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Lithos 81 (2005

Deformation-driven differentiation of granitic magma:

The Stepninsk pluton of the Uralides, Russia

F. Beaa,*, G.B. Fershtaterb, P. Monteroa, V.N. Smirnovb, J.F. Molinaa

aDepartment of Mineralogy and Petrology, Fuentenueva Campus, University of Granada, 18002 Granada, SpainbInstitute of Geology and Geochemistry, Russian Academy of Sciences, Pochtovi per. 7, Ekaterinburg, 620219 Russia

Received 25 April 2004; accepted 29 October 2004

Available online 1 January 2005

Abstract

The wide compositional spectrum of the Variscan batholiths of the Urals, a continuum from gabbros (or diorites) to

leucogranites, was produced by crystal fractionation, but the physical mechanisms involved in formation of the bodies remains

obscure. To test whether syn-magmatic deformation was essential to enhance the efficiency of the process, we studied the Main

Series of Stepninsk, a pluton emplaced into an active crustal-scale strike-slip shear zone. The Main Series (N95 vol.% of total

granitoids) is high-K calc-alkaline, comprises rocks with SiO2 from 51 to 77 wt.%, and stands out because most major and

many trace elements yield excellent linear or curvilinear correlations with silica. It includes deformed gabbrodiorites to

monzogranites, and undeformed syenogranites to alaskites. Deformed and undeformed rocks are coeval (283F2 Ma). All rocks,

irrespective of their silica content, have the same initial Sr and Nd isotope ratios (87Sr/86Sr283 Ma=0.70488F0.000131; e(Nd)283Ma=�0.79F0.49), and contain amphibole and biotite with the same compositions. Based on thermodynamic and trace-element

fractionation simulations, we propose a model of deformation-driven filter-pressing differentiation consistent with these

features. The Main Series is derived from a hydrous high-K granodioritic magma which intruded containing ~0.3 of early-

formed solids. These accumulated locally by flow differentiation to produce the gabbrodiorites. The crystallization continued

until the fraction of solids was higher than ~0.55, after which different magma batches were efficiently squeezed by differential

stress coupled with the opening of tensile fractures in the shear zone. This process produced a range of residua and segregates,

the composition of which depended on the fraction of early-formed solids, the fraction of solids present when squeezing

occurred and, especially, the efficiency of melt segregation. The monzodiorites and quartz–monzodiorites represent efficiently

squeezed residua, the granodiorites to monzogranites represent unfractionated or little fractionated magma batches, and the

leucogranites represent melt segregates with a few entrained crystals of amphibole and biotite. We proposed that wide-spectrum

fractionation of granite magmas mainly occurs when they crystallize under compressive regimes, and is caused by deformation-

driven filter-pressing differentiation.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Granitoids; Magmatic differentiation; Filter pressing; Geochemical modeling

0024-4937/$ - s

doi:10.1016/j.lit

* Correspondi

E-mail addr

) 209–233

ee front matter D 2004 Elsevier B.V. All rights reserved.

hos.2004.10.004

ng author. Tel.: +34 958246176; fax: +34 958243368.

ess: fbea@ugr.es (F. Bea).

F. Bea et al. / Lithos 81 (2005) 209–233210

1. Introduction

One of the most outstanding features of the Uralides

orogen is the wide compositional spectrum of granite

batholiths. Despite their modest size (Fig. 1) almost

every pluton comprises a great variety of rocks, a

continuum from gabbros or diorites to leucogranites

Fig. 1. Geological sketch of Stepninsk pluton. Based on R. N. Shagina

elaborated by the Chelyabinsk Geological Survey (personal communicatio

(Fershtater, 1984; Fershtater et al., 1994; Bea et al.,

2002; Gerdes et al., 2002). Though this notable

geochemical diversity might have resulted from differ-

ent factors, compelling evidence firmly points to

crystal fractionation as the main cause of differentiation

(Fershtater, 1987). The way whereby this occurred,

however, remains obscure. The high viscosity of

(in Fershtater et al., 1994) and the magnetic and gravimetric maps

n).

F. Bea et al. / Lithos 81 (2005) 209–233 211

parental melts, granodioritic or tonalitic, and the

moderate density contrast between the melt and

precipitating phases—mostly labradorite–andesine,

hornblende and biotite—would have decreased the

efficiency of crystal/melt segregation to the point that

either chemically driven convective diffusion, gravity-

driven crystal settling, or any other conceivable

mechanism of magmatic differentiation involving no

external forces, except gravity, could hardly have

produced such highly fractionated rock series. On the

other hand, the fact that most melanocratic facies of

every differentiated Uralian pluton are commonly

strongly foliated whereas most leucocratic facies are

undeformed suggests that syn-magmatic deformation

might have been essential for enhancing the differ-

entiation mechanism.

To explore this idea, we carried out a detailed study

on Stepninsk, a composite pluton syn-kinematically

emplaced into a crustal-scale strike-slip shear zone

(Fershtater et al., 1994; Bea et al., 2002). Using Rb–Sr

and single-zircon stepwise-evaporation 207Pb/206Pb

methods, we first show that the deformed gabbrodior-

ites to granodiorites, and the undeformed granites and

leucogranites are coeval. Then, using major and trace-

element data, and Sr and Nd isotopes we show that

most of them are cogenetic and derived from a high-K

granodioritic magma. Whole-rock geochemistry and

major and mineral trace element data indicate that the

initial magma did not differentiated by sequential

fractional crystallization. The combination of thermo-

dynamic modeling using the MELTS software

(Ghiorso and Sack, 1995; Asimow and Ghiorso,

1998) with trace-element fractionation equations per-

mitted us to build a model of deformation-enhanced

filter-pressing differentiation which explains many

features of the orogenic granitoids of the Urals. Finally,

we suggest that deformation-driven filter pressing may

be the dominant fractionation mechanism of granite

magmas crystallizing in compressive regimes.

2. Geological setting and field relationships

The Variscan orogen of the Urals was the locus of

sustained granite magmatism from ~350 to ~250 Ma,

first related to subduction and then to collision (Bea

et al., 2002, and references mentioned therein). In

contrast to the Variscan belt of western Europe, the

rate of formation of granite magmas did not increase

perceptibly after the Upper Carboniferous collision

(Montero et al., 2000), probably because the Uralian

orogen did not undergo any major post-collisional

extensional collapse (Brown et al., 1998). According

to age and field relationships, Bea et al. (2002)

classified the post-collisional granite bodies of the

Urals into three groups: (1) medium-sized batholiths

composed of 300 to 290 Ma granitoids, apparently

unrelated to the main tectonic structures, (2) small-

sized plutons composed of 285 to 275 Ma granitoids,

emplaced into active crustal-scale strike-slip shear

zones, (3) medium-sized batholiths composed of 260

to 250 Ma granitoids, related to migmatitic com-

plexes. The Stepninsk pluton (54800VN, 60820VE),located within a N1608E dextral shear zone (Fig. 1),

probably represents the best example of the second

category.

The outcropping shape of the Stepninsk pluton is

roughly elliptical, with a ~19 km N–S axis and ~14 km

E–W axis (Fig. 1). It has a nearly concentric structure

formed of alternating bands of strongly deformed

granodiorites to gabbrodiorites, undeformed granites

and leucogranites, and numerous septa of country

rocks: metapelites, limestones and metavulcanites

(Figs. 1 and 2). The units of deformed granitoids

comprise dominant granodiorites, monzogranites and

quartz–monzodiorites, and subordinate monzodiorites

and gabbrodiorites (Fig. 2E). The latter occur inter-

mingled with the former as metric to hectometric

patches with convoluted, locally gradational, contacts.

Rock fabrics are planar or plano-linear, and range from

coarse-grained gneissose to fine-grained milonitic (Fig.

2A–D). The main foliation is parallel to the external

contacts and dips to the center of the body, from ~808near the outer contact to ~608 near the center (Fig. 1),defining a funnel-shaped geometry. The stretching

lineation is subhorizontal, and the shearing microfabric

parallel to the outer contact is always dextral (Fig. 1).

The units of undeformed granitoids comprise dominant

leucocratic monzogranites, syenogranites and alas-

kites, hereafter collectively called the leucogranites

(Fig. 2E). They occur either as bent dikes parallel to the

units of deformed granitoids or as scarce unbent

subvertical N140E dikes that cut all other granite units.

The tabular septa of metamorphic rocks make up

about 20% of the exposed surface. Their external

contacts and internal foliation are parallel to the

Fig. 2. (A–D) Field aspects of the Main Series: (A) gabbrodiorites with granodioritic segregates showing dextral shear folds. (B) Leucogranite

segregates in a milonitic gabbrodiorite; the segregates are accumulated in shear folds and may be up to meter-sized. (C) Sheared granodiorite

with diffuse patches of undeformed monzogranite and veins of leucogranite. (D) Deformed granodiorites with abundant enclaves of

gabbrodiorites cut by an undeformed granite emplaced into a tensile shear zone; note the bending of mafic streaks at the contact; the undeformed

granite still has tenuously defined schlierens of micas parallel to the main foliation. These textures indicate that the deformation ceased before

the total crystallization of the magma and was not recorded by the magmas with the highest melt fraction. (E) Q–A–P plot of Stepninsk rocks.

Note how the deformed are the less silicic varieties. Arrows represent the paths followed by simulations made with the MELTS software

(Ghiorso and Sack, 1995; Asimow and Ghiorso, 1998) for the average composition of the batholith with different percentages of H2O,

represented by the numbers associated with each arrow. Note how the trend is close to the simulation with 4.5 wt.% H2O.

F. Bea et al. / Lithos 81 (2005) 209–233212

foliation of the host granites. Both the septa and the

host rocks show a marked contact metamorphism. The

mineral assemblage of contact metapelites consists of

quartz, plagioclase (An18-40), biotite, fibrolite, and rare

garnet and cordierite. Our geobarometric estima-

tions indicate they were equilibrated at 550–580 8C(biotite/garnet geothermometer) and ~5 kbar (GASP

geobarometer).

