GeolCarp_Vol52_No3_127

GEOLOGICA CARPATHICA, 52, 3, BRATISLAVA, JUNE 2001

127—138

PHYLLOSILICATES FROM HYDROTHERMALLY ALTERED

GRANITOID ROCKS IN THE PEZINOK Sb-Au DEPOSIT,

WESTERN CARPATHIANS, SLOVAKIA

DANIEL MORAVANSKÝ

, MARTIN CHOVAN

and JOZEF LIPKA

Department of Mineralogy and Petrology, Faculty of Science, Comenius University, Mlynská dolina, 842 15 Bratislava, Slovak Republic

Department of Nuclear Physics and Technology, Slovak University of Technology, Ilkovičova 3, 812 19 Bratislava, Slovak Republic

(Manuscript received September 22, 2000; accepted in revised form March 15, 2001)

Abstract: Biotites from altered granitoid rocks on Pezinok-Kolársky vrch Hill Sb-Au deposit belong to the phlogophite-
annite series with the ratio Fe

tot

/(Fe

tot

+ Mg) = 0.44—0.55. Mg-biotites are the dominant type and can be classified as

phlogopites (Phl

30.09—45.79

Ann

26.67—36.00

Eas

22.93—10.22

Sid

20.31—8.02

). Fe-biotites (annites, Phl

34.79—41.13

Ann

39.20—46.39

Eas

12.22—5.86

Sid

13.79—6.62

) are present in the more altered rocks. The chemical composition of the studied biotites is gov-

erned by Tschermak’s substitution, dioctahedral-trioctahedral substitution and A

—1

(Si

—1

)

-substitution. Interlayer

deficient occupancy of A site is probably caused by postmagmatic hydrothermal alteration and can be explained by
A

—1

(Si

—1

)

-substitution. The studied tri-trioctahedral chlorites come from the clinochlore-chamosite isomorphic

series and they originated from biotites. According to the classification of Wiewióra & Weiss (1990) or Weiss (1991),
they can be divided into two groups: ferrous clinochlores (schematic formula Mg

34.39—36.32

28.31—37.07

24.02—35.39

) and

magnesium chamosites (schematic formula Mg

34.33—37.90

34.76—39.32

23.18—28.30

). Ferrous clinochlore is the dominant type

of chlorite. The chemical composition of chlorites is governed by dominant FeMg

—1

-substitution and also by Tschermak’s

substitution and dioctahedral substitution. The content of impurities (K, Ca, Na) is very low. White K-micas were formed
mainly by alteration of alkali feldspars or plagioclases and can be divided into: phengitic muscovites and illites. Phengitic
muscovites in the phengitic component (Brigatti et al. 2000) range from 0.17 to 0.25. The content of Ti is 0.016—0.020
a.p. 11 oxygens and indicates postmagmatic to hydrothermal origin (Miller et al. 1981). The content of interlayer occu-
pancy (K + Ca + Na) ranges from 0.905 to 0.964 a.p. 11 oxygens and agrees with data from Konings et al. (1984), Piantone
et al. (1994) and others. The chemical composition of phengitic muscovites is governed by Tschermak’s substitution,
dioctahedral-trioctahedral substitution and A

—1

(Si

—1

)

-substitution. The phengitic component for illites is 0.02—0.14.

The content of Ti (0.005—0.019 a.p. 11 oxygens) indicates hydrothermal origin (Miller et al. 1981). The content of K
ranges from 0.770 to 0.926 a.p. 11 oxygens and agrees with data from Cathelineau & Izquierdo (1988), Aja et al. (1991
a,b), Środroń & Eberl (1984) and others. The chemical composition of illites is governed by substitutions A

—1

(Si

—1

)

-substitution, dioctahedral-trioctahedral substitution and Tschermak’s substitution.

Key words: Sb-Au deposit, granitoid rocks, illite, phengitic muscovite, magnesium chamosite, ferrous clinochlore,
annite, phlogopite.

Introduction

The Pezinok-Kolársky vrch Hill Sb-Au deposit (Malé Karpaty
Mts, below only PKV Sb-Au deposit) has been studied quite of-
ten in the last decades (Cambel 1959; Chovan et al. 1992 and
others).  Nonetheless,  the  study  of  hydrothermal  alteration  has
been  given  relatively  little  attention.  Some  basic  information
about hydrothermal alterations and processes connected with al-
teration  were  published  in  Andráš  (1983).  The  present  paper
presents some results of our study of the chemical composition
of phyllosilicates from hydrothermally altered granitoid rocks.

The PKV Sb-Au deposit is situated in the Hrubá dolina Val-

ley on the NE edge of the Pezinské Karpaty Mts (part of the
Malé Karpaty Mts) approximately 7 km NNW from Pezinok.
From 1790 to 1811, and from 1914 to 1991, it was an impor-
tant source of Sb. Approximately one million tons of Sb ore
has been mined from this deposit.

Methods

The samples of granitoid rocks were collected from the Pyritová

adit and New Alexander adit. They represent less (A.I. < 0.35),
more (A.I. = 0.35—0.50) and strongly (A.I. > 0.50) altered rocks (see

A.I. – alteration index defined by Hashiguchi et al. (in Vivallo
1987) and evaluated as the ratio: (MgO + K

O)/(Na

O + K

O + CaO

+ MgO) in Table 1).

Polished thin sections of these samples were studied in transmit-

ted light and were then prepared for microprobe analysis. Electron
microprobe analyses were performed by JEOL 733 Superprobe
(Geological Survey of the Slovak Republic) at 15 kV accelerating
voltage, 1.1—1.4 mA sample current and the following standards:
Si – SiO

, Ti – TiO

, Al – Al

, Fe – Fe

, Mn – rodonite,

Mg – MgO, Ca – wollastonite, Na – albite, K – orthoclase and
Cr – chromite.

Formulae for micas and chlorites were calculated on the basis of

11 oxygen (Bailey 1984; Rieder et al. 1998) and 14 oxygen atoms
(Bailey 1988) using the software Minfile (Afifi & Essene 1988). We
have applied Mössbauer spectroscopy in the study of samples PYH-
15, PYH-9 and PYH-3 for averaging the ratio of Fe

/Fe

and the

amount of Fe

(Lipka 1999). Mössbauer spectra were measured at

room temperature using conventional Mössbauer spectrometer (De-
partment of Nuclear Physics and Technology of the Slovak Univer-
sity of Technology) with source

Co in Rh matrix. Spectra were fit-

ted using the NORMOS program. The accuracy of the IS and QS
data is 0.04 mm/s and the accuracy of the relative area is 0.4 %.

X-ray diffraction analyses of clay material were performed on a

Phillips PW 1710 diffractometer (Geological Institute of the Slovak
Academy of Sciences). Oriented samples were prepared by sedi-
mentation of clay suspension (10 mg/cm

) on glass plates. They

*Corresponding author: dmoravan@nic.fns.uniba.sk

128 MORAVANSKÝ, CHOVAN and LIPKA

were analyzed in the range 2°—50° 2

, with CuK

radiation, Ni fil-

ter, 35 kV generator voltage, 20 mA tube current and divergence
slit, fixed, 1.0 and receiving slit 0.2. The scan step size was 0.02°
and scan step time was 0.8 s.

Bulk rock chemistry was determined by wet chemical analysis

(major oxides) and combination of AAS and ICP (trace elements,
Geological Survey of the Slovak Republic).

Geological setting and petrology

The Pezinok-Kolársky vrch Hill Sb-Au deposit (Fig. 1) lies

among crystalline schists (amphiboles, actinolite schists, gneisses,
phyllites) intersected by vein bodies of Variscan granitoids. The Sb-
Au mineralization is associated with a large fault zone and petro-
graphically variable black schists and phyllites.

