GeolCarp_Vol51_No1_19

GEOLOGICA CARPATHICA, 51, 1, BRATISLAVA, FEBRUARY 2000

1936

CALCULATIONS OF TECTONIC, MAGMATIC AND RESIDUAL STRESS

IN THE TIAVNICA STRATOVOLCANO, WESTERN CARPATHIANS:

IMPLICATIONS FOR MINERAL PRECIPITATION PATHS

MICHAL NEMÈOK

1a,b*

, PATRIK KONEÈNÝ

and ONDREJ LEXA

Slovak Geological Survey, Mlynská dolina 1, 817 04 Bratislava, Slovak Republic

Department of Geology and Paleontology, University of Salzburg, Hellbrunner Strasse 34, 5010 Salzburg, Austria

Department of Petrology, Charles University, Albertov 6, 128 43 Prague, Czech Republic

(Manuscript received September 1, 1998; accepted in revised form September 28, 1999)

Abstract: This paper describes the dynamics controlling the development of the tiavnica stratovolcano and the distri-

bution of related ore deposits during the BadenianPannonian (16.56.2 Ma) using combined structural, geobarometric

and geothermometric analyses. During the early Badenianearlier late Badenian (16.515 Ma) the tiavnica area expe-

rienced NE-SW extension. It controlled NW-SE striking normal- and N-S trending sinistral strike-slip faults. Later,

during the early Badenianearlier late Badenian, the extension rotated to W-E and controlled N-S striking normal faults,

NE-SW striking sinistral and NW-SE striking dextral strike-slip faults. Later, the extension rotated to NW-SE. The tec-

tonic stress interacted with the rapid overburden removal and magmatic stress of the granodiorite intrusion, during the

late Badenianearly Sarmatian. The tectonic stress controlled NE-SW trending normal faults. Overburden reduction

input changed normal faulting to strike-slip faulting. Magmatic input caused the opening of subhorizontal veins above

the intrusion and anastomosing fracture pattern around its boundary. Later during the late Badenianearly Sarmatian, the

stress had a plane stress character (

), indicating neither collapse nor distinct regional extension. NE-SW

striking normal faults, ENE-WSW striking sinistral and N-S striking dextral strike-slip faults were active. Normal faults

and releasing bends of the strike-slip faults were places of ore deposit precipitation. During the SarmatianPannonian

(13.610.7 Ma), the stress progressively gained a strong oblate character (

≥

), indicating strong regional NW-

SE extension, controlling the same fault pattern. Later during the Pannonian, counterclockwise stress rotation towards

N-S oriented

and E-W oriented

led to NE-SW oriented sinistral strike-slip faulting.

Key words: Western Carpathians, tiavnica stratovolcano, stress, temperature, fault, vein.

Introduction

During the Miocene, the Carpathian orogen migrated to-

wards the NE and E, synchronously with the subduction of

the remnant Carpathian Flysch Basin, underlain by an oce-

anic/paraoceanic crust and located between the overriding

Carpathians and European platform. This migration ceased

when the Carpathians made contact with the passive margin

of the European Platform (e.g. Royden & Baldi 1988) and

subduction stopped. Another control of the Carpathian de-

velopment was the eastward lateral extrusion of the Eastern

Alps after the Apulia-Europe collision (e.g. Ratschbacher et

al. 1991a,b; Csontos et al. 1992). The shortening in the fron-

tal, accretionary, parts of the Carpathians and subduction

rollback were compensated by the intra- and back-arc exten-

sion, accompanied by the asthenosphere elevation in the

back-arc region (e.g. Royden et al. 1982, 1983a,b; Bergerat

1989; Stegena et al. 1975). This geotectonic situation and

the distribution of the old continental crust in the Inner Car-

pathians controlled the space- and time distribution and com-

position of the Miocene-Quaternary Carpathian volcanism

(Lexa et al. 1993).

The temporal development of the tiavnica stratovolcano

(Fig. 1) is as follows (Lexa et al. 1999; Mao et al. 1996;

Koneèný et al. 1983; Koneèný & Lexa 1984; Kantor & Ïur-

kovièová 1985; Kantor et al. 1988).

The stratovolcano started to form during the early Bade-

nian (16.515.5 Ma) (Fig. 2a), built by pyroxene/horn-

blende-pyroxene andesite lava flows, extrusive domes, py-

roclastic flow deposits and epiclastic breccias. Andesite/

andesite porphyry sills and lacoliths were emplaced in its

central and basal parts. Related alterations comprised chlor-

itization and hematitization of intrusions. The stratovolcano

then underwent denudation roughly synchronous with em-

placement of the quartz-diorite body in its center during the

Badenianearly Sarmatian (16.512.7 Ma) (Fig. 2b).

During the late Badenianearly Sarmatian (1512.7 Ma),

activity of the high sulphidation fluids began causing the

argillitic alteration of host rocks and mezothermal mineral-

izing events. Denudation of the stratovolcano further con-

tinued roughly synchronous with the emplacement of the

large granodiorite intrusion (Fig. 2c). The intrusion, which

had flat roof and outward dipping margins, was situated in

the crystalline basement and its Mesozoic sedimentary cover.