According to field relationships, it seems that the

oldest rocks are the gabbrodiorites and monzonites

F. Bea et al. / Lithos 81 (2005) 209–233 213

followed first by the granodiorites and deformed

monzogranites, then by the undeformed monzogran-

ites and leucogranites of bent dikes and, at last, by the

leucogranites of unbent dikes. At first glance, it might

seem that the emplacement and crystallization of this

rock sequence should have required a considerable

time span. Radiometric dating, however, indicates

they are coeval (see below).

To discuss in detail the emplacement mechanism of

the pluton is beyond the scope of this paper. For our

purposes, it is enough to consider that all available

geophysical and structural evidence (R. N. Shagina

and the Chelyabinsk Geological Survey, personal

communication of unpublished data) indicates that

the pluton grew by diking, hence the numerous septa

of country rocks and the tabular shape of most

individual granite units, and acquired its roughly

concentric structure by syn-emplacement rotation

within the strike-slip crustal-scale shear zone. This is

supported by the subhorizontal stretching lineation

and widespread, consistently dextral, shearing micro-

fabric of granitoids parallel to the pluton’s external

contact all around its perimeter.

3. Petrography of granitoids

According to major and trace elements and Nd

isotopes (Figs. 4–7), Stepninsk granitoids may be

clearly divided into two groups, the Main Series and

the Minor Gabbros. The Main Series is comprised of

both deformed and undeformed rocks and makes up

the bulk of the body; the Minor Gabbros are limited to

a couple of neighboring outcrops in the middle of the

pluton.

The Main Series is composed of gabbrodiorites,

monzodiorites, quartz–monzodiorites, granodiorites,

monzogranites, syenogranites and alaskites (Fig. 2).

All these rocks are composed of plagioclase, K-

feldspar, quartz, hornblende and biotite, accompanied

by conspicuous euhedral crystals of titanite. The

plagioclase from the gabbrodiorites and monzonites

ranges from An51 to An37, and commonly shows

patchy zoning. The plagioclase from the granodiorites

display concentric diffuse oscillatory zoning, from

An43 to An21. The plagioclase from the monzogran-

ites and leucogranites is zoned from An25-20 to An10-5.

K-feldspar is abundant: in the most mafic members

of the series it is limited to anhedral interstitial

crystals; in the granodiorites and granites, it also

appears as euhedral or subhedral porphyritic crystals.

The only ferromagnesian minerals are edenitic

hornblende and biotite. Hornblende dominates over

biotite in gabbrodiorites, but it is subordinate in

granodiorites and is just residual or totally disappears

in the leucogranites which may also contain a few

flakes of late-magmatic to secondary muscovite.

Titanite is ubiquitous; it appears as euhedral crystal

of variable size, up to 1–2 mm, commonly with

abundant inclusions of apatite, ilmenite and magnet-

ite; its modal abundance may reach 2% in the

monzodiorites and granodiorites, decreasing with

increasing silica. Besides titanite, all members of

the Main Series contain abundant apatite, zircon,

magnetite and minor ilmenite. Scarce crystals of Fe-

sulfides, monazite and cheralite appear in rocks with

SiO2z62%. Xenotime, which is very rare, appears

limited to rocks with SiO2N74%. Primary-looking

epidote and allanite are extremely rare and confined

to monzogranites. Secondary allanite is also found

locally as anhedral masses replacing monazite.

The Minor Gabbros are all deformed. Their

mineral assemblage comprises edenitic hornblende

and plagioclase (An54-42) with minor biotite and very

scarce clinopyroxene. As accessories they contain

titanite, apatite, zircon and abundant magnetite and

Ti-magnetite.

4. Samples and methods

For this study, we used 27 samples representative

of the Main Series, 2 Minor Gabbros, and 3

metapelites. All of them were analyzed for major

and 38 trace elements, including 14 REE, Th and U.

Eleven samples of the Main Series, the two Minor

Gabbros and the three metapelites were also analyzed

for Sr and Nd isotopes. Thin sections of the whole

collection were studied by optical and scanning

electron microscopy with backscattered electrons (to

identify accessories) and electron microprobe. Eight

thin-sections of the Main Series were also studied by

laser-ablation ICP-MS to determine the trace-element

composition of ferromagnesian minerals. Zircons

from two samples of the Main Series: a strongly

deformed quartz–monzodiorite and an undeformed

Fig. 3. (A) Rb–Sr isochron of Stepninsk fitted assuming that al

scatter is analytical (model 1 of York, 1969). (B) Single zircon207Pb/206Pb ages of zircons. Note how deformed and undeformed

rocks yield exactly the same age.

F. Bea et al. / Lithos 81 (2005) 209–233214

granite were separated for single grain evaporation207Pb/206Pb analysis.

Major-element and Zr determinations were per-

formed by X-ray fluorescence after fusion with

lithium tetraborate. Typical precision was better than

F1.5% for an analyte concentration of 10 wt.%, and

F5% for 100 ppm Zr. Trace elements, except Zr and

Hf, were determined by ICP-mass spectrometry (ICP-

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

powder in a Teflon-lined vessel at ~180 8C and ~200

psi for 30 min, evaporation to dryness, and subsequent

dissolution in 100 ml of 4 vol.% HNO3; the precision

was better than F5% for analyte concentrations of 10

ppm. Samples for Sr and Nd isotope analysis were

digested in the same way using ultraclean reagents

and analyzed by thermal ionization mass spectrometry

(TIMS) in a Finnigan Mat 262 spectrometer after

chromatographic separation with ion-exchange resins.

Normalization values were 86Sr/88Sr=0.1194 and146Nd/144Nd=0.7219. Blanks were 0.6 and 0.09 ng

for Sr and Nd. The external precision (2s), estimated

by analyzing 10 replicates of the standard WS-E

(Govindaraju et al., 1994), was better than F0.003%

for 87Sr/86Sr and F0.0015% for 143Nd/144Nd.87Sr/86Rb and 143Sm/144Nd were directly determined

by ICP-MS following the method developed by

Montero and Bea (1998), with a precision better than

F1.2% and F0.9% (2s), respectively.

Zircons were conventionally separated using mag-

netic separators and heavy liquids. Representative

grains were dated with the 207Pb/206Pb stepwise-

evaporation method (Kober, 1986, 1987). Data

acquisition was performed in dynamic mode using a

secondary electron multiplier (SEM) as detector with

the 206–204–206–207–208 mass sequence. The

mass-ratio 204/206 was monitored to detect and, if

necessary, correct for common lead. Factors for

common lead correction were calculated by iteration

from the 204Pb/206Pb and 204Pb/207Pb ratios provided

by the Stacey and Kramers (1975) model at the

calculated age, until convergence to a constant value.

Mass fractionation in the detector was corrected by

multiplying by M(207/206). Standard errors for each

step were calculated according to the formula:

S.E.=2s/Mn. However, the 2s confidence interval for

the final age is given by (X�t (0.025)s /Mn ,

X+t(0.025)s/Mn), where X and s are the average

and the standard deviation of measured steps, n the

number of steps, and t(0.025) is the upper (0.025)

point of the t distribution for n�1 degrees of freedom.

Major-element analyses of minerals were obtained

by wavelength dispersive analyses with a CAMECA

electron microprobe. Accelerating voltage was 20 kV

and beam current was 20 nA. The precision was close

to F4% for an analyte concentration of 1 wt.%. LA-

ICP-MS analyses were performed with a 213 nm

Mercantek Nd-YAG laser at coupled to an Agilent

7500 ICP-MS with a shielded plasma torch, using the

NIST-610 glass as standard. The ablation was carried

out in an He atmosphere. The laser beam was fixed to

a 95 Am wide square section. The spot was pre-

ablated for 45 s using a laser repetition rate of 10 Hz

and 40% output energy. Then the spot was ablated for

60 s at 10 Hz with a laser output energy of 75%. To

keep the laser focused during ablation, the sample

stage was set to move upwards 5 Am every 20 s. A

typical session of analysis of a single thin section

l

F. Bea et al. / Lithos 81 (2005) 209–233 215

began and ended with the analysis of the NIST-610

glass (about 450 ppb of each element), which was

also analyzed every nine spots to correct for drift.

Silicon was used as internal standard. Data reduction

was carried out with a homemade software imple-

mented as an ado file (freeware available from F.

Bea) of the STATA commercial package. This

software permits identification and discarding out-

liers, blank subtraction, drift correction, internal

standard correction and conversion to concentration

units. The precision, calculated on the five to seven

replicates of the NIST-610 measured in every session,

is in the range F3% to F7% for most elements.

Whole-rock analysis of Hf was also done by LA-ICP-

MS on the fused disks used from XRF analysis of

Table 1

Rb–Sr and Sm–Nd isotope composition of selected samples

Sample MG

STB-6 STB-39 ST-42 ST-40

Rb (ppm) 10.7 16.3 73.9 111

Sr (ppm) 417 423 1392 169487Rb/86Sr 0.074 0.111 0.154 0.1987Sr/86Sr 0.705497 0.705805 0.705638 0.70587Sr/86Sr (t) 0.705199 0.705357 0.70502 0.704

e(Sr)t 14.6 16.9 12.1 10.9

Nd (ppm) 12.8 12.9 68.3 74.9

Sm (ppm) 3.14 3.06 10.51 10.64147Sm/144Nd 0.148 0.1439 0.093 0.085143Nd/144Nd 0.512857 0.512865 0.512388 0.512143Nd/144Nd(t) 0.512583 0.512598 0.512216 0.512

e(Nd)t 6.04 6.34 �1.13 �1.2

TC(Nd) Ma 691 633 968 936

Sample MS-D MS-U

STB-11 STB-17 STB-20 STB-

Rb (ppm) 113.5 130.3 200.2 275.9

Sr (ppm) 421 562 219 16087Rb/86Sr 0.78 0.67 2.641 4.98387Sr/86Sr 0.707966 0.707455 0.715331 0.72487Sr/86Sr (t) 0.704824 0.704756 0.704698 0.704

e(Sr)t 9.3 8.4 7.5 7.7

Nd (ppm) 31.1 30.5 15.6 22.1

Sm (ppm) 3.79 4.04 1.69 3.23147Sm/144Nd 0.0735 0.0802 0.0655 0.088143Nd/144Nd 0.512377 0.512439 0.51236 0.512143Nd/144Nd(t) 0.512241 0.512291 0.512239 0.512

e(Nd)t �0.64 0.33 �0.68 �1.1

TC(Nd) Ma 846 817 818 945

The time for calculations is 283 Ma, the estimated crystallization age of th

represents the Minor Gabbros. MS-D and MS-U represent deformed and

metapelites.

major elements and Zr, using this element as internal

standard and the same conditions as for mineral

analysis; the precision was better than F5% for

analyte concentration of 1 ppm.