Mineral parageneses have been studied by Cambel (1959). On the

basis of Cambel’s results as well as his own studies, Andráš (1983)
distinguished  four  mineralization  stages  during  which  the  hydro-
thermal Sb-Au mineralization formed: stage 1: quartz—arsenopyrite,
with  gold  associated  with  arsenopyrite  and  pyrite,  stage  2:  quartz—
pyrite—arsenopyrite  ±  löllingite,  tetrahedrite,  chalcopyrite,  stage  3:
quartz—carbonate—stibnite  ± gudmundite,  pyrrhotite,  pyrite,  sphaler-

ite, Pb—Sb sulphosalts, berthierite, stage 4: stibnite—kermesite ± an-
timony, valentinite, schafarzikite. The principal ore mineral of Sb-
mineralization is stibnite, in some zones there is an increased con-
tent of berthierite, or sometimes gudmundite and primary kermesite
(Chovan et al. 1992).

Granitoid rocks associated with PKV Sb-Au deposit are represent-

ed  by  the  peraluminous  two-mica  granites  to  granodiorites  of  the
Bratislava  granitoid  massif.  These  rocks  are  enriched  in  quartz,  con-
tain monazite (± ilmenite) and they are characterized as S-type (Cam-
bel & Petrík 1982; Petrík et al. 1994). The rock-forming mineral as-
semblage of unaltered rocks comprises Fe-biotite, primary muscovite,
plagioclase,  sericitized  K-feldspar  and  quartz.  Plagioclases  (17—28
vol. %) are represented by oligoclase with An

16—25

, in leucocratic type

with An

12—15

. They are slightly saussuritized and are zoned with basic

cores and acidic rims. Perthitic K-feldspars (18—33 vol. %) are repre-
sented by microcline or orthoclase in varying amounts. K-feldspars
are only very slightly sericitized. Quartz is abundant (27—28 vol. %).
Biotite (4—16 vol. %, with Fe

tot

/(Fe

tot

+ Mg) = 0.64—0.68) is more

abundant than muscovite (6 vol. %). Biotites often show incipient
chloritization, epidotization, or are altered to muscovites. The chemi-
cal composition of primary celadonitic muscovites (Al

/(Al

+ Si) =

0.17—0.24 and TiO

= 0.48—0.97 wt. %) is close to phengites (Cambel

& Vilinovič 1987; Petrík 1985).

Geochronological data from granitoid rocks of the Malé Karpaty

Mts yield the age 347 ± 4 Ma (Rb-Sr isochrone method) and (

Sr/

Sr)

= 0.7076 ± 0.0013 for granitoids of the Bratislava Massif.

Results

Granitoid rocks from the PKV Sb-Au deposit can be divided

according  to  alteration  index  (Hashiguchi  et  al.  in  Vivallo
1987)  into  three  groups:  less  altered  rocks  (Pyritová  adit,
PYH-15, A.I. = 0.33), more altered rocks (PYH-9, A.I. = 0.38,
PYH-3,  A.I.  =  0.41)  and  strongly  altered  rocks  (New  Alex-
ander adit, NAŠH-19, A.I. = 0.56, NAŠH-20, A.I. = 0.59).

The rock-forming mineral assemblage in less and more al-

tered rocks comprises plagioclase (andesine, oligoclase to al-
bite), Mg- to Fe-biotite and quartz. K-feldspars are probably
albitized or completely altered to white K-micas. Mg- to Fe-
biotites are gradually altered to chlorites. Rutile or sphene are
a by-product of this alteration. Plagioclases are altered to
white K-micas and carbonates. White K-micas intergrow with
fine-grained carbonates and they are dominant in more altered
rocks with A.I. = 0.41. Remnants of allotriomorphic grains of
feldspar and quartz are preserved locally.

The original rock-forming mineral assemblage in the strong-

ly altered rocks (A.I. = 0.56—0.59) could not be identified.
Apart from newly-formed micas and fine-grained carbonates,
the rocks also contain allotriomorphic remnants of quartz,
± feldspar and primary/postmagmatic muscovite.

The chemical composition of altered granitoid rocks is giv-

en in Table 1. Analyses of less altered rock (PYH-15) and
more altered rock (PYH-9) in the multicationic diagram
Q

(de la Roche 1980) lie near average granodiorite (see

Fig. 2) and this finding agrees with the results of Cambel &
Petrík (1982). Strongly altered rocks show that SiO

, Na

were removed and Al

, Fe

, TiO

, CaO, MgO, K

O and

were added to the rocks during hydrothermal alteration.

The bulk rock chemistry and the rock-forming mineral assem-
blage of strongly altered samples (NAŠH-19, NAŠH-20) indi-
cate different magmatic precursor as metaaluminous to alumi-
nous granodiorites of Modra granitoid massif. We suppose that

Fig.  1.  Schematic  geological  map  of  the  Pezinok-Pernek  deposit  crys-
talline complex with marked deposits and occurrences of raw materials
(Chovan et al. 1992). Legend: 1 – metamorphic rocks of the central de-
posit  area,  2  –  Mesozoic,  3  –  granites  of  the  Bratislava  Massif,  4  –
granodiorites  of  the  Modra  Massif,  5  –  Quarternary,  6  –  productive
zones  with  black  schists  and  stibnite—gold  and  pyrite  mineralization,  7
– deposits of Sb, Au, and Fe ores (1 – Pezinok-Kolársky vrch Hill, 2
–  Pernek),  8  –  occurrences  of  Sb,  Au  ores  (3  –  Trojárová,  4  –
Kuchyňa), 9 – gold occurrence (Staré Mesto), 10 – occurrences of py-
rite  ores,  11  –  occurrences  of  polymetallic  and  copper  ores,  12  –
gold placers. Inset shows localization of the Pezinok Sb-Au deposit in
Slovakia.

PHYLLOSILICATES FROM HYDROTHERMALLY ALTERED GRANITOID ROCKS 129

precursor of sample NAŠH-19 and NAŠH-20 was peralumi-
nous two-mica granite of the Bratislava granitoid massif.

Chemical composition of biotites

Biotites from less altered rocks (see Fig. 10A) have a higher

content of Mg (1.386—1.580 atoms per 11 oxygens), a lower con-
tent of Fe

tot

(1.266—1.232 a.p. 11 oxygens), ratio Fe

tot

/(Fe

tot

Mg) = 0.43—0.48 and the content of Ti is 0.136—0.223 a.p. 11
oxygens. The amount of Fe

in the biotite sample PYH-15 is

7.9 %. The ferrous doublets QS

and QS

should be attributed

to iron in ideal M1 and M2 sites. For the good fit in the spec-
trum for sample PYH-15 it was necessary to add one more
doublet QS

with the relative area of about 2.1 %. The QS and

IS parameters of this doublet are between those which are usu-
ally observed for Fe

and Fe

. Corresponding Mössbauer

spectrum is given in Fig. 5A and the quantitative data are sum-
marized in Table 6. The main type of substitution appears to be
combined Al- and Ti-Tschermak’s substitution (Fig. 3A,B).
Al-Tschermak’s substitution according to the reaction:

)

+ (Si

)

= (Al

)

+ (Al

)

(1),

is the dominant mechanism of Al-substitution in biotites. The
effect of this substitution is maximized if biotites are found to
coexist with Al-saturated phases (Tracy & Robinson 1978) and
the final result of this substitution is the end-member eastonite
KMg

AlAl

(OH)

The Al content in biotites from rocks of the PKV Sb-Au de-

posit ranges from 1.373 to 1.559 a.p. 11 oxygens. If we sup-
pose, that the content of Si ranges from 2.606 to 2.655 a.p.
11 oxygens, then redundant Al (

∑

T site = 4) has to be associat-

ed with an octahedral site. Attendance of Al in octahedral site

Fig. 2. The quartz-muscovite-feldspars-biotite rhombus, extension of the Q

triangle (de la Roche 1980). Rocks and minerals: g – granites, gbi –

biotite granites, ad – adamellites, gd – granodiorites, t – tonalites, d – diorites, and – andesites, B – basic rocks, UB – ultrabasic rocks, DS –
detritic sedimentation, qu – quartz, mu – muscovite. The grey arrow shows increasing degree of hydrothermal alteration.