A number of related porphyritic granodiorite apophyses were

emplaced into the Mesozoic rocks. At the contact of the gra-

nodiorite and the Mesozoic carbonates, magnetite skarns

were formed. Local fluids in the upper central parts of the

*Present address: Energy and Geoscience Institute, University of Utah, 423 Wakara Way, Salt Lake City, UT 84108, USA;

mnemcok@egi.utah.edu

CALCULATIONS OF TECTONIC, MAGMATIC AND RESIDUAL STRESS 21

granodiorite intrusion and surrounding andesite led to the

formation of the stockwork base metal mineralization ac-

companied by quartz, pyrite, pyrophylite and kaolinite al-

terations.

D values of inclusion fluids in the quartz and

sphalerite ranging from 52 to 67 SMOW, about 75

in chlorite and kaolinite; isotopic composition of the Pb-rich

ores (

206

Pb/

204

Pb = 18.81018.839,

207

Pb/

204

Pb = 15.659

15.682,

208

Pb/

204

Pb = 38.92239.011) indicate an important

role of magma-derived fluids (Háber et al. 1997). Later on,

during the continuing denudation of the stratovolcano, gran-

odiorite/quartz diorite porphyry stocks and dyke clusters

were emplaced around the granodiorite intrusion (Fig. 2d).

Some of the dykes penetrated by bore hole B-1 (Fig. 3) were

re-opened either through their median lines or along their

margins (tohl et al. 1990a). Fluids related to different

stocks/dyke clusters gave rise to copper porphyry/skarn ore

deposits and host rock alterations.

Later, during the late Badenianearly Sarmatian (1512.7

Ma), it has been postulated that a caldera was formed

(Koneèný & Lexa 1984), filled by sediments, reworked tuffs,

biotite-hornblende andesite/dacite extrusive domes, dome

flows and pyroclastic flows and was accompanied by the em-

placement of quartz-diorite porphyry sills and dykes (Fig. 2e).

During the Sarmatian (13.611.5 Ma), andesitic volcanic

activity from dispersed centers on the slopes of the strato-

volcano revived and the uplift of the horst structure in the

central parts began (Fig. 2f). The continuing uplift was ac-

companied by the emplacement of rhyolitic and granite por-

phyry bodies along the horst boundary faults, during the late

Sarmatianearly Pannonian (1210.7 Ma). Low sulphida-

tion fluids gave rise to base and precious metal vein-type

mineralization and adularia, sericite alterations. K/Ar ages

of 12.313.3 Ma (±0.4 ±1.2) from various sericite samples

constrain the time of this ore deposition (Tschernyschev et

al. 1995).

D values of inclusion fluids in the quartz, baryte

Fig. 1. a) Structural scheme of the central zone of the tiavnica stra-

tovolcano (after tohl et al. 1990b) with localities of kinematic data

collection and fault pattern. 1 rhyolite dyke and extrusive dome,

upper Sarmatianlower Pannonian; 2 rhyolite volcanoclastic

rock, upper Sarmatianlower Pannonian; 3 post-caldera andesite

and volcanoclastic rock, Sarmatian; 4 hornblende-biotite andes-

ite caldera filling and hornblende-biotite andesite porphyry-dyke,

upper Badenianlower Sarmatian; 5 quartz-diorite porphyry

sill, upper Badenianlower Sarmatian; 6 quartz-diorite porphy-

rydyke, upper Badenianlower Sarmatian; 7 cluster of dykes

and stock of granodiorite porphyry, upper Badenianlower Sarma-

tian; 8 granodiorite, Badenianlower Sarmatian; 9 diorite,

Badenianlower Sarmatian; 10 pre-caldera propylitized andesite,

Badenian; 11 pre-caldera andesite and volcanoclastic rock, Bade-

nian; 12 pre-volcanic basement, CarboniferousKarpatian; 13

secondary quartzite, upper Badenianlower Sarmatian; 14

caldera marginal fault; 15 fault; 16 locality. b) Main structures

with ore deposits with original names. caldera marginal fault,

NW-dipping fault, ! SE-dipping fault, " NE-dipping fault,

# subvertical fault. Note the use of classical vein names, used

traditionally but incorrectly for ore deposit bodies formed along

normal and strike-slip faults.

Fig. 2. Cartoon of the development of the tiavnica stratovolcano from 16.5 to 10.5 Ma (after Lexa et al. 1999). Explanation in text.

▲

22 NEMÈOK, KONEÈNÝ

and LEXA

configurations against the fault population from each loca-

tion. Each stress tensor causes certain reactivation of faults,

that is calculated displacement vectors. Calculated displace-

ment vectors are compared with the measured striations after

each test cycle. The stress tensor compatible with a fault

population is accepted as the stress configuration which

caused their activity. It has a form of reduced stress tensor

(Angelier 1989), providing only vectors of principal stresses

and the ratio of their magnitudes. The coefficient of internal

friction and cohesion (0.4 and 0 MPa) for the purpose of this

calculation for ruptures were taken from the literature (e.g.