5. Results

5.1. Radiometric dating

The 11 samples of the Main Series analyzed for Sr

isotopes included 7 deformed rocks, ranging from

gabbrodiorites to granodiorites, and 4 undeformed

rocks, ranging from monzogranites to alaskites.

Despite the seemingly apparent age differences based

MS-D

STB-4 STB-29 STB-5 ST-43

85.1 102 94.4 95.1

1065 1212 977 766

0.231 0.243 0.28 0.359

696 0.705810 0.705854 0.706219 0.706478

933 0.704879 0.704874 0.705093 0.705032

10.1 10 13.2 12.3

79 73.7 56.3 31.5

9.94 10.45 7.21 4.7

9 0.0761 0.0857 0.0774 0.0901

37 0.512383 0.512367 0.512387 0.512419

211 0.512242 0.512208 0.512244 0.512252

2 �0.62 �1.28 �0.59 �0.42

855 938 859 907

MP

10 STB-23 ST-14 STB-32 ST-15

247.6 131.6 80.7 57.7

130 638 130 21.5

5.531 0.597 1.801 7.797

775 0.727136 0.712964 0.722105 0.747933

709 0.704863 0.710559 0.714855 0.716539

9.9 90.8 151.8 175.7

5.7 35.3 26.1 30

0.61 4.81 4.68 6.55

4 0.0639 0.0824 0.1084 0.1319

378 0.512329 0.51211 0.512076 0.512055

214 0.512211 0.511957 0.511875 0.511811

6 �1.23 �6.17 �7.78 �9.04

840 1211 1558 2040

e pluton. TC (Nd) is the Nd model age for a depleted mantle. MG

undeformed granitoids of the Main Series. MP represents the host

F. Bea et al. / Lithos 81 (2005) 209–233216

on field relations, the 11 samples plot as a single Rb–

Sr isochron at 281F4 Ma, with initial 87Sr/86Sr=

0.704933F0.000084 and MSWD=0.45 (Fig. 3A, raw

data in Table 1). The 207Pb/206Pb age of zircons is also

the same: 18 determinations in 6 zircon grains from a

deformed quartz–monzodiorite yielded 283F2 Ma; 10

determinations in three grains from an undeformed

leucogranite yielded 283F3 Ma; it is important to

emphasize that we never detected zircon grains with

inherited cores (Fig. 3B, raw data in Table 2). The

average of both samples is 283F2 Ma, identical

within the error range to the Rb–Sr age; therefore, we

interpret it to be the crystallization age of the body.

These data indicate that all rocks forming the Main

Series are coeval and were emplaced when the shear

Table 2

Single-zircon stepwise-evaporation 207Pb/206Pb data of a deformed quartz

Grain Step 207Pb/206Pb Error (2s)

Deformed quartz–monzodiorite

1 1 0.05223450 0.39

1 2 0.05244037 0.26

2 1 0.05296366 0.30

3 1 0.05227222 0.11

3 2 0.05234211 0.31

3 3 0.05237189 0.10

3 4 0.05221159 0.14

4 1 0.05401003 0.23

4 2 0.05418523 0.16

4 3 0.05756705 0.26

4 4 0.05416777 0.40

5 1 0.05287511 0.30

5 2 0.05219240 0.12

5 3 0.05270589 0.16

6 1 0.05496898 0.29

6 2 0.05304725 0.35

6 3 0.05257099 0.21

6 4 0.05275534 0.33

Undeformed syenogranite

7 1 0.05256540 0.21

7 2 0.05246177 0.18

7 3 0.05231274 0.20

8 1 0.05686312 0.17

8 2 0.05224630 0.24

9 1 0.05329306 0.27

9 2 0.05268398 0.44

9 3 0.05258720 0.42

10 1 0.05246541 0.19

11 1 0.05213494 0.15

Note they yield exactly the same age.

zone was active; they also show that the deformation

ceased when the most felsic members of the series

were still molten and, therefore, did not record the

deformation.

Minor Gabbros present neither enough Rb/Sr

dispersion to build a whole-rock isochron nor contain

suitable zircon crystals for dating with the 207Pb/206Pb

stepwise-evaporation method. Field relationships,

however, reveal they are coeval with the Main Series.

5.2. Chemical and isotopic composition

5.2.1. Main Series

The Main Series is composed of high-K subalka-

line to calc-alkaline rocks (Fig. 4A; Table 3) with

–monzodiorite and an undeformed syenogranite from Stepnisnk

204Pb/206Pb c207Pb/206Pb Age

0.00003780 0.0518097 277.0F9

0.00003811 0.0520116 285.9F6

0.00006813 0.0520985 289.7F7

0.00002818 0.0519879 284.8F3

0.00002451 0.0521116 290.3F7

0.00003774 0.0519484 283.1F2

0.00003027 0.0518967 280.8F3

0.00014734 0.0519923 285.0F5

0.00016808 0.0518651 279.4F4

0.00039726 0.0519141 281.6F6

0.00016493 0.0518937 280.7F9

0.00008504 0.0517628 274.9F7

0.00002644 0.0519332 282.4F3

0.00006485 0.0518877 280.4F3

0.00021540 0.0519611 283.6F6

0.00007991 0.0520105 285.8F8

0.00005250 0.0519327 282.4F5

0.00006501 0.0519351 282.5F8

Average= 283F2

0.00005857 0.0518384 278.2F5

0.00004267 0.0519666 283.9F4

0.00002757 0.0520375 287.0F4

0.00034890 0.0519133 281.5F4

0.00002917 0.0519475 283.0F5

0.00010246 0.0519279 282.2F6

0.00005383 0.0520267 286.5F10

0.00004582 0.0520465 287.4F9

0.00003816 0.0520360 286.9F4

0.00003175 0.0517982 276.5F3

Average= 282F3

Fig. 4. (A) Borodin (1987) plot of Stepninsk granitoids; the parameter Ac (acidity) is calculated as Ac=4Si/(1.5Ti+6Al+4Fe3++5-

Fe2++5Mn+5.5Mg+7.5Ca+17Na+18K) where Si, Ti et cetera are expressed in milliatoms. The Main Series is subalkaline to calc-alkaline, but

the Minor Gabbros are tholeiitic. (B) N-MORB-normalized plots of the gabbrodiorites to granodiorites of the Main Series and Minor Gabbros.

Whereas the former are similar to high-K calc-alkaline, the later are akin to IAT or low-K calc-alkaline magmas. Normalization values are those

of Hofmann (1988).

F. Bea et al. / Lithos 81 (2005) 209–233 217

SiO2 in the range 51–77 wt.% and elevated LIL and

LREE. NMORB-normalized patterns of the less

silicic varieties (Fig. 4B) have a negative anomaly

of Nb and positive anomalies of Sr, Ba and Pb; they

are notably similar to high-K subduction magmas

(e.g., Tatsumi and Eggins, 1995) except for the

presence of small positive anomalies of Zr and Y.

Chondrite-normalized REE patterns of gabbrodior-

ites, monzodiorites and granodiorites are almost

identical, with LaNc300–400, LuNc10 and no Eu

anomaly (Fig. 5). Monzogranites yield parallel

patterns but at lower REE concentration, with

LaNc200 and LuNc5–6. The concentration of

REE reaches a minimum in the most differentiated

leucogranites, which have LaNc70–90, LuNc3–4, a

negative Eu anomaly (Eu/Eu*c0.4–0.5) and a

marked depletion in Er–Ho attributable to titanite

fractionation, according to our LA-ICP-MS analysis

of minerals (Table 4).

All rocks of the Main Series have notably uniform

initial Sr and Nd isotope ratios (Fig. 6; Table 1):87Sr/86Sr283 Ma ranges from 0.704698 to 0.705093,

with an average of 0.70488 and standard deviation of

0.000131; e(Nd)283 Ma ranges from �1.28 to 0.33

with an average of �0.79 and standard deviation of

0.49.

5.2.2. Minor Gabbros

According to the major elements, the Minor

Gabbros are tholeiitic (Fig. 4A). Compared to similar

rocks of the Main Series, they are richer in MgO and

CaO but poorer in Al2O3, TiO2 and, especially, Na2O,

K2O and P2O5 (Fig. 7; Table 3). The trace element

patterns are characteristic of arc magmas, either

tholeiitic or low-K calc-alkaline magmas, and are

notably lower in LIL and LREE than similarly silicic

rocks from the Main Series (Fig. 4B). Chondrite-

normalized REE plots show little LREE/HREE

fractionation (LaN/LuNc3) and no Eu anomalies

(Fig. 5). The initial 87Sr/86Sr (0.705278F0.000119)

is close to the Main Series, but the e(Nd)283 Ma is

sharply different: 6.19F0.21 (Fig. 6). The pronounced

differences in Nd isotopes and major and trace

elements indicate that the Main Series and the Minor

Gabbros are not cogenetic. These, therefore, will not

be considered in the discussion about the mechanisms

of differentiation that produced the Main Series.