SAMPLE

PYH-15

(L.A.)

PYH-9

(M.A.)

PYH-3

(M.A.)

NAŠH-19

(S.A.)

NAŠH-20

(S.A.)

SiO

63.83

55.86

47.94

63.05

61.15

15.93

17.99

18.34

15.39

15.76

4.50

7.30

5.15

3.60

3.78

TiO

0.759

1.293

1.52

0.649

0.651

CaO

3.97

4.26

6.93

4.22

4.20

MgO

1.63

2.89

3.78

1.64

1.89

MnO

0.067

0.059

0.144

0.100

0.124

0.26

0.41

0.35

0.24

0.23

4.54

4.05

3.23

0.12

0.13

2.64

2.13

3.41

4.09

4.38

tot

0.18

0.14

<0.01

0.86

1.02

<0.01

0.02

0.01

0.41

0.48

0.43

0.31

0.29

0.31

0.28

0.35

0.28

0.32

0.04

1.00

6.98

4.64

4.77

Total

99.076

98.152

98.574

99.209

98.705

124

181

199

1331

1378

709

613

687

804

801

181

101

111

114

204

130

264

393

224

246

242

4323

1627

174

<0.005

1.90

0.74

0.01

<0.01

0.01

151

163

172

0.9

1.2

0.4

1.6

1.9

A.I.

0.33

0.38

0.41

0.57

0.60

Table  1:    Bulk  chemical  composition  of  altered  granitoid  rocks  from
the Pezinok Sb-Au deposit. Oxides in wt. %, trace elements in ppm, Au
in  g/t, and A.I. = alteration index (see text). Analyses of altered grani-
toids  are  arranged  in  the  order  of  increasing  A.I.  L.A.  –  less  altered
rocks, M.A. – more altered rocks, S.A. – strongly altered rocks.

130 MORAVANSKÝ, CHOVAN and LIPKA

cannot be interpreted by substitution (1). Foster (1960a,b) sug-
gested that the additional Al is incorporated into biotite by
means of a dioctahedral-trioctahedral substitution:

3(R

)

= 2(Al

)

(2),

which could be viewed as a muscovite component in biotite
and results in the formation of octahedral vacancies (

). The

low content of octahedral Al

in biotites from the PKV Sb-Au

deposit (0.077—0.165 a.p. 11 oxygens) indicates complementa-
ry character of substitution (2) to substitution (1).

Ti-substitution in biotite is a problematic type of substitu-

tion. While, Engel & Engel (1960) suggested that Ti

replaces

on an octahedral site, many other authors have suggested

the attendance of Ti

(Aubrecht & Hewitt 1980; Dymek

1983; Hewitt & Aubrecht 1986 and others). Ti

-substitution

in an octahedral site is complicated by its high charge and
small cation radius (0.605

—1

nm) compared to Mg

(0.72

—1

nm) and Fe

(0.78

—1

nm) (ionic radii from

Shannon & Prewitt 1969).

Several substitution mechanisms have been suggested for

interpretation of this type of substitution. The first involves a
coupled substitution for an octahedral and tetrahedral site ac-
cording to the reaction:

)

+ 2(Si

)

= (Ti

)

+ 2(Al

)

(3).

Substitution (3) can be viewed as a Ti-Tschermak’s com-

ponent (see Fig. 3A,B) leading to the theoretical end-mem-
ber Ti-eastonite KMg

TiSiAl

(OH)

(Czamanske &

Wones 1973). A second coupled substitution would be ac-
cording to the reaction:

(Al

)

+ (Si

)

= (Ti

)

+ (Al

)

(4).

Direct substitution of Ti for Al on octahedral site accord-

ing to Forbes & Flower (1974), Dymek & Albee (1977) and
also substitution according to the dehydrogenation reaction:

)

+ 2(OH)

—

= (Ti

)

+ 2(O

) + H

(5),

are probably not for biotites from the PKV Sb-Au deposit.
This type of substitution leads to the theoretical end-member
Ti-oxybiotite KMg

TiSi

AlO

The range of Fe

-Tschermak’s substitution appears to be

very small. An additional important mechanism relevant to
Fe

in biotite involved in situ oxidation Fe

, is represented

by the following reaction:

+ (OH)

—

= Fe

+ O

+1/2 H

(6).

Many aspects of the resulting oxyannite end-member

KFe

AlO

were discussed by Eugster & Wones

(1962), Wones (1963a,b) and Wones & Eugster (1965). In the

Fig. 3. Substitutions in biotites, phengitic muscovites and illites from the Pezinok Sb-Au deposit. A – (Si — 2.5) — (Fe + Mg + Mn) plot of bi-
otites, x = Si — 2.5, d = Fe + Mg + Mn, A = K + Ca + Na. B – Ti — and Al

— contents of biotites. C – (Si — 3) — (Fe + Mg + Mn) plot of phen-

gitic muscovites and illites, x = Si — 3, d = Fe + Mg + Mn, A = K + Ca + Na. D – (Si — 3) — IC (interlayer charge) plot of phengitic muscovites
and illites, the unbroken line gives the theoretical trend corresponding to the (A

—1

(Si

—1

)) substitution, A = K + Ca + Na.

PHYLLOSILICATES FROM HYDROTHERMALLY ALTERED GRANITOID ROCKS 131

examined rocks from PKV Sb-Au deposit we consider that
substitution (6) could be more important than Fe

-Tscher-

mak’s substitution.

Substitutions (1), (2), (3), and possibly also substitution

(4), are combined with interlayer deficient occupancy on the
A site. Chemical variation on the A site in biotite may arise
from at least two sources. Substitution H

for K

could ex-

plain a low concentration of interlayer cations. This type of
substitution in biotites from the PKV Sb-Au deposit could
not be confirmed.

We consider that interlayer deficient occupancy on the A

site (0.826—0.942 a.p. 11 oxygens) can be explained accord-
ing to the reaction:

)

+ (Al

)

+ (Si

)

(7),

which relates phlogopite—annite to a talc component

(MgFe

)

(OH)

(Dymek 1983; Hewitt & Aubrecht

1986). Average formula of biotites from PKV Sb-Au deposit is:
(K

0.910

)(Mg

1.477

1.143

0.108

0.011

0.188

0.104

)(Si

2.633

1.367

)

(OH)

Biotites from more altered rocks (see Fig. 10B) have a high-

er content of Fe

tot

(1.461—1.749 a.p. 11 oxygens) and lower

content of Mg (1.185—1.464 a.p. 11 oxygens). The ratio Fe

tot

(Fe

tot

+ Mg) = 0.53—0.55 and the content of Ti is 0.107—0.211

a.p. 11 oxygens. The relative amount of Fe

in the biotite

sample PYH-9 is only 2.2 %. The ferrous doublets QS

and

should be attributed to iron in ideal M1 and M2 sites. Cor-

responding Mössbauer spectrum is given in Fig. 5B and the
quantitative data are summarized in Table 6. Substitutions in
these biotites are similar to substitutions in biotites from less
altered rocks. Whereas the greater portion of analyses in the
Ti-Al

diagram lie below the Ti/(Al

—1) = 1/2 line (Fig. 3B),

we consider that Al-Tschermak’s substitution according to re-
action (1) is the dominant type of substitution. Interlayer defi-
cient occupancy is higher than in the previous case (0.595—
0.771 vs. 0.826—0.942 a.p. 11 oxygens) and indicates the
presence of substitution (7). The average formula is:
(K

0.512

)(Mg

1.346

1.559

0.035

0.015

0.166

0.119

)(Si

2.549

1.451

)

(OH)

Biotites from unaltered/less altered granitoid rocks (Petrík

1985, ZK-50, ZK-51, ZK-53, ZK-60, ZK-126 (Bratislava
granitoid massif) and ZK-5 (Modra granitoid massif)) have a
higher content of Fe

tot

(1.186—1.572 a.p. 11 oxygens) and Mg

(0.684—1.331 a.p. 11 oxygens). The ratio Fe

tot

/(Fe

tot

+ Mg) =

0.47—0.68 and content of Ti is 0.119—0.209 a.p. 11 oxygens.
Our study of granitoid rocks from borehole PT-58 (locality
Trojárová) yielded similar conclusions. These samples have a
higher content of Fe

tot

(1.494—1.674 a.p. 11 oxygens) and Mg

(0.746—0.848 a.p. 11 oxygens). The ratio Fe

tot

/(Fe

tot

+ Mg) =

0.65—0.83 and content of Ti is 0.080—0.146 a.p. 11 oxygens.
These conclusions are not in compliance with our results from
altered granitoid rocks from Pezinok-Kolársky vrch.