Jaeger & Cook 1976; Hardcastle 1989). Friction and cohe-

sion for fractures were derived by our own rock mechanics

tests.

However, measured fault data from the outcrops were fre-

quently a mixture of several sub-sets, each of them related to

a different stress tensor, characterized by principal stress ori-

entations and the ratio of their magnitudes. In such a case, in-

compatible faults with first resulting stress tensor were sepa-

rated and testing repeated. This continued until all subsets

were separated and related tensors calculated. The separation

was controlled by the field observations of the structures of

various relative ages. The relative chronology of tectonic

events, represented in our study by their driving force-stress

tensors, at each location was indicated either by cross-cutting

relationships of different faults and veins or by cross-cutting

relationships of different striations on the same fault plane.

All relative chronologies of stress tensors from locations

were correlated and compared with the stratigraphy of the af-

fected rocks in order to infer time periods when the stresses

were large enough to cause veining and/or faulting. Each ten-

sor was interpreted as the cause of the deformation which oc-

curred during the certain time interval.

Separated tectonic events, unlike in sedimentary terrains,

slightly overlap, due to the broader, and occasionally overlap-

ping, age intervals of studied formations (see Koneèný & Lexa

Fig. 3. Profile of the eastern part of the tiavnica stratovolcano (after tohl et al. 1990b). The figure shows the 14

floor of the Rozália

Mine in Banská Hodrua, which is shown in more detail in Fig. 6g.

and carbonates ranging from 94 to 113 SMOW, calcu-

lated oxygen isotope composition of water in equilibrium

with carbonate and barite varying between 3 and 11

indicate aqueous solutions of predominantly meteoric origin

(Háber et al. 1997).

Previous studies, with the exception of the local study of

Mao et al. (1996) in the Rozália Mine, have not used any

structural data to infer either ore deposition models or the

development stages of the stratovolcano. The structural re-

connaissance has shown that collected structural data do not

always support the existing models. The combined structur-

al, geobarometric and geothermometric analyses presented

in this paper aim to test the existence of the postulated

caldera (Koneèný & Lexa 1984), to determine the tectonic

stress configurations active during the activity of ore fluids,

to determine how the magmatic stress interfered with the

tectonic stress, to determine how the overburden changes

due to fast erosion interfered with the tectonic stress and to

determine which faults were active in certain time periods,

that is potential migration paths during the activity of cer-

tain ore fluids.

Methods

Fault-striae data and extensional vein readings were col-

lected and used to determine paleostress configurations us-

ing the computer programs of Sperner et al. (1993) and

Hardcastle & Hills (1991). The former is a Pascal program

for IBM PC compatible computers. The software is based on

Turner (1953) and constructs P and T axes lying in the plane

comprising the fault plane normal and striation vector. The P

axis is inclined 45° from the fault plane and the T axis is per-

pendicular to the P axis. The latter program is a Quickbasic

program. It requires at least 4 faults with determined dis-

placement sense. The method tests a large variety of stress

CALCULATIONS OF TECTONIC, MAGMATIC AND RESIDUAL STRESS 23

Fig. 4. Mohr-circle envelope determined by triaxial testing of the

upper Badenian-lower Sarmatian granodiorite. The determined co-

hesion is 38.9 MPa, the angle of internal friction varies from 40 to

20°, depending on the confining pressure.

1984). This does not mean that separated tectonic events over-

lap, only that their limits are understood as broader bound-

aries.

In order to determine the angle of the internal friction and

cohesion for the reduced stress tensor calculation and the

stress relationship for the calculation of stress magnitudes,

the rock mechanics tests have been carried out on granodior-

ite samples. The samples were cylindrical, 3.78 cm in diame-

ter and 5.755.98 cm in height.

The principal stress magnitudes were computed applying

the following equations of Angelier (1989):

(1)

R = (

)/(

)

(2)

(3)

where R is a stress ratio, calculated by the Hardcastle & Hills

(1991) method. For a given lithology, the

ratio (

) was

derived from the Mohr circle envelope graph. The Mohr en-

velope was constructed as a line inclined to the x axis (nor-

mal stress) at the angle of internal friction, and cutting the y

axis (shear stress) at the value of cohesion. For normal fault-

ing, the value for

equalled the overburden load

, which is

given by

gh, where

is the density of stratovolcanic rocks,

g the acceleration of gravity (rounded to 9.812 ms

), and h

the thickness of overburden. This thickness was given by the

height of the stratovolcanic cone, reconstructed on the basis

of inclinations of preserved lava and pyroclastic flows, and

estimated at about 3000 m (V. Koneèný & J. Lexa 1995, pers.

commun.). The central parts of the missing stratovolcanic

cone were built by intrusive rocks, lava flows and minimally

by volcanoclastics (1520 %; V. Koneèný 1996; pers. com-

mun.). The density of the volcanic basement composed domi-

nantly of calcic dolomite is 27502850 kgm

, density of

andesite 25002600 kgm

, density of volcanoclastics 2000

2200 kgm

, density of quartz-diorite porphyry 2630 kgm

density of diorite porphyry 2630 kgm

, density of diorite

2680 kgm

and density of granodiorite 2640 kgm

(efara

et al. 1976; Ibrmajer et al. 1989). The mineralogical densities,

corrected for porosity, are 2700 kgm

for quartz-diorite por-

phyry, 2680 kgm

for diorite porphyry, 2730 kgm

for dior-

ite and 2710 kgm

for granodiorite.

was calculated by

combining equations 1 to 3:

(R+

-R

), where R, de-

termined from reduced stress calculation, is 0.4 and

, de-

termined from Mohr circle envelope graph (Fig. 4), is

0.0062893.

