourmaline

will cease when the solute cannot escape from the

dissolution cell and saturation is reached with subse-

quent precipitation. Removal of solute from the dis-

solution front may proceed by diffusion through

structural defects and intracrystal compositional boun-

daries, but in a tectonically active environment, this

can be achieved by fracturing of crystals and turbu-

lence of fluids.

Tourmaline affected by dissolution processes shows

a variety of elements with different diffusion profiles

because diffusion rates vary. Mostly, the fluids have

leached Al, Fe, Ca, and Ti with concomitant forma-

tion of tourmaline enriched in Mg and Na (Figs.

3B,D,E,F and 11C). The leaching-reaction process

mainly involved ion-exchange mechanisms with par-

allel coupled substitutions. Occurrence of other com-

positional trends, even at thin section scale, is be-

lieved to reflect complex fluid–mineral reactions

controlled by local paragenetic variations. The over-

growths that developed on tourmaline porphyroclasts

in the gneisses illustrate the relation between compo-

sitional zoning patterns and paragenesis: overgrowths

associated with tourmaline porphyroclasts in mica-

free gneisses are relatively richer in Fe and Na,

whereas those in mica-bearing gneisses are richer in

Mg and Ca (Fig. 10). It appears from the chemical

and textural differences between porphyroclasts and


overgrowths that these tourmalines record an ‘‘uncon-

formity’’ corresponding to a transition from hydro-

thermal metasomatism to regional metamorphism.

All of the textural features described above, result-

ing from post-growth dissolution–diffusion processes,

formed during the second metamorphic phase under

amphibolite facies conditions. Support for this inter-

pretation derives from the systematic change in the

Mg/Fe ratio of biotite and muscovite developed during

the second metamorphic phase, with which the pale

reaction rims and cellular tourmaline appear to be in

equilibrium (Fig. 12). It has been observed that tour-

maline rims show a systematic compositional variation

with biotite and muscovite in rocks (e.g., Henry and

Guidotti, 1985). Such a compositional relation is

consistent with temperatures attained during the sec-

ond metamorphic episode in the Sierra Nevada area.

Tourmaline–biotite partitioning data (Fig. 12) are

consistent with an approximate temperature of 550

jC, using the expressions of Colopietro and Friberg

(1987) and Henry and Dutrow (1996).

Textural and compositional relationships suggest a

detrital origin for Cr-free turmaline cores in the

metasediments. However, Cr-bearing orange tourma-

line cores occurring in some of the metasediments

may represent tourmaline formed during the first

metamorphic phase. Thin layers of marbles and/or

calc-schists containing Cr-rich silicates have been

reported in similar stratigraphic sequences from the

Nevado–Filabride Complex. Such Cr-rich minerals

are believed to have formed through reactions involv-

ing detrital Cr-rich spinel during the first metamorphic

phase (Lopez Sanchez-Vizcaıno et al., 1995).

8.2. Boron metasomatism

Tourmalinites from the Sierra Nevada are essen-

tially composed by tourmaline and quartz. It is

unlikely that a rock with subequal amounts of tourma-

line and quartz is a clastic sediment affected by

isochemical metamorphism, even assuming an alumi-

nous pelite with a high boron content (e.g., Slack et

al., 1984). In addition, the wide variation in tourma-

line compositions (Fig. 7) suggests a lack of simple

chemical control by whole-rock chemistry.

The amount of boron in psammopelitic rocks varies

typically from a few tens to 1000 ppm (Harder,

1975a,b; Leeman and Sisson, 1996). Under closed-

system conditions, the boron required for tourmaline

formation in psammopelitic sediments derives from

preexisting tourmaline and/or can be released by the

breakdown of B-bearing minerals during prograde

metamorphism. In metamorphic materials, micas rep-

resent an important boron source (Christ and Harder,

1978; Nabelek et al., 1990; Bebout et al., 1992). In the

schists of Pelona (southern California), for example,

Moran (1993) showed that over 60% of the whole-rock

boron resides in abundant phengite that coexists with

minor tourmaline. It appears that the boron content

fixed in tourmaline increases significantly at upper

greenschist facies (Sperlich et al., 1996), so that in the

amphibolite facies, most the boron are accommodated

in tourmaline. If the tourmalinites (z 15–20% tour-

maline) developed from psammopelitic rocks via dia-

genetic and metamorphic processes, it is obvious that

their formation would need an additional amount of

boron relative to that contained in the sedimentary

protolith, i.e., an external supply of boron is required.