The chemical composition of biotites with regard to end-

member phlogopite, annite, eastonite and siderophyllite is
shown in Fig. 4 (de Albuquerque 1973). Whereas biotites from
less altered rocks can be classified as phlogophites (Phl

30.09—45.79

Ann

26.67—36.00

Eas

22.93—10.22

Sid

20.31—8.02

), biotites from more al-

tered rocks have more Fe than Mg and can be classified as an-
nites (Phl

34.79—41.13

Ann

39.20—46.39

Eas

12.22—5.86

Sid

13.79—6.62

). The

ratio of Fe

tot

/(Fe

tot

+ Mg), the rock-forming mineral assem-

blage and the multicationic Q

(de la Roche 1980) indi-

Fig. 4. Chemical composition of biotites from the Pezinok Sb-Au de-
posit with regard to end-members phlogopite, annite, eastonite and si-
derophyllite (de Albuquerque 1973). The grey arrow shows increasing
degree of hydrothermal alteration.

Fig. 5. Room temperature Mössbauer spectra of studied samples from
the Pezinok Sb-Au deposit. A – biotite sample PYH-15, B – biotite
sample PYH-9 and C – chlorite sample PYH-3.

132 MORAVANSKÝ, CHOVAN and LIPKA

cates that the studied granitoid rocks are close to metaalumi-
nous to peraluminous granodiorites of Modra granitoid massif.
Granodiorites from the Modra granitoid massif are enriched in
plagioclase, contain amphibole, accessory epidote, titanite, al-
lanite and they are characterized as I-type (Cambel & Petrík
1982; Petrík et al. 1994). Selected electron microprobe analy-
ses of biotite and biotite-like material, the ratio in percent of
single components and ratio Fe

tot

/(Fe

tot

+ Mg) are given in Ta-

ble 3 and Table 5.

Chemical composition of chlorites

Chlorite-like material from less altered rocks (see Fig. 10A)

differs from the theoretically possible chemical composition of
chlorites  according  to  Wiewióra  &  Weiss  (1990)  or  Weiss
(1991).  It  is  caused  by  incomplete  alteration  of  biotites  into
chlorites  (

R = 6.114—6.258 a.p.f.u.). Because these “chlo-

rites“ do not have octahedral vacancies, we can eliminate the
presence of dioctahedral substitution (according to the reaction
2). The Fe

tot

content ranges from 1.920 to 2.458 a.p.f.u., the

Mg content ranges from 2.639 to 3.099 a.p.f.u. and the ratio
Fe

tot

/(Fe

tot

+ Mg) is 0.53—0.62. We suppose that the relative

amount of Fe

in chlorite sample PYH-15 is similar to the rel-

ative amount of Fe

in biotite sample PYH-15. The average

formula is: (Mg

2.886

2.092

0.197

0.028

0.971

0.038

)

(Si

2.664

1.336

(OH)

and this formula corresponds to fer-

rous clinochlore. Magnesium chamosites have not been identi-
fied in less altered rocks. FeMg

—1

-substitution and Tscher-

mak’s substitution (according to the reaction 1 and 3) are

present only over a small range. Octahedral Al ranges from
0.933 to 1.008 a.p.f.u.

Identified tri-trioctahedral chlorites (

R = 5.510—5.917,

= 0.090—0.483 a.p.f.u.) belong to clinochlore-chamosite

series  and  originated  from  biotites.  According  to  the
Wiewióra & Weiss (1990) or Weiss (1991) classification they
can  be  divided  into  two  groups:  ferrous  clinochlores  (sche-
matic  formula  Mg

34.39—36.32

28.31—37.07

24.02—35.39

) and

magnesium chamosites (schematic formula Mg

34.33—37.90

34.76—39.32

23.18—28.30

, see Fig. 6).

Chlorites from more altered rocks, with A.I. = 0.38—0.41

(see Fig. 10D), range in the Fe

tot

from 1.561 to 2.301 a.p.f.u,

whilst the content of Mg ranges from 1.929 to 2.291 a.p.f.u..
The ratio Fe

tot

/(Fe

tot

+ Mg) = 0.48—0.56. We suppose that the

relative amount of Fe

in chlorite sample PYH-9 is similar to

the relative amount of Fe

in biotite sample PYH-9. The rela-

tive amount of Fe

of chlorite sample PYH-3 is 9.3 %, the

rest of the iron is in two-valence state. Mössbauer spectrum
was fitted with two doublets (QS

and QS

), which correspond

to Fe

, and with doublet QS

that corresponds to Fe

. Corre-

sponding Mössbauer spectrum is given in Fig. 5C and the
quantitative data are summarized in Table 6. The contents of
impurities are very low (Ca = 0.007—0.047, K = 0.035—0.278
a.p.f.u.). The average formula is: (Mg

2.160

1.831

0.139

0.005

1.381

0.098

0.386

)(Si

2.965

1.035

(OH)

, for ferrous

clinochlores and (Mg

2.083

1.991

0.205

0.006

1.406

0.042

0.267

)(Si

2.944

1.056

(OH)

, for magnesium chamosites.

The presence of ferrous clinochlores and the presence of

magnesium chamosites in our samples indicates dominant
FeMg

—1

-substitution. The ratio Fe

tot

: Mg in more altered rocks

REF. SAMPLE

OBS. SAMPLE

Chl

(K)

Ill

2M1

PYH–15 (L.A.)

PYH–9 (M.A.)

PYH–3 (M.A.)

NAŠH–19 (S.A.)

NAŠH–20 (S.A.)

(tab.)

(obs.)

14.300

14.239

36 (Chl)

14.283

42 (Chl)

14.259

31 (Chl)

13.484

3 (Chl)

14.254

1 (Chl)

10.000

10.049

96 (Ill)

10.025

100 (Ill)

10.055

100 (Ill)

10.078

100 (Ill)

10.098

100 (Ill)

9.900

7.120

100

7.096

66 (Chl)

7.109

98 (Chl)

7.095

78 (Chl)

7.068

2 (Chl)

7.111

2 (Chl)

6.375

14 (Pl)

6.392

13 (Pl)

5.030

5.004

30 (Ill)

5.015

30 (Ill)

5.019

30 (Ill)

4.900

4.999

28 (Ill)

4.990

29 (Ill)

4.723

19 (Chl)

4.728

21 (Chl)

4.723

17 (Chl)

4.630

4.520

4.535

10 (Ill?)