Having calculated the principal stress magnitudes of the

tectonic stress, the interplay of the tectonic stress, overbur-

den removal and magmatically induced stress was studied.

Increments of overburden removal by erosion were directly

subtracted from the

value. Vertical magmatic stress, cal-

culated by buoyancy equation (e.g. Price & Cosgrove

1991), was also subtracted directly from

. Subtractions of

load and magmatic pressures from

and

were equal to

gh(1-

)) following Jaeger & Cook (1976), where

Poisons ratio,

is the density, g is the acceleration of gravi-

ty and h is the depth. Calculations, which used magmatic

stress, were also made to test the values provided by

geobarometric and geothermometric study, the precision of

which is at the limits of the applied methods.

The effect of the temperature change, related to the over-

burden removal, on the development of the residual stress

was calculated from the equation following Suppe (1985):

∆σ

= [

E/(1-

)] * (dT/dz)

∆

(4)

where

= 7*10

°C

is the linear thermal expansion coef-

ficient, E is the Youngs modulus, dT/dz is the thermal gradi-

ent and

∆

z is the change in depth. The volume change was

calculated from:

dT = v/(100-v)

(5)

where v is the volume change of the intrusion in %.

The T-t path modeling of the contact aureole of the grano-

diorite intrusion was made by the software of Peacock

(1990), using an explicit finite difference algorithm to solve

the one-dimensional heat transfer problem. Granodioritic

magma was assumed to be intruded instantaneously, at tem-

perature 750 °C (own data) and crystallized over a 100 °C

temperature interval. Other assumptions for the modeling are:

during crystallization, the magma released 100 kJkg

, the

temperature of the country rock was 100 °C, the width of the

intrusion is 5 km, the thermal conductivity was 2.75 Wm

(taken from Cermak & Rybach 1982), the heat capacity was

1050 Jkg

K (taken from Cloetingh et al. 1995). Following the

T-t path modeling, the temperature gradient that develops the

cooling-related residual strength higher than the tensional

strength of the granodiorite/andesite was calculated from:

∆

T > [

(1-

)]/

(6)

where

is the tensional strength. Fracturation events were

determined at each point of T-t cooling paths where the ten-

sional strength of the rock was overcome.

In order to determine the magmatic pressures in the mag-

ma chamber and in the upper parts of the granodiorite intru-

sion and the depths of the chamber and top of the intrusion,

26 NEMÈOK, KONEÈNÝ

and LEXA

Fig.

) Cross-cutting

relationships

various

structures

localities

the

Rozália

ine

14th

floor

Banská

Hodrua.

The

14th

floor

located

Fig.

Field

sketches

ere

made

sites

localized

Fig.

6g:

12,

) 14th

floor

the

Rozália

ine

with

localities

kinematic

data

collection.

The

grid

eters,

localities

are

numbered

and

codes

indicate

bore

holes.

Locations

are

different

from

surface

locations

Fig.

CALCULATIONS OF TECTONIC, MAGMATIC AND RESIDUAL STRESS 27

Data

The data comprise fault-striae, extension vein readings

(Fig. 5) and granodiorite sample analyzes.

Most of the extensional veins are filled by idiomorphic

minerals (stretched type sensu Ramsay & Huber 1983) origi-

nated by elastic fracturation accompanied by insufficient flu-

id flow. This type indicates only a rough direction of

, par-

allel to their opening vector, assumed to be perpendicular to

vein walls. Stretched-type veins are not present in the Sarma-

tian andesites, rare in upper Sarmatian-lower Pannonian

rocks (only at location 4 with NW-SE strikes), frequent in

upper Badenian-lower Sarmatian rocks (locations 6e, 33,

39a, 72, 98 with usually NE-SW oriented veins) and abun-

dant in lower Badenian-early upper Badenian rocks (loca-

tions 6a, b, c, 16, 20, 22, 24b, 34b, 35, 47, 74, 75b, 87 with

dominantly NE-SW oriented and subordinate NW-SE veins).

The only fibrous veins, which indicate opening with suffi-

cient fluid flow, and which indicate the opening vector exact-

ly, are present at location 87 (Badenian rocks) and along the

Rozália vein (Figs. 1, 3) having W-E to WNW-ESE oriented

fibers (parallel to

All faults were formed and/or reactivated in the brittle en-

vironment, with the exception of narrow zones, which under-

went a temporal ductile regime due to the thermal and chem-

ical activity of migrating fluids. Faults along the NW and SE

side of the central tiavnica zone dip to the NW and SE, re-

spectively (Figs. 1, 3). The central zone itself is deformed by

a complicated fault pattern containing faults with NW, SE

and vertical dips. Meso-scopic faults from locations are

shown in Fig. 5.