Boron-rich hydrothermal fluids exsolved from a

magmatic source and expelled into the surrounding

rocks is a likely source on the basis of the following

considerations: (1) tourmalinites from the Sierra

Nevada alternate with tourmaline-bearing metasedi-

ments and felsic orthogneisses (Fig. 1). (2) Modal

volume of tourmaline decreases away from the

gneisses. (3) Occurrence of unzoned tourmaline por-

phyroclasts in the gneisses suggests that the magma

acquired enough boron to crystallize tourmaline. Dis-

seminated tourmaline that lacks fine-scale chemical

zonation is the most compelling evidence for boron

enrichment in magma (London et al., 1996). (4)

Locations where late-stage hydrothermal/magmatic

fluids are enriched in boron, tourmalinite may form.

Most wallrocks are reactive to boron-rich fluids to

precipitate tourmaline or other boron minerals (Mor-

gan and London, 1987). When a magmatic system

becomes open to Al-bearing wall rocks, the interac-

tion between late-magmatic fluids and host rocks

generates a chemical–potential gradient that induces

the transport of boron via aqueous fluids from melt to

surrounding rocks. Indirect evidence of the activity of

an important vapor phase is provided by mass balance

calculations, suggesting that tourmalinites of mag-

matic affinity formed under relatively high fluid/rock

ratios. (5) Similarity in the chemical trends shown by

tourmalinites and metasediments of the Sierra Nevada


is consistent with psammopelitic precursors for the

tourmalinites. Psammopelitic metasediments contain

enough Al, Mg, Fe, Ca, and Na to form tourmaline

through reaction of boron-rich fluids with feldspar,

phyllosilicates and other minerals (Morgan and Lon-

don, 1989; Fuchs and Lagache, 1994).

The problem, however, is in determining the boron

source; whether the original source of the boron is

magmatic, that is, the boron evolved specifically from

a silicic magma, or whether it was in materials from the

original sedimentary sequence. In this regard, there are

certain aspects that deserve to be considered. (1) The

Sierra Nevada orthogneisses have a felsic, peralumi-

nous character and, based on Y and Nb contents and

other geochemical data, can be related to continental

collisional environments and crustal melting pro-

cesses. Most boron-enriched melts are siliceous and

peraluminous in nature, and typically are generated as

products of anatexis in continental collision settings

(London et al., 1996; London, 1999). (2) B/Be ratios in

orthogneisses are significantly higher than typical

crustal values (V 6, Leeman et al., 1992), thus denot-

ing relative boron enrichment. Granites or rhyolites

with high B/Be ratios have in common the fact that B-

rich silicic melts formed in areas characterized by

significant crustal thickness (Leeman and Sisson,

1996). The boron content of melts will depend on the

ability that boron-bearing minerals have to retain the

boron until anatexis. If the magma formed by melting

of psammopelitic metasediments that have not under-

gone prior dehydration or melting, the amount of boron

will be relatively high in the initial melts (Leeman and

Sisson, 1996). Low degrees of melting ( < 10%) of

relatively boron-rich protoliths (f 100 ppm) can gen-

erate high boron melts (f 1000 ppm) assuming that

all available boron goes into melts (London et al.,

1996). Such a quantity of boron is nevertheless lower

than that observed in most granites and rhyolites. The

origin of boron-rich magmas depends on the presence

and stability of boron-rich phases at their source and/or

extensive fractionation of magmatic bodies (London et

al., 1996). Boron enrichments also may occur during

emplacement of magma by contamination or hydro-

thermal leaching of boron-rich sequences. (3) Tourma-

line in the orthogneisses has 87Sr/86Sr ratios (c 0.712)

comparable to the initial 87Sr/86Sr value (0.714 that

was regarded by Nieto (1996) to be representative of

orthogneisses from the Mulhacen Complex. This sim-

ilarity suggests that Sr in tourmaline was derived from

the same source as Sr in the gneisses. Most peralumi-

nous igneous rocks have radiogenic initial Sr isotopic

ratios that reflect derivation from a crustal source

(London et al., 1996). (4) The higher 87Sr/86Sr ratios

of tourmalines in the metasediments and tourmalinites

may be due to isotopic exchange during the interaction

of B-rich fluids with wall rocks. By this model, the87Sr/86Sr ratios of the tourmaline would be inherited

from the magmatic-hydrothermal fluids, which in turn

acquired a more radiogenic character by infiltration

into the surrounding rocks.