4.460

100

4.498

2 (Ill)

4.489

2 (Ill)

4.290

4.250

15 (Qtz)

4.258

2 (Qtz)

4.253

1 (Qtz)

4.110

4.102

2 (Pl)

4.038

24 (Pl)

4.035

18 (Pl)

4.008

1 (Pl)

3.880

3.879

10 (Pl)

3.771

18 (Pl)

3.770

11 (Pl)

3.770

5 (Pl)

3.630

3.650

3.665

15 (Pl)

3.668

12 (Ill)

3.660

5 (Pl)

3.654

2 (Ill)

3.671

2 (Ill)

3.560

3.540

24 (Chl)

3.543

39 (Chl)

3.538

35 (Chl)

3.537

2 (Chl)

3.350

100

3.360

100

3.342

74 (Ill)

3.322

60 (Ill)

3.333

50 (Ill)

3.338

47 (Ill)

3.340

46 (Ill)

3.239

34 (Pl)

3.193

10 (Pl)

3.100

3.190

100 (Pl)

3.192

57 (Pl)

2.900

2.936

12 (Pl)

2.860

2.834

2.836

12 (Chl)

2.829

9 (Chl)

2.792

2 (Chl)

2.648

2.685

1 (Chl)

2.600

100

2.570

100

2.548

2.552

12 (Chl)

2.562

9 (Chl)

2.561

2 (Chl)

2.470

2.491

10 (Ill)

2.498

7 (Ill)

2.504

5 (Ill)

2.505

4 (Chl)

2.435

2.450

2.379

2.390

11 (Chl)

2.270

2.277

12 (Ill)

2.255

2.240

2.180

2.140

2.000

2.002

16 (Ill)

2.003

15 (Ill)

1.990

100

1.992

1.988

17 (Ill)

1.989

22 (Ill)

1.997

19 (Ill)

1.940

Table 2: The d spacing data (in nm

10) and intensites from XRD patterns of clay material from altered granitoid rocks of the Pezinok Sb-Au de-

posit. Ref. sample: Chl

– clinochlore, JCPDS table 19 – 749, Ill

– illite, JCPDS table 2 – 0462 and Ill

2M1

– illite, JCPDS table 9 – 334.

Chl – chlorite, Ill – illite, Qtz – quartz, Pl – plagioclase. L.A. – less altered rocks, M.A. – more altered rocks, S.A. – strongly altered rocks.

PHYLLOSILICATES FROM HYDROTHERMALLY ALTERED GRANITOID ROCKS 133

Table 3: Selected electron microprobe analyses of mica and mica-like material from the Pezinok Sb-Au deposit. Oxides in wt. %, analyses are recal-
culated on the basis of 11 oxygen atoms. Phl – phlogopite, Ann – annite, Eas – eastonite, Sid – siderophyllite, Fe/(Fe + Mg)* – Fe as Fe

tot

, Ph

– phengitic component (see text), Fe

* – calculated on the basis of data from Mössbauer spectroscopy. L.A. – less altered rocks, M.A. – more

altered rocks.

SAMPLE

PYH–15 (L.A.)

PYH–9 (M.A.)

Mineral

Bt-Chl

Ms?

Grain

Point

A10

A11

A12

A28

A10

A11

A19

A23

A24

A25

SiO

34.61

34.56

34.67

34.94

28.63

57.61

34.55

34.66

34.77

34.59

31.11

29.94

50.51

52.53

50.87

30.85

TiO

3.59

3.47

3.91

3.32

1.27

0.86

3.54

3.39

2.63

2.39

3.39

3.03

0.37

0.32

0.40

1.75

16.43

16.17

16.08

15.89

17.53

17.73

16.14

16.06

17.09

17.56

15.60

15.91

30.50

27.46

30.12

18.11

0.00

0.09

1.53

1.48

1.54

1.50

1.88

0.00

1.49

1.52

1.49

1.47

0.50

0.52

0.00

0.51

FeO

17.78

17.28

17.90

17.52

21.95

4.70

17.33

17.67

17.40

17.19

22.17

22.96

1.88

3.03

1.82

22.57

MnO

0.22

0.16

0.00

0.22

0.24

0.00

0.19

0.27

0.00

0.11

0.24

0.20

0.00

MgO

12.21

13.16

12.78

13.16

17.03

3.30

13.09

12.95

12.94

14.07

10.41

11.85

2.17

2.65

2.38

11.59

CaO

0.33

0.00

0.64

0.00

0.25

0.00

0.12

0.33

0.22

0.50

0.25

0.10

0.00

0.69

0.00

0.14

1.90

0.69

0.00

9.29

9.72

9.59

3.30

12.66

9.28

9.66

9.74

8.59

5.55

3.56

10.39

8.20

10.11

3.94

Total

95.99

95.54

96.60

96.14

91.83

97.50

95.61

96.18

96.17

96.02

89.09

88.39

96.18

96.59

96.64

89.51

2.635

2.628

2.655

2.286

3.830

2.634

2.638

2.636

2.606

2.568

2.482

3.321

3.436

3.330

2.502

1.365

1.372

1.345

1.649

0.170

1.366

1.362

1.364

1.394

1.432

1.518

0.679

0.564

0.670

1.498

0.000

0.065

0.000

T site

4.000

0.109

0.088

0.064

0.077

0.000

1.219

0.084

0.078

0.163

0.165

0.086

0.036

1.685

1.555

1.654

0.233

0.206

0.199

0.223

0.189

0.012

0.043

0.203

0.194

0.150

0.136

0.211

0.189

0.019

0.016

0.020

0.107

0.000

0.006

0.106

0.103

0.106

0.104

0.137

0.000

0.103

0.105

0.103

0.101

0.034

0.036

0.000

0.034

1.124

1.093

1.126

1.105

1.454

0.262

1.097

1.116

1.095

1.075

1.531

1.592

0.104

0.166

0.100

1.531

0.014

0.010

0.000

0.014

0.017

0.000

0.013

0.018

0.007

0.010

0.017

0.014

0.000

1.386

1.496

1.444

1.491

2.027

0.327

1.488

1.469

1.463

1.580

1.281

1.464

0.213

0.259

0.232

1.401

O site

2.942

2.989

2.963

2.980

3.647

1.851

2.988

3.110

2.981

3.067

3.160

3.331

2.021

1.996

2.006

3.312

0.027

0.000

0.046

0.000

0.011

0.030

0.016

0.035

0.018

0.009

0.000

0.018

0.241

0.088

0.000

0.903

0.900

0.940

0.930

0.336

1.074

0.903

0.938

0.942

0.826

0.585

0.386

0.872

0.684

0.845

0.408

A site

0.930

0.900

0.940

0.930

0.336

1.120

0.903

0.938

0.942

0.826

0.596

0.416

0.906

0.960

0.951

0.417

Phl

30.09

45.79

43.90

43.04

45.79

44.98

45.77

35.32

34.79

41.13

Ann

26.67

36.00

37.39

35.97

36.78

34.49

43.17

39.20

46.39

Eas

22.93

10.22

10.11

8.98

10.22

10.04

11.23

9.69

12.22

5.86

Sid

20.31

8.02

8.60

8.26

8.02

8.20

8.51

11.82

13.79

6.62

Fe/(Fe+Mg)*

0.47

0.44

0.46

0.48

0.44

0.45

0.43

0.55

0.53

0.34

0.17

0.22

0.17

is approximately 1 : 1, as shown in Fig. 6. Since these chlorites
have

∑

R = 5.510—5.917 a.p.f.u. in an octahedral site, we have

suggested the present of the dioctahedral substitution
(Wiewióra & Weiss 1990; Wiess 1991; Zane & Weiss 1998)
according to reaction (2). This type of substitution causes octa-
hedral vacancies (

= 0.090—0.483 a.p.f.u.) leading to the

Fig. 6. Triangular diagram Fe

tot

, Mg and R

+ (Weiss 1991) for

chlorites from Pezinok Sb-Au deposit. 1 – chamosite, 2 – cli-
nochlore.

formation of tri-dioctahedral chlorites as indicated in some
analyses, on Fig. 7.