The cross-cutting relationships of various structures in the

field indicate that the oldest fracture pattern is formed by

the anastomosing vein pattern, the network of small frac-

tures with quartz fill, without any preferred orientation (Fig.

6ae). It is accompanied by quartz, pyrite, pyrophylite and

kaolinite alterations. The anastomosing vein system is

cross-cut by a subhorizontal system of quartz and carbonate

veins (Fig. 6ac), called the Svetozár vein system (sensu

Mao et al. 1996). Some of them show evidence of cycles of

increased fluid pressure (Fig. 6f), indicated by repeated epi-

sodes of hydraulic fracturing. Fig. 6f indicates that the

andesite host rock was opened by hydraulic fracturing.

Fracture patterns were conduits for fluids, which migrated

away from the area of the fluid overpressure and caused its

decrease. Decreased fluid pressure triggered mineral precip-

itation, which sealed rock fragments in precipitated quartz

fill (Fig. 6f). Sealing of the escape paths caused a second

cycle of the overpressure that triggered the second fractura-

tion event. It is indicated by pieces of rock lined by quartz

fill that remained in precipitated Mn-carbonate and ame-

thyst fill (Fig. 6f). Sometimes subvertical quartz veins

opened by W-E extension cross-cut the pre-existing subhor-

izontal quartz and carbonate vein system and are cross-cut

by younger subhorizontal quartz-sulphidic vein system

(Fig. 6d). All these structures are older than the quartz dior-

ite porphyry sills and dykes. This is indicated by the intru-

sive contact of the quartz-diorite porphyry with pre-existing

silicified andesite deformed by the anastomosing vein sys-

tem and subhorizontal (Svetozár system-type sensu Mao et

al. 1996) veins (Fig. 6b,c). The relationship of the various

petrographic rock types in dykes indicates multiple intrusion

from differentiated magma chamber. All above mentioned

structures are deformed by the NE-SW normal faults. The

dip of normal faults varies between 50° and 90°. They are

connected by subhorizontal detachments, as is documented

in a few cases (Fig. 6ce). Subhorizontal detachments fre-

quently formed along earlier subhorizontal veins, causing

their boudinage in zones wrapped in mylonite (Fig. 6e).

Two samples were taken from the central part of the gran-

odiorite body, with a coarse granular texture. One was col-

lected from the bore hole (B1/1368) and the other one from

the outcrop. Both samples do not indicate any alteration.

Granodiorite consists of plagioclase, orthoclase, amphibole,

biotite, quartz and Fe-Ti oxides. Some euhedral minerals of

plagioclase have zone of resorption near their rims. Inten-

sive resorption edges are present in amphiboles, less in bi-

otites, indicating that the magma was not oversaturated with

water (only 25 %).

The parameters of the granodiorite determined by the rock

mechanics tests (Fig. 4) are as follows: tensional strength =

2.54.9 MPa, cohesion = 38.9 MPa, angle of the internal fric-

tion = 4020°, Youngs modulus = 6.5 * 10

MPa, Poisons

ratio = 0.22. It has to be said that the tensional strength of the

granodiorite is very low due to its alteration. That is why the

fresh-rock value of 20 MPa (taken from Suppe 1985) was used

for the modeling of fracturation developed by the residual

stress due to cooling.

Results

Amphibole and plagioclase from granodiorite samples are

mineral phases, which crystallized in the early stages of so-

lidification of granodiorite and hence yield information re-

garding physical conditions in the magma chamber. Depen-

dence of the aluminium content on pressure has been used for

the formulation of the amphibole geobarometer (Hammar-

strom & Zen 1986; Hollister et al. 1987; Johnson & Ruther-

ford 1988; Schmidt 1992; Anderson & Smith 1995). Amphib-

ole barometer application was limited by several conditions.

The mineral assemblage had to involve all nine mineral phas-

es: amphibole, biotite, plagioclase, orthoclase, quartz, mag-

netite and titanite. tiavnica granodiorite is pure on titanite

which otherwise has a negligible impact on amphibole com-

position. The anorthite content in plagioclase outer zones

should be in the range An

. The bazicity of studied

plagioclases apparently decreases from core to rim from An

to An

. The role of iron and its ferric/ferrous ratio is funda-

mentally important. More reduced amphiboles with Fe

(Fe

+Fe

) less then 0.25 were not utilized as well as Fe

tot

(Fe

tot

+Mg) outside the 0.40.65 range. Using the equation of

Anderson & Smith (1995) and an average temperature of

about 730 °C, the pressures were found to range from 1.5 to

4.1 kbar with average 2.8 kbar (18 estimations).

Assuming that the total pressure equals the fluid pressure

in a fluid saturated system, the depth of the magma chamber

was estimated. In the case of the tiavnica granodiorite, the

28 NEMÈOK, KONEÈNÝ

and LEXA

system did not achieve a fluid saturated state, indicated, for

example, by resorption of phenocrysts in the late stages of

crystallization. The probable magma chamber is estimated to

be located at depths greater then 10 km. The essential part of

the magma chamber passively intruded to a level about 2.5

km below the surface including the eroded cover (V.