5.3. Element variation patterns in the main series

5.3.1. Bulk rock variations

Fig. 7 reveals an excellent linear correlation

(Rb�0.995) of SiO2 with TiO2, FeOtot., CaO and

Table 3

Chemical composition of selected samples of Stepninsk

Ref. STB-39 STB-6 ST-42 ST-40 STB-4 STB-38 STB-29 STB-5 ST-43 STB-17

SiO2 50.73 50.89 51.06 52.95 56.93 59.9 61.01 62.04 64.34 67.2

TiO2 1.18 1.25 1.57 1.49 1.09 0.99 1.01 0.91 0.82 0.63

Al2O3 17.52 18.03 18.51 18.31 17.99 17.23 16.87 16.76 15.48 15.72

FeOtot 8.60 8.64 8.94 8.32 6.78 5.79 5.44 5.19 4.71 3.53

MgO 6.59 6.37 4.43 3.99 3.10 2.45 2.18 2.17 1.77 1.43

MnO 0.16 0.16 0.14 0.13 0.11 0.09 0.08 0.08 0.07 0.05

CaO 10.31 10.73 7.79 6.87 5.81 4.84 4.91 4.16 3.74 2.71

Na2O 2.69 2.59 4.07 3.76 3.75 3.67 3.43 3.66 3.20 3.14

K2O 0.50 0.33 2.05 2.57 2.74 3.68 3.65 3.26 4.18 3.8

P2O5 0.17 0.21 0.79 0.74 0.60 0.52 0.49 0.47 0.41 0.03

Li 19 13 24 26 33 31 31 36 40 31

Rb 16 11 73 111 85 95 102 94 92 130

Cs 0.6 0.5 2 1.8 1.5 1.8 1.8 2.1 1.7 4.6

Be 0.8 0.7 2.4 2.6 3.4 2.9 2 3 2.2 2.5

Sr 423 417 1593 1494 1065 1022 1212 977 766 695

Ba 131 117 1351 1470 1383 1710 1633 1427 1471 1252

Sc 31 29 19 18 14 12 11 11 10 8

V 168 191 185 162 135 111 116 102 91 72

Cr 343 340 62 52 85 66 72 67 48 46

Co 33 33 25 23 19 16 14 14 12 9

Ni 77 73 52 48 40 33 31 28 26 19

Cu 11 8 25 23 31 26 23 25 20 18

Zn 60 70 94 99 102 88 87 83 78 67

Ga 17 17 24 24 23 22 22 22 20 21

Y 23 24 26 20 21 23 20 18 13 12

Nb 4.1 4.1 23.2 24.1 30.7 38.9 35.7 31.4 26.6 25.1

Ta 0.3 0.3 1.8 1.9 2.2 3.3 2.5 2.3 1.9 1.9

Zr 111 111 118 252 292 346 324 294 259 213

Hf 2.8 2.9 3 6.6 7.3 8.9 8.2 7.4 6.5 5.7

Mo 0.3 0.2 1.9 1.0 1.2 1.2 1.4 0.9 1.6 0.4

Tl 0.17 0.09 0.22 0.75 0.40 0.49 0.59 0.52 0.46 1.07

Pb 4.14 5.79 12.7 12.8 17.2 22.5 20.9 20.5 26.8 24.0

U 0.34 0.30 3.50 1.72 5.33 3.54 3.03 3.24 3.42 4.37

Th 1.42 1.02 12.2 10.9 13.2 17.6 18.8 14.8 15.0 15.7

La 8.69 8.58 103 104 88.8 92.7 82.3 77.7 61.4 49.6

Ce 20.8 21.0 219 195 194 199 175 160 101 105

Pr 2.91 2.81 25.9 21.2 21.0 21.1 18.6 16.8 9.62 11.8

Nd 12.9 13.2 96.2 74.9 79.0 71.56 66.3 56.3 31.5 37.8

Sm 3.60 3.68 14.8 10.6 9.94 10.6 9.4 7.21 4.70 4.96

Eu 1.37 1.32 3.87 2.85 2.55 2.25 2.38 1.91 1.38 1.35

Gd 3.88 4.04 9.42 7.14 6.84 7.11 6.28 5.52 3.56 3.59

Tb 0.67 0.69 1.20 0.85 0.86 0.87 0.79 0.67 0.47 0.46

Dy 4.08 4.38 5.54 4.14 4.01 4.05 3.84 3.32 2.46 2.41

Ho 0.89 0.90 0.93 0.70 0.70 0.77 0.67 0.60 0.49 0.43

Er 2.21 2.4 2.14 1.70 1.75 1.94 1.65 1.48 1.19 1.08

Tm 0.36 0.36 0.30 0.23 0.27 0.29 0.25 0.22 0.17 0.17

Yb 2.11 2.21 1.85 1.38 1.56 1.77 1.43 1.30 1.09 1.05

Lu 0.32 0.33 0.26 0.2 0.22 0.26 0.22 0.19 0.18 0.14

TSZ 664 665 681 755 790 814 810 814 802 810

STB-39 and STB-6 are Minor Gabbros; the rest belongs to the Main Series. TSZ is the zircon saturation temperature (8C). Mean is the grand

average of the Main Series, calculated weighting every facies according to its cartographic representation: 10% of gabbrodiorites and

monzodiorites, 20% of quartz–monzodiorites, 50% of granodiorites, 20% of monzogranites, 10% syenogranites, 10% of alaskites.

F. Bea et al. / Lithos 81 (2005) 209–233218

STB-22 STB-11 STB-18 STB-10 STB-1 STB-20 STB-37 STB-23 STB-21 Mean

68.03 72.35 74.09 74.28 74.53 74.87 76.28 76.3 77.35 63.2

0.61 0.37 0.26 0.20 0.28 0.27 0.17 0.16 0.09 0.86

16.03 14.25 13.77 14.58 13.75 13.54 13.19 12.91 12.51 16.6

3.55 2.17 1.78 1.18 1.63 1.47 1.25 0.94 0.83 4.96

1.13 0.57 0.38 0.27 0.31 0.30 0.12 0.18 0.01 1.98

0.06 0.05 0.05 0.03 0.04 0.03 0.04 0.03 0.02 0.08

2.99 1.56 1.09 1.12 0.83 0.96 0.41 0.57 0.48 4.08

3.84 3.35 3.21 2.93 3.45 3.51 3.5 3.07 3.35 3.57

3.48 4.61 4.78 4.85 4.25 4.26 4.45 4.82 4.35 3.77

0.29 0.13 0.1 0.09 0.08 0.08 0.05 0.05 0.01 0.40

40 41 23 43 26 47 16 33 11 34

241 114 136 276 120 200 174 248 171 118

3.8 1.5 1.6 10.1 1.5 4.1 5.8 2.3 1.9 2.2

3.5 3.2 3 4.6 4.6 4.2 3.5 4.2 2.8 3

782 421 315 160 117 219 69 129 23 854

1278 869 655 378 201 425 132 212 27 1286

7 4 2 3 2 3 2 1 1 10

58 33 22 13 18 15 11 10 6 94

30 17 11 5 7 10 4 5 2 62

9 4 3 1 2 2 1 0 0 12.5

19 10 6 4 7 6 2 4 2 28

11 11 18 11 4 5 3 6 5 20

70 44 36 38 30 43 29 23 19 75

21 19 19 19 18 19 18 18 17 21

21 10 7 11 5 7 7 3 3 18.5

48.1 25.5 23.9 26.4 35.4 36.9 32.7 22.9 22.8 30.8

3.5 1.8 1.9 1.9 2.4 2.9 2.5 1.7 1.9 2.4

272 184 180 141 164 138 110 150 108 249

7 5.1 5.4 4.4 5.3 4.4 4.2 5.3 4.3 6.5

0.6 0.5 0.4 0.1 0.4 0.4 0.3 7.1 0.3 1.1

1.09 0.7 0.99 1.50 0.83 1.03 0.92 1.03 1.1 0.71

27.8 26.8 31.2 38.4 36.0 36.3 31.6 31.9 37.9 23.5

7.24 4.84 3.43 4.91 2.55 4.01 4.54 6.82 8.97 3.90

22.4 26.9 27.3 40.8 30.3 34.8 35.5 20.6 37.7 18.1

64.4 48.1 39.3 46.6 23.0 44.5 24.9 26.2 25.1 75.0

143 95.2 81.3 87.8 51.9 80.7 56.6 49.0 31.0 151

14.0 9.86 6.58 8.59 4.01 6.12 4.04 3.27 2.29 15.7

46.0 31.1 20.1 28.4 12.0 17.6 11.1 8.93 5.38 55.0

6.61 3.79 2.57 4.15 1.70 1.9 1.56 0.94 0.69 7.55

1.26 0.93 0.61 0.46 0.31 0.4 0.18 0.22 0.09 1.90

4.28 2.93 1.54 3.37 1.10 1.36 1.06 0.73 0.42 5.33

0.63 0.4 0.23 0.44 0.17 0.2 0.17 0.11 0.06 0.69

3.19 1.89 1.29 2.25 0.97 1.06 1.02 0.56 0.31 3.41

0.65 0.33 0.26 0.39 0.18 0.2 0.23 0.11 0.06 0.63

1.77 0.86 0.66 0.92 0.51 0.59 0.73 0.33 0.24 1.58

0.32 0.13 0.1 0.14 0.09 0.1 0.12 0.06 0.04 0.24

2.07 0.86 0.65 0.85 0.54 0.67 0.85 0.4 0.32 1.46

0.31 0.12 0.1 0.12 0.09 0.11 0.13 0.07 0.06 0.22

824 800 804 791 802 783 769 794 766 –

F. Bea et al. / Lithos 81 (2005) 209–233 219

Fig. 5. Chondrite-normalized REE patterns (normalization values from McDonough and Sun, 1995). Note the marked difference between the

Minor Gabbros and the gabbrodiorites of the Main Series. The progressive developments of an Eu anomaly and a strong depletion in Dy–Er are

both caused by titanite fractionation, according to the partitioning of REE between minerals (in preparation).

F. Bea et al. / Lithos 81 (2005) 209–233220

P2O5. The correlation with MgO (R=�0.994) and

Al2O3 (R=�0.983) are also excellent but curvilinear.

The correlation with the alkalis is much poorer:

curvilinear positive but notably scattered with K2O

and, if the whole series is considered, nonexistent

with Na2O. The saturation in alumina increases with

silica, so that the series becomes peraluminous for

SiO2N67–68 wt.% (Fig. 10B). Fetot./(Fetot.+Mg) also

increases with silica, shallowly at first in the

gabbrodiorites to the monzogranites, then more

steeply from the monzogranites to the leucogranites

(Fig. 7).

Sc, V, Co and Ni (RV�0.995) and Sr (R=�0.988)

show excellent linear negative correlations with silica

(Figs. 8 and 9). The plots of Cr, Cu, Zn, Ba, Zr and the

REE against silica are slightly scattered and curved,

either flat or positive up to SiO2c57 wt.% and

negative from this point onward. Positive correlations

are never so good as the negative ones (Fig. 9); the

highest coefficient (R=0.952) belongs to Pb, followed

by Th (R=0.827), Tl (R=0.715) and Be (R=0. 605).