The Al content in analysed chlorites ranges from 2.162 to

2.750 a.p.f.u. If the content of Si ranges from 2.838 to 3.219
a.p.f.u., then the excess of Al has to be associated with octahe-
dral site (Al

= 1.229—1.526 a.p.f.u.). The presence of Al on

an octahedral site can be explained only with the help of Ts-
chermak’s substitution (Foster 1962) according to the reaction
(1 and 3) where R

= Mg, Fe

, Mn and R

= Al, Fe

and

Cr.  Chemical  composition  of  chlorites  according  to  the
Wiewióra  &  Weiss  (1990)  or  Weiss  (1991)  classification  is
shown in Fig. 7. Ferrous clinochlore is the dominant type of
chlorite.  The  present  of  clinochlore  has  also  been  confirmed
by X-ray study of clay material (see Table 2). Selected elec-
tron microprobe analyses of chlorite and chlorite-like material
and the ratio Fe

tot

/(Fe

tot

+ Mg) are given in Table 4.

Chemical compositions of white K-micas

White K-micas were formed mainly by the alteration of al-

kali feldspars or plagioclases and can be divided into: phengit-
ic muscovites and illites.

Phengitic muscovites from more altered rocks (see

Fig. 10C) have a lower content of Fe

tot

(0.100—0.161 a.p.

11 oxygens) and Mg (0.213—0.259 a.p. 11 oxygens). The phen-
gitic component (Brigatti et al. 2000) evaluated as a percent-
age of the ratio (Fe

+ Fe

+ Mg + Ti + Mn)/(Fe

+ Fe

Mg + Ti + Mn + Al) is 0.17—0.25. The Ti ranges from 0.016 to

134 MORAVANSKÝ, CHOVAN and LIPKA

0.020 a.p. 11 oxygens. This agrees with its hydrothermal ori-
gin  (see  Fig.  8).  The  dominant  relation  among  the  chemical
constituents  of  phengitic  muscovites  is  given  by  the  Tscher-
mak’s  substitution  (Si

—1

)

(Al

—1

) according to reac-

tion (1) and is demonstrated by the strong positive correlation
of Si

/Al

(Fig. 3C). This substitution is combined with the

important dioctahedral-trioctahedral substitution (2,

0.864—0.994 a.p. 11 oxygens) and with the low A

—1

(Si

—1

)

Table 4: Selected electron microprobe analyses of chlorite and chlorite-like material from the Pezinok Sb-Au deposit. Oxides in wt. %, analyses are
recalculated on the basis of 14 oxygen atoms. Fe/(Fe + Mg)* – Fe as Fe

tot

., Fe

* – calculated on the basis data from Mössbauer spectroscopy.

L.A. – less altered rocks, M.A. – more altered rocks, S.A. – strongly altered rocks.

SAMPLE

PYH–15 (L.A.)

PYH–9 (M.A.)

PYH–3 (M.A.)

Mineral

Chl

Grain

Point

A15

A12

A20

A19

A20

A21

SiO

23.61

26.54

24.01

27.28

28.61

30.54

26.28

27.62

27.71

29.45

28.25

25.99

27.97

27.60

27.14

25.88

TiO

0.11

1.08

0.00

0.85

4.16

0.80

0.10

0.17

0.60

3.80

0.00

0.10

0.00

0.23

19.22

18.21

18.27

18.66

17.25

17.82

19.78

19.97

20.08

20.70

19.76

19.96

19.38

19.83

20.61

21.75

2.04

1.90

2.14

0.53

0.46

0.52

2.39

2.24

2.25

1.69

2.26

2.37

2.24

2.29

2.30

FeO

23.77

22.10

24.99

23.43

20.30

23.29

23.27

21.89

21.92

16.53

22.09

23.07

21.86

22.30

22.43

MnO

0.45

0.38

0.37

0.12

0.00

0.26

0.11

0.00

MgO

18.24

16.67

17.93

14.33

12.27

12.28

13.70

13.79

12.69

13.09

12.64

13.92

14.61

13.40

14.63

12.93

CaO

0.00

0.10

0.00

0.09

0.23

0.42

0.07

0.00

0.06

0.00

0.15

0.07

0.00

0.06

0.00

0.17

2.38

0.00

0.64

1.82

0.26

0.00

0.44

0.62

1.03

0.70

0.00

0.28

0.31

0.00

0.46

Total

87.61

89.36

87.17

85.93

85.10

86.19

85.70

86.12

85.87

85.34

85.70

85.46

86.41

85.83

87.17

85.62

2.530

2.768

2.601

2.956

3.057

3.219

2.838

2.933

2.954

3.016

3.017

2.811

2.959

2.946

2.851

2.776

1.470

1.232

1.399

1.044

0.943

0.781

1.162

1.067

1.056

0.984

0.983

1.189

1.041

1.054

1.149

1.224

T site

4.000

0.958

1.007

0.933

1.288

1.229

1.432

1.355

1.433

1.476

1.515

1.504

1.355

1.376

1.441

1.402

1.526

0.009

0.085

0.000

0.068

0.334

0.063

0.008

0.014

0.048

0.293

0.000

0.008

0.000

0.019

0.199

0.180

0.211

0.047

0.041

0.046

0.239

0.221

0.222

0.161

0.224

0.237

0.220

0.226

0.224

0.225

2.114

1.914

2.247

2.078

1.814

2.053

2.078

1.922

1.932

1.400

1.951

2.064

1.913

1.969

1.948

1.961

0.041

0.034

0.011

0.000

0.023

0.010

0.000

2.914

2.639

2.808

2.266

1.954

1.929

2.205

2.183

2.016

1.999

2.012

2.244

2.305

2.132

2.291

2.068

O site

6.258

6.187

6.233

5.855

5.646

5.628

5.895

5.833

5.778

5.510

5.786

5.917

5.850

5.818

5.873

5.862

0.000

0.011

0.000

0.010

0.026

0.047

0.010

0.000

0.007

0.000

0.017

0.000

0.008

0.000

0.023

0.317

0.000

0.087

0.248

0.035

0.000

0.060

0.084

0.135

0.095

0.000

0.038

0.042

0.000

0.063

Fe/(Fe+Mg)*

0.44

0.47

0.48

0.49

0.52

0.51

0.50

0.52

0.44

0.41

0.37

0.39

0.40

0.49

0.51

Table 5: Selected electron microprobe analyses of mica and mica-like material from the Pezinok Sb-Au deposit. Oxides in wt. %, analyses are
recalculated on the basis of 11 oxygen atoms. Phl – phlogopite, Ann – annite, Eas – eastonite, Sid – siderophyllite, Fe/(Fe + Mg)* – Fe as
Fe

tot

, Ph – phengitic component (see text), Fe

* – calculated on the basis of data from Mössbauer spectroscopy. L.A. – less altered rocks,

M.A. – more altered rocks, S.A. – strongly altered rocks, UA. – unaltered rocks.

SAMPLE

PYH–9 (M.A.)

NAŠH–19 (S.A.)

NAŠH–20 (S.A.)

PT–58 (UA.)