Koneèný & J. Lexa 1999, pers. commun.), determined by

cross-section restoration.

Temperatures in the magma chamber have been evaluated

using the amphibole-plagioclase semi-empirical geother-

mometer of Blundy & Holland (1990). The use of this geo-

thermometer places some limitations on the composition of

the mineral phases. Amphiboles cannot exceed 7.8 silica at-

oms pfu and the plagioclases must be less calcic then An

This thermometer is valid over the temperature range 500

1100 °C with deviation of around 75 °C. The fundamental

condition is to determine equilibrium pairs. The anorthite

content of plagioclases decreases more or less from core to

rim. Small euhedral plagioclases around An

are of-

ten enclosed in amphiboles which crystallized together with

amphiboles. Calculated equilibrium temperatures range

from 690 to 790 °C, indicating crystallization above water

saturated granitoid solidus (

≈

700 °C, Piwinski 1975).

The calculated paleostress configurations were separated

into four groups representing four tectonic regimes during

the: early Badenianearlier late Badenian (16.515 Ma),

late Badenianearly Sarmatian (1512.7 Ma), Sarmatian

(13.611.5 Ma) and late Sarmatianearly Pannonian (12

10.7 Ma) (Table 1). The limits of several of these periods

overlap, which is a consequence of the overlapping ages of

the studied rocks.

The early Badenianearlier late Badenian (16.515 Ma) pe-

riod was characterized by stress changes, that is clockwise ro-

tation of the maximum horizontal stress. NE-SW oriented ex-

tension acted at the beginning of this period (Fig. 7). Evidence

of its activity is recorded at locations 6d, 23, 64 and 72 (Table

1). The stress configuration is characterized by the oblate

stress ellipsoid (

≥

). The magnitude of

was simi-

lar to

and the magnitude of

was distinctly lower than

and

. This caused a distinct regional extension driven by the

active subduction in front of the Carpathians to the NE of the

tiavnica region (Jiøíèek 1979; Ksi¹¿kiewicz 1960; Vialov

1974). Subvertically oriented

and NW-SE

had close

enough magnitudes to experience the

exchange in the

case of regional stress pulses and/or the subtraction of the

overburden load. Regional stress pulses or the regional stress

interplay with varying vertical forces occasionally activated

strike-slip faulting, (e.g. evidence at location 47, Table 1).

Varying vertical forces were caused by episodes of magma

emplacement (Fig. 2) and/or geologically fast changes of the

stratovolcanic relief, that is fast development of volcanic

cones and fast selective erosion.

Later, during the early Badenianearlier late Badenian,

both mentioned stress configurations, with NE-SW oriented

extension, progressively rotated towards the configuration

with W-E oriented

and subvertical

(Fig. 7). Evidence

of such a changed stress is present at locations 6c, 27, 30, 55,

84b, 84d and 109 (Table 1). Calculated stress ellipsoids from

data allowing the use of the Hardcastle & Hills (1991) meth-

od have either an oblate (sites 6c, 55) or prolate character (

≥

) (site 109). At the end of the early Badenianearli-

er late Badenian period, the stress field further rotated clock-

wise towards the position with NW-SE oriented

and sub-

vertical

(Fig. 7). Evidence of its activity is recorded at

locations 6a, 24a, 27, 68, 74, 75, 88a, 107 (Table 1). Stress

ellipsoids determined from locations 74, 88a and 107 have

prolate shapes. Pulses of the regional stress or its interplay

with varying vertical forces caused the

exchange and

occasionally activated strike-slip faulting, (e.g. evidence at

location 71).

Fig. 7 shows which faults were active during the early

Badenianearlier late Badenian. To discuss its value, we

emphasize that the activity of faults is indicated by field evi-

dence and non activity is indicated by lack of evidence. This

means that, despite the study of all suitable outcrops along

faults in Fig. 7, the small chance remains that some faults

could have been active. The same applies to Figs. 8, 9, 10

described below.

The late Badenianearly Sarmatian (1512.7 Ma) period

was characterized by the lack of distinct regional tectonic

stress changes. Extension was W-E to NW-SE oriented

(Fig. 8) and this dominantly controlled the normal faulting.

The

was vertical. Stress ellipsoids, calculated for loca-

tions 10, 16, 36, 78, 92, 109, 121, 122 and 123 (Table 1) have

shapes close to the plane stress ones (

). The

Fig. 7. Structural scheme of the central zone of the tiavnica stra-

tovolcano (after tohl et al. 1990b) with paleostress configurations

and faults determined as active during the early Badenian to earli-

er late Badenian. Divergent arrows indicate normal faulting, diver-

gent arrows coupled with convergent ones indicate strike-slip re-

gime. Arrows show the exact orientation of related principal

stresses, unlike the Table 1.