The elements that increase with silica, especially those

with a high tendency to be mobile in hydrothermal

solutions such as Li, Cs and U, show a marked

dispersion in the silicic end of the series, this same

effect is displayed by Na2O.

Fig. 9 also shows the variation of some element

ratios. Nd/Th decreases regularly with increasing

silica contents. Zr/Hf stays around 39, slightly higher

than the chondritic value (c37; McDonough and

Sun, 1995), up to the monzogranites and then

decreases linearly to Zr/Hf=24 in the alaskites, a

value characteristic of highly fractionated felsic rocks

(Wang et al., 1996; Linnen, 1998; Linnen and

Keppler, 2002). Al/Ga behaves similarly to Zr/Hf.

However, K/Rb, classically considered as one of the

most reliable indicators of magmatic fractionation,

does not show any clear pattern.

Table 4

LA-ICP-MS analyses (averages) of selected elements in biotite and

hornblende of rocks from the Main Series

SiO2

(wr)

STB-

42

STB-

4

STB-

38

STB-

29

STB-

5

STB-

17

STB-

18

STB-

19

51.06 56.93 59.90 61.01 62.04 67.20 74.09 74.73

Biotite

n spots 11 12 14 12 12 10 11 9

Li 272 268 343 132 261 206 416 1218

Rb 554 571 669 511 601 735 752 1263

Cs 9 10 17 12 11 34 13 32

Ba 3242 3223 2192 2536 3344 994 2835 907

Sc 7.2 7.3 5.4 12 7.1 18.9 26.6 26.4

V 336 328 349 347 334 348 176 202

Co 57 57 45 59 59 49 54 54

Ni 119 120 107 97 123 84 100 111

Ga 107 104 111 69 58 68 102 71

Amphibole

n spots 14 10 13 10 12 10 8

Sr 80 71 58 89 53 17 38

Ba 67 69 48 61 39 13 33

Sc 67 67 85 136 117 93 104

V 307 306 255 267 227 237 221

Co 41 41 37 41 37 35 37

Ni 64 64 61 54 62 101 75

Ga 23 23 21 19 19 18 22

The percentage of SiO2 in the whole rock is also given. Every spot

was done in a different crystal of the same rock section. Sample

STB-18 does not contain hornblende.

Fig. 6. e(Nd)283 Ma vs. e(Sr)283 Ma for Stepninsk rocks. Note that the

11 samples of the Main Series, from gabbrodiorites to alaskites, plo

into a very small area, quite different from the Minor Gabbros

notably enriched in radiogenic 143Nd. Note also the lack of any

tendency of the Main Series toward host metapelites. This reveals

first that all terms of the main Series were derived from a common

source; second, that there is no perceptible assimilation of hos

rocks; third, that Minor Gabbros and Main Series are not cogenetic

F. Bea et al. / Lithos 81 (2005) 209–233 221

5.3.2. Variations of major and trace elements in

Fe–Mg minerals

The compositions of hornblende and biotite do

not follow the systematic variations of their host

rocks. These minerals are equally magnesian in

gabbrodiorites and leucogranites (Fig. 10A) and

contain similar, even increasing concentrations of

strongly compatible trace element such as Ni, Co and

Sc (Fig. 11). V in hornblende decreases slightly from

gabbrodiorites to granodiorites and then keeps nearly

constant to leucogranites; V in biotite keeps constant

up to the monzogranites but then decreases by a

factor of 0.5 in the leucogranites. Li and Rb in biotite

also keep roughly constant from gabbrodiorites to

monzogranites, but then increase substantially in

leucogranites, probably as a consequence of sub-

solidus reactions between biotite crystals and Li–Rb

bearing late-magmatic to post-magmatic hydrother-

mal solutions (Robert et al., 1983; Harlov and

Melzer, 2002).

6. Discussion: a model of deformation-driven

fractionation

6.1. Chemical patterns

Neither the whole-rock geochemistry nor the

composition of rock-forming minerals is compatible

with a sequential process of fractional crystallization,

in which increasingly silicic rocks precipitated from

progressively differentiated residual melts. Had such

a process occurred, the changing modal proportions

of crystallizing phases and the progressive increase of

the partition coefficients of elements such as Co, Ni,

Sc, V and Sr as the liquid become more silicic (Ewart

and Griffin, 1994; Sisson, 1994) would have caused

markedly curvilinear, even discontinuous or angular,

element vs. SiO2 variation patterns. Instead, we

observe clear linear trends, as shown in Fig. 8. It

would equally have caused a smooth variation in the

ratios of elements fractionated by the same minerals

with different partition coefficient such as Zr/Hf and

Al/Ga instead of nearly constant values from

SiO2=51% to 70% and a then steep almost linear

decrease to the alaskites. In addition, sequential

fractional crystallization would have also caused the

composition of ferromagnesian minerals to change in

t

t

.

Fig. 7. Major elements vs. SiO2 plots. The Minor Gabbros plot outside the trends defined by the Main Series, which are excellent linear negative

for TiO2, FeOtot., CaO and P2O5 and slightly curvilinear but also excellent for Al2O3 and MgO. Note the scatter of Na2O in the most silicic

rocks. Compare the variation of Fe/(Fe+Mg) in rocks with that of amphibole and biotite shown in Fig. 10A. Compare also the values of ASI in

rocks with VIAl in biotites (Fig. 10B).

F. Bea et al. / Lithos 81 (2005) 209–233222

parallel with their host rocks, instead of the notable

lack of variation revealed in Figs. 10 and 11.

The linear patterns (Figs. 7 and 8) might have

resulted from the mixing of two components: one

mafic, represented by the gabbrodiorites and mon-

zodiorites, and another felsic, represented by the

alaskites. Using this reasoning Popov et al. (1999)

proposed, for Stepninsk, an origin of hybridation

between a mantle-derived mafic magma, represented

by the Minor Gabbros, and a crust-derived felsic

magma represented by the leucogranites. This hypoth-

esis, however, is unsupportable from both isotope and

element geochemistry. First, it is obvious from Figs.

4–7 that the Minor Gabbros and Main Series are not

cogenetic and there is no mixing line between them;

neither Fig. 6 reveals any perceptible contamination

Fig. 8. Trace element vs. SiO2 plots for the Main Series (1). Note the excellent linear negative patterns of Sc, V, Co, Ni and Sr and the curvilinear

patterns of Ga, Cu, Zn, Ba, Zr, Y and the REE.

F. Bea et al. / Lithos 81 (2005) 209–233 223

Fig. 9. Trace element and selected interelement ratios vs. SiO2 plots for the Main Series (2). Positive correlations are much poorer than negative

ones. Note the scatter in the most silicic edge rocks shown by the elements highly mobile in hydrothermal solutions. Note also the smooth

variations in element ratios except K/Rb. Zr/Hf changes from near chondritic to low values characteristic of highly fractionated granites, see text

for details.

F. Bea et al. / Lithos 81 (2005) 209–233224

with host metapelites. Second, all rocks of the Main

Series, from gabbrodiorites to alaskites, have almost

the same initial 87Sr/86Sr and 143Nd/144Nd (Fig. 6, see

also Fig. 3A) and must, therefore, have been derived

from a common source. Third, neither mixing nor

assimilation can explain the highly differentiated

character of the leucocranites revealed, for example,

by the decrease in Zr/Hf, Al/Ga and Nd/Th.

A mechanism capable of producing linear varia-

tions in element concentrations with constant

isotope ratios is filter pressing acting on a mushy

magma. Provided that solids and melt were in

isotopic equilibrium, squeezing different batches of

the same magma can produce a suite of isotopically

identical residua and segregates, the chemical

composition of which will vary between a melt-

depleted residuum and a totally liquid segregate.

This is roughly the picture displayed by the Main

Series. Considering, besides, that Stepninsk crystal-

lized within the most suitable geodynamic environ-

Fig. 10. (A) Fe/(Fe+Mg) in amphibole and biotite of selected rocks of the Main Series arranged according to increasing silica in the rock in

which both minerals are hosted. Note the lack of any systematic variation and compare with Fig. 7. (B) Hexahedral Al in biotite arranged in the

same way. Compare with the evolution of ASI in Fig. 7.

F. Bea et al. / Lithos 81 (2005) 209–233 225

ment for efficient squeezing, i.e. an active strike-slip

shear zone, we assumed as a starting hypothesis that

the Main Series resulted from differentiating a

partially crystallized magma by deformation-driven

filter pressing. The basics of the proposed mecha-

nism are as follows.

6.2. Proposed mechanism

The separation of tiny crystals of amphibole,

biotite or andesine from a viscous, probably convect-

ing, magma is an inefficient task when the melt

fraction is sufficiently high to prevent aggregation of

the suspended crystals. However, when the amount of

crystals increases to form a tridimensional framework,

the intergranular liquid may be efficiently expelled, by

differential stress, into early formed tensile fractures

in a way very much as proposed by Davidson et al.

(1994) and Sawyer (2000) (see also Bagdassarov et

al., 1996, and Bagdassarov and Dorfman, 1998). The

minimum fraction of solids required to form a

tridimensional framework in a crystallizing magma

has been called the Rigid Percolation Threshold

(RPT) by Vigneresse et al. (1996). Though the exact

value of the RPT depends of the size, morphology and

relative abundance of crystals, values close to 0.55

may well be representative for most granite magmas

(Vigneresse et al., 1996). Once a magma reaches the

RPT, it can be efficiently squeezed until the solid

fraction increases to 0.75–0.80 (the Particle Locking

Threshold—PLT—of these authors), after which the

system becomes totally locked.

In addition to addressing the physics of the

process, we also need to consider the chemical state.

To build a thermodynamically consistent fractionation

model, first we have to determine the major element

composition and water content of the initial magma.

Bearing in mind that isotope data indicate the Main

Series evolved as a closed system, we assumed a

major element composition equal to the weighted

Fig. 12. Color index vs. percentage of solids obtained in MELTS

simulations of the crystallization of a magma with major elemen

composition identical to the Main Series average and different H2O

contents. RTP is the Rigid Percolation Threshold (see text) and the

horizontal grey area represents the color index of the Main Series

gabbrodiorites to quartz–monzodiorites. Only for H2ON4% the

residua of squeezing at the RTP (when it is physically possible)

would have color index matching that of the real rocks.