Mineral

Ill

Grain

Point

A26

A27

A33

A34

A35

A38

A39

A25

A31

SiO

33.69

29.28

48.45

46.28

49.56

48.98

48.77

47.96

48.79

48.43

46.05

47.06

34.29

33.65

46.49

45.80

TiO

3.02

2.17

0.32

0.14

0.18

0.11

0.10

0.38

0.09

0.31

0.33

2.29

2.06

0.48

0.52

15.86

15.40

34.72

37.54

34.79

34.99

34.15

32.83

33.50

33.14

30.85

30.01

18.55

18.40

34.84

35.63

0.00

0.15

0.00

0.48

0.54

0.00

FeO

21.43

23.87

0.78

0.35

0.66

0.73

0.72

1.15

0.77

1.17

1.82

1.56

23.23

25.13

1.73

1.41

MnO

0.00

0.19

0.16

0.00

0.02

0.57

0.55

0.00

MgO

9.97

10.96

1.10

0.24

1.16

1.30

1.21

2.00

1.53

1.89

4.02

4.00

6.74

6.88

1.10

0.98

CaO

0.00

0.22

0.00

0.42

0.50

7.58

3.80

10.43

11.01

9.27

9.53

10.00

10.25

9.78

10.22

11.01

11.26

9.78

8.88

10.32

11.02

Total

92.03

86.21

95.80

95.60

95.62

95.64

95.10

94.57

94.46

95.14

94.08

95.27

95.45

95.77

95.38

95.86

2.686

2.508

3.178

3.052

3.224

3.195

3.212

3.197

3.231

3.205

3.133

3.160

2.675

2.634

3.093

3.043

1.314

1.492

0.822

0.948

0.776

0.805

0.788

0.803

0.769

0.795

0.867

0.840

1.325

1.366

0.907

0.957

0.000

T site

4.000

0.176

0.064

1.862

1.969

1.892

1.886

1.863

1.776

1.846

1.790

1.607

1.612

0.381

0.331

1.824

1.835

0.181

0.140

0.016

0.007

0.009

0.006

0.005

0.019

0.005

0.016

0.017

0.135

0.121

0.024

0.026

0.000

0.008

0.000

0.032

0.039

0.000

1.429

1.710

0.043

0.020

0.036

0.040

0.039

0.064

0.043

0.065

0.104

0.087

1.516

1.645

0.096

0.080

0.000

0.014

0.000

0.037

0.036

0.000

1.185

1.400

0.108

0.024

0.113

0.126

0.119

0.199

0.151

0.186

0.408

0.400

0.784

0.803

0.109

0.097

O site

3.003

3.367

2.029

2.020

2.050

2.058

2.034

2.058

2.045

2.057

2.136

2.116

3.853

2.133

2.053

2.038

0.000

0.019

0.000

0.054

0.065

0.771

0.416

0.873

0.926

0.770

0.793

0.840

0.872

0.827

0.861

0.956

0.964

0.974

0.887

0.876

0.935

A site

0.771

0.416

0.873

0.926

0.770

0.793

0.840

0.872

0.827

0.861

0.956

0.964

0.974

0.906

0.930

1.000

Phl

37.90

33.97

28.48

26.96

Ann

46.34

41.53

55.27

54.74

Eas

7.09

11.04

5.52

6.04

Sid

8.67

13.46

10.73

12.26

Fe/(Fe+Mg)*

0.55

0.66

0.67

0.08

0.02

0.08

0.14

0.13

0.25

0.24

0.08

0.07

PHYLLOSILICATES FROM HYDROTHERMALLY ALTERED GRANITOID ROCKS 135

-substitution of Al by Si together with a decrease in interlayer
occupancy on A site according to reaction (7). The content of
octahedral Al ranges from 0.778 to 0.843 a.p. 11 oxygens and
content of A site (K + Ca + Na) ranges from 0.905 to 0.964 a.p.
11 oxygens. The average formula is: (K

0.869

0.056

0.014

)

(Al

1.623

0.018

0.112

0.302

0.003

0.942

)(Si

3.276

0.724

)

(OH)

Illites from strongly altered rocks (see Fig. 10D) have a low-

er content of Fe

tot

(0.020—0.065 a.p. 11 oxygens), whilst the

content of Mg ranges from 0.024 to 0.187 a.p. 11 oxygens.

Fig. 8. Distinction between magmatic, post to late-magmatic and hy-
drothermal white K-mica from the Pezinok Sb-Au deposit by means of
TiO

, FeO and MgO (Miller et al. 1981). The grey arrow shows in-

creasing degree of hydrothermal alteration.

Table 6: Parameters (QS, IS) and relative areas (A

rel

) of the components from biotite samples and chlorite sample of the Pezinok Sb-Au deposit.

L.A. – less altered rocks, M.A. – more altered rocks, S.A. – strongly altered rocks.

rel

SAMPLE

Mineral

(mm/s)

(%)

(mm/s)

(%)

(mm/s)

(%)

(mm/s)

(%)

/Fe

PYH–15 (L.A.)

2.67

1.02

43.5

2.31

1.00

46.5

1.14

0.72

2.1

0.62

0.35

7.9

0.086

PYH–9 (M.A.)

2.72

1.02

49.8

2.46

1.03

48.0

0.62

0.10

2.2

0.022

PYH–3 (M.A.)

Chl

2.67

1.02

77.7

2.02

0.91

13.0

0.62

0.10

9.3

0.103

Fig. 7. General representation of chlorite chemical composition from
the Pezinok Sb-Au deposit based on the octahedral cations R

, R

, tet-

rahedral cations R

3+(IV)

and vacancy in octahedral position (marked by ).

The field confined by the thick line represents theoretically possible
chemistry of chlorites. Explanations to isoline descriptions: R

3+(VI)

= 4

represents four trivalent octahedral cations and R

3+(IV)

= 1 represents one

trivalent tetrahedral cations, etc. Points designated by letters represents
the following general chemical composition of chlorites: a –
(R

)(R

(OH)

–

)(R

(OH)

–

)(R

(OH)

, d – (R

2.8

0.4

)(R

) O

(OH)

, e

– (R

2.4

1.2

) (R

(OH)

, f – (R

)(R

(OH)

, g

– (R

4.7

1.3

) (R

(OH)

, h – (R

0.5

1.5

) (R

(OH)

i – (R

)(R

(OH)

(Wiewióra & Weiss 1990; Weiss 1991). The

grey arrow shows increasing degree of hydrothermal alteration.

Fig. 9. Principal chemical characteristics of hydrothermal white K-mi-
cas from the Pezinok Sb-Au deposit. M + Al

vs. Si - R

+ Ti. M = K +

Na + 2. (Ca + Ba + Sr), R

= Fe + Mg + Mn. Plotted end-members

(filled  circles)  are  muscovite  (Musc.),  celadonite  (Cela.),  high-charge
beidellite  (Beid.  0.66),  and  phengite  substitution  maximum  (Phen.
max.), 1  –  idioblastic  muscovite,  2  –  illite-phengitic muscovite,  3  –
phengitic  muscovite,  4  –  illite  (Piantone  et  al.  1994).  The  grey  arrow
shows increasing degree of hydrothermal alteration.

136 MORAVANSKÝ, CHOVAN and LIPKA

Content of Ti, which may indicate magmatic, postmagmatic or
hydrothermal origin of micas (Miller et al. 1981) ranges from
0.005 to 0.019 a.p. 11 oxygens. These results are in compli-
ance with the data for hydrothermal micas (see Fig. 8). The
phengitic component (Brigatti et al. 2000) is very low and
ranges from 0.02 to 0.14.

The main type of substitution involves formation of inter-

layer vacancies (A

—1

(Si

—1

)) according to reaction (7)

where

is an interlayer vacancy in the A site (Konings et al.

1985, see Fig. 3D). This type of substitution leads to formation
of illite K

0.65

3.35

(OH)

(Cathelineau & Iz-

quierdo 1988) and later to the creation of pyrophylite end-
member Si

(OH)

. The content of K in samples from

the PKV Sb-Au deposit, range from 0.770 to 0.926 a.p. 11 ox-
ygens and it is close to ideal content of K in illite (Cathelineau
& Izquierdo 1988). The present of illite has also been con-

firmed by X-ray study of clay material (see Table 2). Another
important substitution is dioctahedral-trioctahedral substitu-
tion according to reaction (2) which causes octahedral vacan-
cies (

= 0.942—0.980 a.p. 11 oxygens).