CALCULATIONS OF TECTONIC, MAGMATIC AND RESIDUAL STRESS 29

value is close to the mean of the remaining two. Locations

29, 33, 46, 60 and 98 indicate activity of the strike-slip fault-

ing. The location 33 with NW-SE oriented compression and

NE-SW extension (Table 1) is anomalous. It may indicate

stress rotation inside a strike-slip fault zone as known from

other regions (e.g. Bogen & Seeber 1986; Freund 1970,

1971, 1974; Nur et al. 1986; Ron et al. 1984; Scoti et al.

1991). Fig. 8 shows active faults of this period.

Stress magnitude calculation, using the stress ratio of

Hardcastle & Hills (1991) (equal to 1-R ratio of Angelier

1989) R = 0.4 and

= 0.0062893, yields values of subverti-

cal

= 75.8 MPa, subhorizontal

= 45.7 MPa and

0.47 MPa. The fast removal of the upper parts of the strato-

volcanic cone above the intrusions (Fig. 2ac) would remove

roughly 42.5 MPa of the load during the late Badenianearly

Sarmatian. This would result in subhorizontal

= 45.7 MPa

and

= 0.47 MPa and subvertical

= 33.3 MPa, changing

the stress ellipsoid from the plane stress to an oblate one and

exchanging

and

(locations 29, 46, 60, 98 in Fig. 8).

Thus, the related kinematic regime changed from normal to

strike-slip faulting. During the same time an emplacement of

the granodiorite intrusion occurred.

Taking the density of the load (calculated proportionally to

the content of lava flows, volcanoclastics and intrusives) to

equal 2577 kgm

, the density of magma 2585 kgm

, s

vertical

(load) 33.3 MPa and the magmatic pressure and intrusion

height estimates mentioned earlier, the resulting magmatic

pressure in the top of the intrusion by buoyancy calculation

(Price & Cosgrove 1991) was about 37.7 MPa, nearly equal to

the addition of s

vertical

and the tensional strength of the sur-

Table 1: Paleostress configurations determined from rocks in the Banská tiavnica area. Note that the table does not show the result of a

tectonic event separation. It is a list of available stratigraphies with inventory of all paleostress configurations determined from these

stratigraphies at related locations. The decrease of the number of paleostress configurations in younger stratigraphies is caused by the

fact that they were deformed by a smaller number of tectonic events than older stratigraphies. Note that the orientations of paleostresses

are generalized in order to demonstrate regional correlation. Codes S and H indicate the use of stress inversion software of Sperner et al.

(1993) and Hardcastle & Hills (1991). The numbers in parentheses indicate the ratio of paleostress magnitudes. Divergent arrows indicate

normal faulting, divergent arrows coupled with convergent ones indicate strike-slip regime.

30 NEMÈOK, KONEÈNÝ

and LEXA

rounding rocks. Magmatic pressure further reduced the prin-

cipal stresses in the surroundings of the intrusion and

changed vertical

= 33.3 MPa to

= 4.4 MPa. Said dif-

ferently, taking the determined tensional strength of the gran-

odiorite = 2.54.9 MPa, the magmatic pressure was in most

places larger than the sum of the vertical

and tensional

strength. The change of subhorizontal

= 45.7 MPa and

subhorizontal

= 0.47 MPa is neglectable, only

changes

. The calculated effective subvertical extension was

large enough to open a subhorizontal Svetozár-type vein sys-

tem above the intrusion, that is in accordance with field evi-

dence (e.g. Fig. 6ac). The intrusion was fed for some time

as indicated by the fact that: 1) subhorizontal veins under-

went several overpressure cycles/episodes of reopening con-

trolled by fluids escaping from the intrusion (Fig. 6f) and flu-

ids were of magmatic origin, as shown by the isotopic data of

Háber et al. (1994), 2) subhorizontal veins are cross cut by

veins opened by tectonic W-E oriented extension, which are

later cross cut by subhorizontal veins (Fig. 6d).

Previous calculations show that the hydrostatic stress state,

which opened the anastomosing vein system (Fig. 6ae), the

oldest fracture system, could not have been created by tec-

tonic stress/overburden removal stress/magmatic stress inter-

play. The calculated differential stress was too high to open

fractures without a preferred orientation (see Cosgrove 1995

for details). The anastomosing vein system had to be opened

by residual stress. It could be accumulated either by an

abrupt temperature drop due to fast overburden removal (Fig.

2ac) or by rather fast temperature drop during cooling and

was independent of the stresses discussed above. Using the

equations (4) and (5), the residual stress

∆σ

would be larger

than the strength of rock after the fast uplift of several tens of

meters if there was no cooling of the intrusion. Taking cool-

ing into account the uplift has to be larger. 0.5 % volume loss

of the intrusion due to cooling would be large enough to gen-

erate this residual stress. However, there is a faster process

of developing the residual stress available. Fig. 11a shows

the modeled cooling T-t paths of points at various distances

from the granodiorite intrusion/andesite country rock bound-

ary. The most dramatic temperature changes, during the first

200,000 years after the intrusion, are recorded by the 1 km

thick granodiorite and andesite zones at their boundary (Fig.

11b,c). Wherever fast cooling occurs over a temperature gra-

dient of more than 35 °C, sufficient residual stress is built up

to overcome the tensile strength of the rock (11d).