Fig. 11. Variation in the content of selected trace elements in amphibole and biotite (LA-ICP-MS data, see Table 3) arranged according to the

silica contents in host rocks. Vertical bars represent standard deviations. See text for discussion.

F. Bea et al. / Lithos 81 (2005) 209–233226

mean of all facies, which corresponds to a high-K

granodiorite (Table 3). The water content of otherwise

identical crystallizing magmas greatly influences the

color index and relative proportions of quartz,

plagioclase and alkali feldspar which initially crystal-

lize. For this reason, it is possible to estimate the water

content of the magma which crystallized to produce

the Stepninsk Main Series by comparing numerical

simulations, detailed below, with the modal compo-

sition of the first solids to crystallize, represented by

the gabbrodiorites to granodiorites. Using the MELTS

software (Asimow and Ghiorso, 1998; Ghiorso and

Sack, 1995), we modeled the crystallization of a melt

with major element concentration equal to the Main

Series average and variable proportions of water. With

hornblende, biotite, titanite, quartz, feldspars and Fe–

Ti oxide as fractionating phases, the quartz–magnet-

ite–fayalite oxygen-fugacity buffer, and a pressure of

5 kb—conditions that may well represent the crystal-

lization of the Stepninsk pluton—we observed that the

simulations with 3.5 and 4.5 wt.% H2O match the Q–

A–P trend of gabbrodiorites and monzodiorites (Fig.

2), and the simulations with 4.5 and 5.5 wt.% H2O

match the color index (Fig. 12). We therefore assumed

that the initial magma contained about 4.5 wt.% H2O,

a value common for the Uralian granite magmatism,

which was notably rich in water (Fershtater et al.,

1994, 1998).

The temperature of intruding magma cannot be

determined directly. From the lack of inherited zircons

(see Fig. 3B) and considering that the emplacement

t

F. Bea et al. / Lithos 81 (2005) 209–233 227

level recorded by contact metamorphism (about 16

km) was considerably higher in the crust that the locus

of magma generation, we can make a rough estimate

from the zircon saturation temperature of the less

fractionated rocks (Table 3), about 825 8C. At thistemperature, considerably lower than the MELTS

liquidus (959 8C), the intruding magma would have

contained about 20–30% of solids in suspension

(hereafter the early solids) (Fig. 13). During dike

feeding of the pluton, the magma could have under-

gone local flow differentiation, producing some minor

early cumulates and complementary solid-depleted

magma batches. The main fractionation event, how-

ever, should have occurred once the cooling magma

reached the rigid percolation threshold. At this point,

distinct batches of mushy magma were squeezed and

produced different residua and segregates, the com-

position of which depended on (i) the fraction of early

solids present in each magma batch before in situ

crystallization, (ii) the fraction of solids (early+newly

crystallized) present when squeezing occurred, (iii)

the efficiency of melt segregation. So, the gabbrodior-

Fig. 13. Percentage of solids vs. temperature obtained in MELTS

simulations of the crystallization of a magma with major element

composition identical to the Main Series average and different H2O

contents. Grey bars represent the results of the zircon saturation

temperature (Watson and Harrison, 1983) and the amphibole–

plagioclase thermometer measured from gabbrodiorites to grano-

diorites (Holland and Blundy, 1994). The zircon thermometer

probably marks the temperature of the magma when intruded; note

that for 4.5 wt.% H2O it should have contained about 25% of solids.

The amphibole–plagioclase thermometer is likely to mark the

temperature of squeezing; note that it gives the temperature at which

a magma with 4.5 wt.% H2O would have reached the fraction of

solids at which it can be squeezed.

ites represent cumulates of the early solids, the

monzodiorites and quartz–monzodiorites represent

efficiently squeezed residua, the granodiorites to

monzogranites represent unfractionated or little frac-

tionated magma batches, and the leucogranites repre-

sent almost pure-melt segregates with a few entrained

crystals of amphibole and biotite. In this way, the

proposed model explains why gabbrodiorites and

monzodiorites are always present inside the grano-

diorites as irregular metric to hectometric patches with

diffuse contacts, whereas the leucogranites crop out as

discordant dikes. It also explains why the hornblende–

plagioclase thermometer yields uniform results around

705 to 720 8C in all rocks more mafic than

granodiorites because, according to MELTS, these

are the temperatures at which the Stepninsk magma

reached the Rigid Percolation Threshold and was

therefore able to undergo filter fractionation (Fig. 13).

Lastly, this model also provides a simple explanation

for the lack of variations in the composition of

ferromagnesian minerals. They have nearly the same

composition, irrespective of the rock in which they are

hosted, simply because, according to MELTS, all

amphibole and most biotite precipitated from the

initial magma before squeezing.

6.3. Checking the model

To check model’s consistency with the trace

element vs. SiO2 variation patterns shown in Figs. 8

and 9, we calculated several simulations composed of

15 points each. One of these points representing the

unfractionated initial magma (labelled as 7a in Figs.

14 and 15), another the initial magma depleted in

early solids (7b), another the early solids accumulated

by flow differentiation (es), and the remaining 12

points are six residuum-segregate couples produced

by filter pressing. Three of these couples were

supposedly formed from the unfractionated initial

magma and thus contained 0.30 of early solids; the

difference between them is the solid fraction at which

they were squeezed: 0.55 (labelled as 1r, the

residuum, and 1s, the segregate in Figs. 14 and 15),

0.60 (2r and 2s) and 0.65 (3r and 3s), so that they

contained 0.25, 0.30, and 0.35 of newly crystallized

material, respectively. The other three couples were

supposedly formed from the early-solids-depleted

initial magma; as in the previous case, the difference

Fig. 14. Comparison between the Co vs. SiO2 pattern of the Main Series (black dots) and the filter-pressing differentiation model (crosses).

Model point labelling is as follows: es represents the early solids; 7a is the original magma; 8a is the original magma totally depleted in early

solids; 1r to 6r are residua and 1s to 6s are the coupled segregates (see text for details). In all cases, we assumed a bulk partition coefficient for

the early solids Kes=3. (A) Model with a bulk partition coefficient during squeezing Ksq=3.5; note the almost perfect fit with the real pattern. (B)

The same with Ksq=4.5; the model fits almost equally well. (C) Idem with Ksq=7. Perceptible deviations only occur in the residuum-segregate

couples squeezed at the lowest solid fraction. During filter pressing differentiation, therefore, highly compatible elements tend to produce well-

correlated linear trends which are largely independent of fluctuations in the bulk partition coefficient.

F. Bea et al. / Lithos 81 (2005) 209–233228

between them is the fraction of solids (here totally

newly crystallized) at which they were squeezed, 0.55

(4r and 4s), 0.50 (5r and 5s) and 0.45 (6r and 6s).

We assumed that early solids had a long residence

time within the melt, and both solids and melt were,

therefore, fully equilibrated. The trace element com-

position of coexisting liquid and solid can thus be

estimated by using the well-known equations for

batch crystallization:

Cl=C0 ¼ 1= K þ F 1� Kð Þð Þ ð1Þ

Cs=C0 ¼ K= K þ F 1� Kð Þð Þ ð2Þ

where K represents the bulk partition coefficient, F is

the melt fraction, Cl and Cs are, respectively, the

composition of coexisting liquid and solid at F, and

C0 represents the composition of the initial magma.

The newly formed solids, by contrast, resided

within the melt for a short time before squeezing so

that their equilibrium with the melt should be

approached with the equations for Rayleigh-type

fractional crystallization;

Cl=C0 ¼ F K�1ð Þ ð3Þ

Cts=C0 ¼ F K�1ð Þ=F � 1 ð4Þ

Fig. 15. Generation of curvilinear negative and positive patterns with the filter-pressing differentiation model (crosses). Point labelling is the

same as in Fig. 14. (A) La vs. SiO2. Kes=1.7; Ksq=1.9 for the batches squeezed at a solid fraction of 0.55 (1 and 4); for the rest Ksq=2.1. (B) Zn

vs. SiO2. Kes=1.4; Ksq=1.9 for the batches squeezed at a solid fraction of 0.55 (1 and 4); Ksq=2.1 for the batches squeezed at 0.60 (2 and 5);

Ksq=2.3 for the batches squeezed at 0. 65. (C) Pb vs. SiO2. Kes=0.45; Ksq=0.55 for the batches squeezed at a solid fraction of 0.55 (1 and 4);

Ksq=0.60 for the batches squeezed at 0.60 (2 and 5); Ksq=0.65 for the batches squeezed at 0. 65. (D) Zr/Hf vs. SiO2. The pattern is reproduced

simply by assuming constant Kbulk=2.1 for Hf and Kbulk=2.5 for Zr. See text.

F. Bea et al. / Lithos 81 (2005) 209–233 229

where Cts is the average composition of the total

solids.

The main difficulty in applying Eqs. (1)–(4) is to

link F with the chemical variable used as variation

index, SiO2 in our case. This difficulty can be

overcome, however, by using MELTS to calculate

both the major elements composition of residual

liquids and the modal composition of precipitates at

every crystallization step. The calculations of the 15

points was done as follows: SiO2 for the segregates is

the MELTS’s liquid composition recalculated to 1%

H2O. SiO2 for the residua was calculated from the

modal composition provided by MELTS for a magma

with (i) a major element composition equal to the

Main Series average in the case of the early-solids

bearing residua 1r, 2r and 3r, and (ii) a major element

composition equal to the average of the Main Series

minus 0.30 of early solids for the residua 4r, 5r and 6r.

The composition of the initial magma, 7a, is the Main

Series average. The composition of the initial magma

depleted in early solids, 8a, was calculated from Eq.