If the boron-rich hydrothermal fluids had a mag-

matic source, an enrichment of tourmalinites in Li and

Be might be expected. The abundant Be in tourma-

linites from the Orobic Alps was interpreted by Slack

et al. (1996) to indicate that the tourmalinizing fluids

had a felsic igneous component. In the Sierra Nevada

area, nevertheless, Be contents are lower than in the

gneisses (Table 3). This is not consistent with a

magmatic provenance, but it may be also due to a

low mobility during metasomatic processes. Be has

much in common chemically with Al, with similar

geochemical behaviour. The boron may have been

also introduced into the magma externally from the

sedimentary sequence, particularly by hydrothermal

leaching of the metaevaporitic formation located in

the cover from the Nevado–Filabride Complex. This

cover underlies the calc-schist formation that includes

the gneisses and tourmaline-rich rocks. Such evapor-

ites could represent a potential additional source of

abundant boron for tourmaline formation, and perhaps

a source of other elements like F and Li. With the

present data, however, it is very difficult to assess the

role that the evaporites may have played in forming of

Sierra Nevada tourmalinites.

9. Summary

The principal features and implications of our

study are as follows:

1. In the Sierra Nevada area, different stages of

tourmaline formation can be established: (1)

tourmaline porphyroclasts of magmatic to post-

magmatic affiliation within the orthogneiss; (2) fine-

grained tourmaline formed by metasomatic replace-


ment of psammopelitic rocks giving rise to

tourmalinites; (3) textural and compositional read-

justments due to the Alpine regional metamorphism.

2. Tourmaline coexists with quartz, feldspars, musco-

vite, biotite, garnet, epidote, and Al-rich titanite.

Except for porphyroclasts, tourmaline is very fine

to fine grained, variable in colour (green colours

predominate), with many crystals displaying a fine-

scale chemical zonation that is interpreted to be of

premetamorphic origin. Tourmaline crystals com-

monly exhibit textural features (e.g., cellular

morphologies, pale discordant reaction rims) that

denote some destabilization during regional meta-

morphism, probably due to the activity of alkaline

and boron undersaturated fluids.

3. Overall, tourmaline shows a wide compositional

variation along the schorl–dravite solid solution

series, with variable proportions of X-site vacan-

cies, Al deviation, and minor amounts of Ca. The

range of Fe/(Fe +Mg) ratios for tourmalines from

the tourmalinites (0.34–0.95) is more limited than

for those from the metasediments (0.16–0.92) and

gneisses (0.28–0.97). However, there are no

differences in the Na/(Na + Ca) ratios, which

generally vary between 0.5 and 0.9. Tourmaline

that formed during regional metamorphism shows

a more restricted compositional field, being

characterized by increases of Mg and Na, and

decreases of Al, Fe, Ca, and Ti. The composition of

tourmaline overgrowths varies depending on the

mineral paragenesis: in contrast to tourmaline

overgrowths from the mica-free gneisses, the

composition of those from mica-bearing gneisses

is enriched in Mg relative to the porphyroclasts.

4. Tourmaline typically shows three main types of

compositional zonation: (1) a discontinuous core-

to-rim zoning wherein detrital cores (and Cr-

bearing cores probably developed during the first

metamorphic phase) are surrounded by tourmaline

overgrowths of different composition that formed

during the second metamorphic phase; (2) tourma-

line porphyroclasts exhibit abrupt, multiple, oscil-

latory chemical and colour zoning which are

surrounded by metamorphic tourmaline over-

growths; the composition of these overgrowths

being paragenetically controlled. Some composi-

tionally homogeneous porphyroclasts could be

inherited magmatic tourmaline. (3) Fine-scale

chemical zonation, commonly patchy in the

interiors and including pale discordant reaction

rims, is characterized by significant fluctuations in

Al, Fe, Mg, Na, Ca, Ti, and F. Fluid mixing in an

open system is believed to be the cause of such

chemical fluctuations.

5. Petrographic and geochemical data are consistent

with a psammopelitic precursor for the Sierra

Nevada tourmalinites. Chemical variations involv-

ing relatively immobile elements (Al, Ti, Zr, Nb,

etc.) show that the metasediments and tourmalin-

ites define the same trend, which is clearly different

from that of the gneisses. The REE patterns of the

tourmalinites are also similar to those of the

metasediments.

6. Tourmaline-rich rocks formed as a result of the

metasomatic replacement of psammopelitic rocks

by boron-rich hydrothermal fluids that were

derived from adjacent felsic magma. The source

of boron is uncertain, however. 87Sr/86Sr ratios of

the tourmaline appear to be inherited from

magmatic-hydrothermal fluids that became more

radiogenic by interaction with surrounding rocks.

The metasomatic processes probably took place

under relatively high fluid/rock conditions where

tourmaline composition reflects the interaction of

both magmatic-hydrothermal and sedimentary

systems.

Acknowledgements

We are grateful to J.F. Slack for the helpful

comments that improved the manuscript. We also

thank D.J. Henry for the constructive review. Support

for this study was received from the Spanish CICYT

(project no. PB 98-0150). [RR]

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