These substitutions are combined with low Tschermak’s

substitution according to the reaction (1) where R

= Al, R

= Mg, Mn, Fe

(Velde 1965, 1967) and is confirmed by the

low content of a phengitic component (0.02—0.14). The con-
tent of octahedral Al ranges from 1.777 to 1.970 a.p. 11 oxy-
gens. The average formula is: K

0.845

(Al

1.861

0.010

0.044

0.128

0.957

)(Si

3.187

0.813

(OH)

The chemical composition of the primary muscovites

(Petrík 1985, ZK-8, ZK-49, ZK-50, ZK-51, ZK-53, ZK-54,
ZK-60 and ZK-126 (Bratislava granitoid massif)) is analogous
to the phengitic muscovites studied. These micas have Fe

tot

content varying from 0.146 to 0.252 a.p. 11 oxygens, the con-

Fig. 10. Back scattered electron images of the thin polished section of phyllosilicates from altered granitoid rocks of the Pezinok Sb-Au deposit. A
– Back scattered electron image of the thin polished section of chloritized biotites (less altered rocks), B – Back scattered electron image of the
thin polished section of chloritized biotites (more altered rocks), C – Back scattered electron image of the thin polished section of altered plagio-
clases with illites and phengites (more altered rocks) and D – Back scattered electron image of the thin polished section of illites (strongly altered
rocks). Legend: Ank – ankerite, Ap – apatite, Bt – biotite, Ca – calcite, Dol – dolomite, Chl – chlorite, Ill – illite, Ms – muscovite, Pl –
plagioclase, Rt – rutile.

Pl + Ill

50 m

Ca + dol

Ill

Chl

Dol + ank

50 m

Chl

Pl + Ms

50 m

Pl + Ms

Chl

50 m

PHYLLOSILICATES FROM HYDROTHERMALLY ALTERED GRANITOID ROCKS 137

tent of Mg ranges from 0.084 to 0.126 a.p. 11 oxygens and the
content Ti from 0.023 to 0.050 a.p. 11 oxygens. The phengitic
component is 0.07—0.16. Our study of granitoid rocks from
borehole PT-55, 58 (locality Trojárová) yielded similar conclu-
sions. These samples have a content of Fe

tot

(0.059—0.110 a.p.

11 oxygens) and of Mg (0.071—0.109 a.p. 11 oxygens). The
phengitic component is 0.06—0.09 and the content of Ti is
0.019—0.080 a.p. 11 oxygens. The content of A site (K + Ca +
Na) ranges from 0.930 to 0.999 a.p. 11 oxygens.

A more exact classification of single analyses according to

Piantone et al. (1994) is shown in the M + Tet. Al vs. Si—R

Ti diagram (see Fig. 9). Analyses from Petrík (1985) and bore-
hole PT-55, 58 (locality Trojárová) lie in the field of musco-
vite, analyses from more altered granitoid rocks lie in the field
of  phengitic  muscovite  and  analyses  from  strongly  altered
rocks  lie  in  the  field  of  illite-phengitic  muscovite.  No  illites
s.s.  were  found.  Selected  electron  microprobe  analyses  of
phengitic  muscovites,  illites  and  the  percentage  value  of  the
phengitic component are given in Table 5.

Conclusions

The ratio of Fe

tot

/(Fe

tot

+ Mg), the rock-forming mineral as-

semblage and the multicationic Q

(de la Roche 1980)

indicate that studied granitoid rocks are close to metaalumi-
nous to peraluminous granodiorites of the Modra granitoid
massif. These conclusions are not in compliance with Cambel
& Vilinovič (1987). We consider that some granitoid bodies in
the Pernek Unit (Putiš 1992) have a very close relationship to
the Modra granitoid massif.

Biotites from altered granitoid rocks on Pezinok-Kolársky

vrch Hill Sb-Au deposit belong to the phlogophite-annite se-
ries with the ratio Fe

tot

/(Fe

tot

+ Mg) = 0.44—0.55. Mg-biotites

are the dominant type and can be classified as phlogopites
(Phl

30.09—45.79

Ann

26.67—36.00

Eas

22.93—10.22

Sid

20.31—8.02

). Fe-biotites

(annites, Phl

34.79—41.13

Ann

39.20—46.39

Eas

12.22—5.86

Sid

13.79—6.62

) are

present in the more altered rocks. Electron microprobe analy-
ses indicate that chemical composition is governed by substi-
tutions (1), (2), (3), (4) and (7). Interlayer deficient occupancy
of A site is probably caused by postmagmatic hydrothermal al-
teration  and  can  be  explained  by  substitution  (7).  From  the
change  of  chemical  composition  of  biotites  across  alteration
zones  it  is  obvious  that  Al

and K

O were gradually re-

moved, whereas FeO was added to the biotites. The relative
amount of Fe

(below 10 %) indicates a decreasing degree of

oxidation with an increasing degree of hydrothermal alter-
ation. The Fe

is removed from biotites and during alteration

is added to the chlorites. The reduction of Fe

into Fe

in the

studied granitoid rocks is connected with the chemical nature
of the wallrocks of the sulphidic PKV Sb-Au deposit.

The studied tri-trioctahedral chlorites are from clinochlore-

chamosite isomorphic series and they originated from bi-
otites. According to classification Wiewióra & Weiss (1990)
or Weiss (1991), they can be divided into two groups: ferrous
clinochlores (schematic formula Mg

34.39—36.32

28.31—37.07

24.02—35.39

) and magnesium chamosites (schematic formula

34.33—37.90

34.76—39.32

23.18—28.30

). Ferrous clinochlore is the

dominant type of chlorite. The chemical composition of chlo-
rites is governed by dominant FeMg

—1

-substitution and also by

substitutions (2), (1) and (3). The content of impurities (K, Ca,
Na) is very low.

White K-micas were formed mainly by alteration of alkali

feldspars  or  plagioclases  and  can  be  divided  into:  phengitic
muscovites  and  illites.  Phengitc  muscovites  in  the  phengitic
component (Brigatti et al. 2000) ranges from 0.17 to 0.25. The
content  of  Ti  is  0.016—0.020  a.p.  11  oxygens  and  indicates
postmagmatic  to  hydrothermal  origin  (Miller  et  al.  1981).
Content of interlayer occupancy (K + Ca + Na) ranges from
0.905 to 0.964 a.p. 11 oxygens and agrees with data from Kon-
ings et al. (1984), Piantone et al. (1994) and others. The chem-
ical composition of phengitic muscovites is governed by sub-
stitutions (1), (2) and (7).

The phengitic component for illites is 0.02—0.14. The con-

tent of Ti (0.005—0.019 a.p. 11 oxygens) indicates hydrother-
mal origin (Miller et al. 1981). The content of K ranges from
0.770  to  0.926  a.p.  11  oxygens  and  agrees  with  data  from
Cathelineau & Izquierdo (1988), Aja et al. (1991a,b), Środroń
& Eberl (1984) and others. The chemical composition of illites
is  governed  by  substitutions  (7),  (2)  and  (1).  The  change  of
chemical  composition  of  white  K-micas  across  alteration
zones indicates that FeO and MgO were gradually removed,
whereas Al

and K

O were added to phengitic muscovites

and illites.

These conclusions agree with data from other Sb-Au depos-

its. Similar newly-formed minerals with chemical composition
close to minerals studied and similar alteration processes,
were described in France, for example: Haut Allier deposit
(Bril & Beaufort 1989), Le Bourneix (Touray et al. 1989), Le
Chatelet (Piantone et al. 1994) and also in the Dúbrava (Slova-
kia) Sb-Au deposit (Orvošová et al. 1998).

Acknowledgments:  The  financial  support,  obtained  from
Grant Agency for Science (VEGA No. 1/5218/98) as well as
from the Geological Survey of the Slovak Republic (ŠGÚDŠ)
No.  0599160,  is  gratefully  acknowledged.  We  would  like  to
thank Mgr. Daniel Ozdin (ŠGÚDŠ) for microprobe analyses
of  phyllosicates,  RNDr.  ubica  Puškelová  (Geological  Insti-
tute of the Slovak Academy of Sciences) for X-ray diffraction
analyses of clay material and Ing. Ignác Tóth (Department of
Nuclear Physics and Technology of the Slovak University of
Technology)  for  Mössbauer  spectroscopy  of  phyllosilicates.
Thanks also go to doc. RNDr. Pavel Fejdi, CSc. for reviewing
an early version of the manuscript and to all reviewers.

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