(2) with Kes as indicated in Table 5 and F=0.75. In the

case of magma batches depleted in early solids, we

calculated the trace element compositions of the

segregates 4s, 5s, and 6s, and the residua 4r, 5r and

6r with Eqs. (3) and (4), respectively, using 8a as

C0, Ksq as bulk partition coefficient, and melt

fractions of 0.45 (4r and 4s), 0.40 (5r and 5s),

and 0.35 (6r and 6s). In the case of magma batches

containing 0.30 of early solids, we calculated the

compositions of the segregates 1s, 2s, and 3s from

Eq. (4) using 8a as C0, Ksq as bulk partition

coefficient, and melt fractions of 0.7 (1s), 0.65 (2s),

and 0.60 (3s). The composition of residua results

from averaging 0.30 of early solids with 0.25 (1r),

0.30 (2r) and 0.35 (3r) of newly formed solids

calculated from Eq. (4) with the same parameters as

the complementary segregates 1s, 2s and 3s. Table 5

and Figs. 14 and 15 show the results for some

representative elements.

Table 5

Results of trace element modeling for Co, Zn, La and Pb

Model points Co (Kes=3) Zn (Kes=1.4) La (Kes=1.7) Pb (Kes=0.45)

SiO2 ppm, Ksq=3.5 ppm, Ksq=4.5 ppm, Ksq=7.5 ppm Ksq ppm Ksq ppm Ksq

es 51 25 25 25 99.3 102.7 12.3

1r 54 19.8 22.2 25.9 108.6 1.9 100.8 1.9 14.4 0.6

2r 55.5 18.5 20.4 22.9 111.7 2.1 102.7 2.1 15.4 0.6

3r 57 17.3 18.7 20.4 113.5 2.3 100.7 2.1 16.6 0.7

4r 58.5 14.2 14.7 15.1 100.6 1.9 89 2.1 17.6 0.6

5r 59.5 13.3 13.7 13.9 100.9 2.1 86 2.1 19.2 0.6

6r 60.5 12.5 12.7 12.8 99 2.3 82 2.1 20.7 0.7

7a 63.5 12.5 12.5 12.5 75 71 23.5

8a 67.7 8.3 8.3 8.3 68.2 60.4 27.2

1s 73.5 3.4 2.4 0.8 51.4 1.9 43.8 1.9 32 0.6

2s 75.2 2.8 1.8 0.5 44 2.1 37.6 2.1 32.4 0.6

3s 76.5 2.3 1.4 0.3 36.5 2.3 34.5 2.1 32.6 0.7

4s 76.5 1.1 0.5 0 34.6 1.9 25 2.1 39 0.6

5s 77 0.8 0.3 0 25.8 2.1 22 2.1 39.3 0.6

6s 77 0.6 0.2 0 18.2 2.3 19 2.1 37.2 0.7

Kes and Ksq are the bulk partition coefficients for the early solids and the newly formed solids, respectively. See text for details.

F. Bea et al. / Lithos 81 (2005) 209–233230

Fig. 14 shows the model for cobalt, one of the

elements that has an excellent linear negative corre-

lation with silica (see Fig. 8). Let us call Kes the bulk

partition coefficient for early solids and Ksq the bulk

partition coefficient during squeezing. With Kes=3,

Eq. (2) yields exactly the composition of the less

silicic gabbrodiorite; if we then assume Ksq=3.5, the

15 points of the model define a nearly perfect linear

trend that matches almost exactly the real one.

Remarkably, shifting Ksq to higher values has little

influence: Ksq=4.5 brings out an almost equally well-

matching model (Fig. 14B) and Ksq=7 causes percep-

tible departures only in the couples squeezed from

early-solids bearing magma at the lowest fraction of

solids (Fig. 14C). These results indicate that the

partition of compatible elements with a bulk distribu-

tion coefficient of 3 or higher is little affected by

fluctuations in mineral/melt K’s and would cause,

therefore, highly correlated linear trends.

Curvilinear negative patterns such those of Ga, Cu,

Zn, Ba, REE, et cetera (Fig. 8) are also well matched

when the bulk distribution coefficient is below 3. The

variation patterns of these elements can be split in

three parts, 51 to 57 wt.% SiO2, 57 to 72 wt.% SiO2,

and 72 to 77 wt.% SiO2 (Fig. 15). The first part can be

either horizontal or positive; the second part is always

negative; the third part is also negative but steeper.

The slope between 51 and 57 wt.% SiO2 depends on

the difference between Kes and Ksq; it is horizontal

when KsqVKes and becomes positive as the difference

increases. The steeper slope from 72 to 77 wt.% SiO2

is owed to the increment in Ksq of magma batches

squeezed at a higher solid fraction. Two examples can

illustrate this effect. The pattern of La, nearly

horizontal from 51 to 57 wt.% SiO2, is almost

perfectly matched if we assume Kes=1.7 and Ksq

increasing from 1.9 to 2.1 (Fig. 15A). The pattern of

Zn, positive 51 to 57 wt.% SiO2, is also reasonably

well matched if we assume Kes=1.4 and Ksq increas-

ing from 1.9 to 2.3 (Fig. 15B). It should be remarked

that changes in the bulk partition coefficient are not an

artifact but something inevitable owing to changing

modal proportions of precipitates and the dependence

of mineral/melt partitioning on melt composition, well

documented for many elements (for example Adam et

al., 1993; Ewart and Griffin, 1994; Sisson, 1994).

Totally positive patterns are also well reproduced.

For example, the model of Pb fits by assuming

Kes=0.45 and Ksq increasing from 0.55 to 0.65 (Fig.

15C). Here small variations in Ksq cause noticeable

scatter in the most silicic rocks. It is worth mentioning

the capability of the model to reproduce element ratios

as well. For example, Zr/Hf (Fig. 15D) which is

almost flat from 51 to 72 wt.% SiO2 and then

decreases linearly, is simply mimicked by assuming

a slightly higher Ksq for Hf.

The only patterns the model cannot reproduce are

those with a marked dispersion in the silicic end such

F. Bea et al. / Lithos 81 (2005) 209–233 231

as Li, Rb, Cs, U, Nb and Th (Fig. 9). Since these

elements are highly mobile in hydrothermal solutions

and Stepninsk magma was notably rich in water, we

suggest they were considerably redistributed during

the pneumatolitic-hydrothermal stage. This is reflec-

ted by the increased Li, Rb and Cs contents of biotite

of leucogranites, otherwise equally magnesian as

biotites from more mafic rocks. This effect may also

explain the scatter in K/Rb (Fig. 9) and the decoupling

of the Na2O vs. SiO2 pattern in leucogranites with

respect to the rest of the Main Series.

Insofar as it can be ascertained, the squeezing

model seems consistent with isotope data, field and

petrographic observations, the thermodynamics of

magmatic crystallization, and the element variation

patterns of the Main Series. Since a similarly consis-

tent interpretation of all this data set challenges any

model based in magma mixing, assimilation, or any

other closed-system fractionation mechanism, we

propose deformation-driven filter pressing as the main

source of the large geochemical diversity displayed in

the Stepninsk pluton.

6.4. Filter-pressing differentiation in granite

petrogenesis

Filter-pressing differentiation, despite some men-

tion in the 1940s to 1970s (see Propach, 1976, and

references therein mentioned), has never been con-

sidered a major factor in the evolution of granite

magmas (Clarke, 1992). During the 1990s, however,

some authors have outlined its potential importance:

for example Bea et al. (1994) explained the vertical

zonation of a 1-km-thick granite sill as a consequence

of filter pressing caused by the compaction of mushy

granite magma, and Sisson and Bacon (1999)

proposed that filter pressing owing to the release of

gas via second boiling is, probably, the dominant

process of fractionation in granitic plutons at depths

below 10 km.

Our work here indicates that filter pressing can

also be the dominant process in granite magmas

crystallizing at greater depths under compressive

tectonic regimes. Stepninsk is not an isolated

example: most plutons of the Urals show similar

wide-spectrum associations of deformed gabbros-to-

granodiorites and undeformed monzogranites-to-alas-

kites with nearly identical isotope signatures, small

or no age differences and linear geochemical

variation patterns, a set of features that hardly can

be explained by any mechanism other than filter-

pressing differentiation. Remarkably, the Uralian

style of plutonism finds no equivalent in the west

branch of the European Variscan Belt, where the

abundant batholiths are mostly composed of sub-

ordinate granodiorites and dominant granites. Here

rocks more mafic than tonalites are very scarce and

do not seem linked to the granites by closed-system

magmatic differentiation, such as in the Urals, but

have a distinctly different origin (for example Bea et

al., 1999; Roberts et al., 2000). The fact that the

Uralides orogen did not undergo post-collision

extension (Brown et al., 1998) and was only affected

by compressive tectonics all throughout granitic

magmatism (Bea et al., 2002), whereas the Variscan

granites of Iberia were for the most part generated

during or after the extensional collapse (Bea et al.,

2003), tempts us to suggest that wide-spectrum

differentiation in granite magmas occurs when they

crystallize under compressive regimes, and is caused

by deformation-driven filter pressing. On the other

hand, granite magmas crystallizing under extensional

regimes are little affected by this mechanism and

consequently tend to form narrow-spectrum differ-

entiation series.

7. Summary and conclusions

The Main Series of the Stepninsk pluton was

generated by filter-pressing differentiation of a

hydrous high-K granodioritic magma. This magma

intruded into a strike-slip shear zone at 5 kb and ~825

8C, and contained about 30% of early-formed solids,

mostly plagioclase, hornblende and biotite with minor

titanite, which were in equilibrium with the melt.

Local accumulations of these solids, probably by flow

differentiation during emplacement through dikes,

gave rise to the gabbrodiorites currently found as

small patches inside more massive facies. The

crystallization continued until the fraction of solids

reached or surpassed the rigid percolation threshold.

Different batches of mushy magma were then

squeezed producing residua and segregates with

different composition depending on the fraction of

early solids present in each magma batch, the fraction

F. Bea et al. / Lithos 81 (2005) 209–233232

of solids at which squeezing occurred and, especially,

the efficiency of melt segregation. The monzodiorites

and quartz–monzodiorites represent efficiently

squeezed residua, the granodiorites to monzogranites

represent unfractionated or little fractionated magma

batches, and the leucogranites represent melt segre-

gates, crystallized once the deformation has totally

ceased. This caused amphibole and biotite, massively

precipitated before squeezing, to have nearly the same

composition irrespective of the rock that they are

hosted in. It also caused different variation patterns of

trace elements, depending on their bulk partition

coefficient. The most compatible, with KbulkN3

yielded excellently correlated (RV0.995) negative

patterns against silica, whereas mildly compatible

(Kbulkb3) yielded curvilinear and more scattered

negative patterns. After crystallization, hydrothermal

activity caused by the water released from the magma

produced a notable redistribution of elements such as

Li, Tb, Cs, U, or Nb.

We conclude that deformation-driven filter pressing

is the dominant process of differentiation of granite

magmas crystallizing at depth under compressive

regimes, and it was responsible of the high geo-

chemical diversity characteristic of Uralian granites.

Acknowledgments

The authors are greatly indebted to Jane H.

Scarrow for her thorough revision of the manuscript,

which greatly contributed to improve it. The very

constructive revisions made by two anonymous

referees are also gratefully acknowledged. This

project has been financially supported by the Spanish

CICYT Project BTE2002-04618-C02-01, and by the

NATO Collaborative Grant EST-CLG-978997.

References

Adam, J., Green, T.H., Sie, S.H., 1993. Proton microprobe

determined partitioning of Rb, Sr, Ba, Y, Zr, Nb and Ta between

experimentally produced amphiboles and silicate melts with

variable F content. Chem. Geol. 109, 29–49.

Asimow, P.D., Ghiorso, M.S., 1998. Algorithmic modifications

extending MELTS to calculate subsolidus phase relations. Am.

Mineral. 83, 1127–1131.

Bagdassarov, N., Dorfman, A., 1998. Granite rheology: magma

flow and melt migration. J. Geol. Soc. 155, 863–872.

Bagdassarov, N.S., Dorfman, A.M., Dingwell, D.B., 1996.

Modeling of melt segregation processes by high-temperature

centrifuging of partially molten granites: 1. Melt extraction

by compaction and deformation. Geophys. J. Int. 127,

616–626.

Bea, F., Pereira, M.D., Corretge, L.G., Fershtater, G.B., 1994.

Differentiation of strongly peraluminous, perphosphorus gran-

ites. The Pedrobernardo pluton, central Spain. Geochim.

Cosmochim. Acta 58, 2609–2628.

Bea, F., Montero, P., Molina, J.F., 1999. Mafic precursors,

peraluminous granitoids, and late lamprophyres in the Avila

batholith: a model for the generation of Variscan batholiths in

Iberia. J. Geol. 107, 399–419.

Bea, F., Fershtater, G.B., Montero, P., 2002. Granitoids of the

Uralides: implications for the evolution of the orogen. In:

Brown, D., Juhlin, C., Puchkov, V. (Eds.), Mountain Building in

the Uralides: Pangea to the Present. American Geophysical

Union, Washington, pp. 211–232.

Bea, F., Montero, P., Zinger, T., 2003. The nature and origin of the

granite source layer of central Iberia: evidence from trace

element, Sr and Nd isotopes, and zircon age patterns. J. Geol.

111, 579–595.

Borodin, L.S., 1987. Petrochemistry of Magmatic Series. Nauka,

Moscow. 261 pp., (in Russian).

Brown, D., Juhlin, C., AlvarezMarron, J., Perez Estaun, A.,

Oslianski, A., 1998. Crustal-scale structure and evolution of

an arc-continent collision zone in the southern Urals, Russia.

Tectonics 17, 158–170.

Clarke, D.B., 1992. Granitoid Rocks. Chapman & Hall, London.

283 pp.

Davidson, C., Schmid, S.M., Hollister, L.S., 1994. Role of melt

during deformation in the deep crust. Terranova 6, 133–142.

Ewart, A., Griffin, W.L., 1994. Application of proton-microprobe

data to trace-element partitioning in volcanic-rocks. Chem.

Geol. 117, 251–284.

Fershtater, G.B., 1984. Eugeosinclinal Gabbro-Granite Rock Series.

Nauka, Moskow. 264 pp., (in Russian).

Fershtater, G.B., 1987. Petrology of Major Intrusive Associations.

Nauka, Moscow. 232 pp., (in Russian).

Fershtater, G.B., Borodina, N.S., Rapoport, M.S., Osipova, T.A.,

Smirnov, B.H., Levin, M.Y., 1994. Orogenic Granitoid Magma-

tism of the Urals. Russian Academy of Sciences, Urals Branch.

247 pp., Miass.

Fershtater, G.B., Bea, F., Borodina, N.S., Montero, P., 1998. Lateral

zonation, evolution, and geodynamic interpretation of magma-

tism of the Urals: new petrological and geochemical data.

Petrology 6, 409–433.

Gerdes, A., Montero, P., Bea, F., Fershtater, G., Borodina, N.,

Osipova, T., Sardakova, T., 2002. Peraluminous granites

frequently with mantle-like isotope compositions: the continen-

tal-type Murzinka and Dzhabyk batholiths of the eastern Urals.

Int. J. Earth Sci. 91, 3–19.

Ghiorso, M.S., Sack, R.O., 1995. Chemical mass transfer in

magmatic processes: IV. A revised and internally consistent

thermodynamic model for the interpolation and extrapolation

F. Bea et al. / Lithos 81 (2005) 209–233 233

of liquid solid equilibria in magmatic systems at elevated

temperatures and pressures. Contrib. Mineral. Petrol. 119,

197–212.

Govindaraju, K., Potts, P.J., Webb, P.C., Watson, J.S., 1994. 1994

Report on Whin Sill Dolerite WS-E from England and Pitscurrie

Micrograbbro PM-S from Scotland: assessment by one hundred

and four international laboratories. Geostand. Newsl. XVIII,

211–300.

Harlov, D.E., Melzer, S., 2002. Experimental partitioning of Rb and

K between phlogopite and a concentrated (K,Rb)Cl brine:

implications for the role of concentrated KCl brines in the

depletion of Rb in phlogopite and the stability of phlogopite

during charnockite genesis. Lithos 64, 15–28.

Hofmann, A.W., 1988. Chemical differentiation of the Earth: the

relationship between mantle, continental crust, and oceanic

crust. Earth Planet. Sci. Lett. 90, 297–314.

Holland, T., Blundy, J., 1994. Non-ideal interactions in calcic

amphiboles and their bearing on amphibole–plagioclase ther-

mometry. Contrib. Mineral. Petrol. 116, 433–447.

Kober, B., 1986. Whole-grain evaporation for 207Pb/206Pb-age

investigations on single zircons using a double filament thermal

ionization source. Contrib. Mineral. Petrol. 93, 482–490.

Kober, B., 1987. Single-zircon evaporation combined with Pb+

emitter-bedding for 207Pb/206Pb age investigations using thermal

ion mass spectrometry and implications to zirconology. Contrib.

Mineral. Petrol. 96, 63–71.

Linnen, R.L., 1998. The solubility of Nb–Ta–Zr–Hf–W in granitic

melts with Li and Li+F: constraints for mineralization in rare

metal granites and pegmatites. Econ. Geol. 93, 1013–1025.

Linnen, R.L., Keppler, H., 2002. Melt composition control of Zr/Hf

fractionation in magmatic processes. Geochim. Cosmochim.

Acta 66, 3293–3301.

McDonough, W.F., Sun, S.S., 1995. The composition of the Earth.

Chem. Geol. 120, 223–253.

Montero, P., Bea, F., 1998. Accurate determination of 87Rb/86Sr and143Sm/144Nd ratios by inductively-coupled-plasma mass spec-

trometry in isotope geoscience: an alternative to isotope dilution

analysis. Anal. Chim. Acta 358, 227–233.

Montero, P., Bea, F., Gerdes, A., Fershtater, G., Zinkova, E.,

Borodina, N., Osipova, T., Smirnov, V., 2000. Single-zircon

evaporation ages and Rb–Sr dating of 4 major Variscan

batholiths of the Urals—a perspective on the timing of

deformation and granite generation. Tectonophysics 317,

93–108.

Popov, V.S., Tevelev, A.A., Bogatov, V.I., 1999. The Stepninsk

pluton on the south Urals: relationships of plutonic rocks

coming from mantle and crustal sources. Izv. Vuzov. Geol.

Razved. 5, 52–68.

Propach, G., 1976. Models of filter differentiation. Lithos 9,

203–209.

Robert, J.L., Volfinger, M., Barrandon, J.N., Basutcu, M., 1983.

Lithium in the interlayer space of synthetic trioctahedral micas.

Chem. Geol. 40, 337–352.

Roberts, M.P., Pin, C., Clemens, J.D., Paquette, J.L., 2000.

Petrogenesis of mafic to felsic plutonic rock associations—the

CALC-alkaline Querigut complex, French pyrenees. J. Petrol.

41, 809–844.

Sawyer, E.W., 2000. Grain-scale and outcrop-scale distribution and

movement of melt in a crystallising granite. Trans. R. Soc.

Edinb. Earth Sci. 91, 73–85.

Sisson, T.W., 1994. Hornblende-melt trace-element partitioning

measured by ion microprobe. Chem. Geol. 117, 331–344.

Sisson, T.W., Bacon, C.R., 1999. Gas-driven filter pressing in

magmas. Geology 27, 613–616.

Stacey, J.S., Kramers, J.D., 1975. Approximation of terrestrial lead

isotope evolution by a two-stage model. Earth Planet. Sci. Lett.

26, 207–221.

Tatsumi, Y., Eggins, S.M., 1995. Subduction Zone Magmatism:

Frontiers in Earth Sciences. Blackwell Science, Cambridge,

Mass, USA. 211 pp.

Vigneresse, J.L., Barbey, P., Cuney, M., 1996. Rheological

transitions during partial melting and crystallization with

application to felsic magma segregation and transfer. J. Petrol.

37, 1579–1600.

Wang, R.C., Fontan, F., Xu, S.J., Chen, X.M., Monchoux, P., 1996.

Hafnian zircon from the apical part of the Suzhou granite,

China. Can. Mineral. 34, 1001–1010.

Watson, E.B., Harrison, T.M., 1983. Zircon saturation revisited:

temperature and composition effects in a variety of crustal

magma types. Earth Planet. Sci. Lett. 64, 295–304.

York, D., 1969. Least squares fitting of a straight line with

correlated errors. Earth Planet. Sci. Lett. 5, 320–324.