Later stages of the late Badenianearly Sarmatian (15

12.7 Ma) are characterized by the lack of distinct influence of

either magmatic stress or overburden removal stress. The em-

placement of quartz-diorite porphyry dykes (see Figs. 1, 3, 6b,c)

postdates the granodiorite intrusion. Calculation of the average

width and length of 128 dykes yields 1:5 ratio; 106 m:36 m.

The dykes are inclined at angles of about 5060° (Fig. 3), indi-

cating that they are hybrid structures not formed by 100 % hy-

draulic fracturing (see Price & Cosgrove 1990). Their driving

stress difference is 19.624.5 MPa, i.e. 4T5T, where T

is the tensile strength of the rock. Taking the calculated value of

as roughly equal to 45.7 MPa,

becomes at least 4.3 times

larger than the tensile strength of granodiorite. Fig. 3 shows that

inclined dykes intruded upwards, experiencing decrease of the

Fig. 9. Structural scheme of the central zone of the tiavnica stra-

tovolcano (after tohl et al. 1990b) with paleostress configurations

and faults determined as active during the Sarmatian. For further

explanations see Fig. 7.

Fig. 8. Structural scheme of the central zone of the tiavnica stra-

tovolcano (after tohl et al. 1990b) with paleostress configurations

and faults determined as active during the late Badenian to early

Sarmatian. For further explanations see Fig. 7.

CALCULATIONS OF TECTONIC, MAGMATIC AND RESIDUAL STRESS 31

vertical load, until conditions for the formation of sills were

met. At the junction of the dykes with sills, it has to be valid that

(

)<(T

⊥

), where T

is the strength of the host rock test-

ed in the vertical direction and T

⊥

is the strength of the host rock

tested in the horizontal direction.

The Sarmatian (13.611.5 Ma) period was also character-

ized by W-E to NW-SE oriented extension (Fig. 9), which

produced dominantly normal faulting. However, the stress

ellipsoid shape, calculated for locations 6b, 15, 20, 24b, 69,

77, 80 and 95 (Table 1), changed to oblate, indicating pro-

gressively stronger regional extension. This is supported by

the fact that there was not any significant erosion, except

the local one of 400 m in the horst region (Háber et al.

1997), but rather the addition of new material on top of the

pre-existing structures (Fig. 2e). It should change the ellip-

soid to prolate, but the value of the R ratio decreased to in-

dicate the oblate ellipsoid. Stress determined for location 40

shows the dextral strike-slip reactivation of suitably orient-

ed faults. Fig. 9 shows active faults of this time period.

During the late Sarmatianearly Pannonian (1210.7 Ma)

period, the W-E to NW-SE oriented extension continued, later

during the Pannonian it was replaced by N-S and E-W com-

pression and extension, respectively (Fig. 10, Table 1). Late

Sarmatianearly Pannonian stress ellipsoids calculated for lo-

cations 114, 116 and 117, with NW-SE oriented extension

(Fig. 10, Table 1), have an oblate character, indicating strong

regional extension. Subsequent fast erosion caused the

exchange, reactivating pre-existing normal faults as strike-slip

faults (location 4 in Fig. 10, 117 in Table 1). A younger stress

pattern, resembling the present stress field (e.g. Gutdeutsch &

Aric 1988), was determined from strike-slip faults at locations

4, 117 and 119. Fig. 10 shows active faults of this time period.

Interpretation

During the early Badenianearlier late Badenian (16.515

Ma), the tiavnica stratovolcano composed of andesite lava-,

pyroclastic-, epiclastic flows and related sills and lacoliths

formed (Fig. 2a). The volcanic cone underwent rather fast

erosion. Synchronously, a quartz-diorite body was emplaced

in its center, during the Badenianearly Sarmatian (1512.7

Ma) (Fig. 2b). During the Badenian, the minimum horizontal

stress

rotated clockwise roughly 90° from NE-SW to NW-

SE (Table 1, Fig. 7). Denudation of the stratovolcano contin-

ued contemporaneously with the emplacement of the grano-

diorite body, 10 km in diameter and 7 km high (Fig. 2c). The

rapid cooling experienced by the upper parts of the intrusion

and adjacent zone of the andesite (Fig. 11) accumulated re-

sidual stress which opened an anastomosing pattern of frac-

tures by hydraulic fracturing (Fig. 6ae). The granodiorite

body was not free to contract soon after releasing fluids due

to its confinement provided by 2 mechanisms. The first one

was the inhomogeneous cooling because the intrusion was

welded to the country rock and the net temperature change

was heterogeneous. The second one was imposed by the fact

that granodiorite is a heterogeneous intergrowth of minerals

of different thermal expansion coefficients and elastic con-

tents.

The theoretical time-position zone of fracturation events

(Fig. 11) does not fit the observation. The system of mineral-

ized anastomosing tensional fractures seems to be present

only in the outermost parts of the granodiorite and adjacent

parts of the andesite. It implies that the fracturation took part

during the earlier stages of the cooling and in the areas of the

most dramatic temperature changes. The fracturation would

be allowed during the small initial volume change due to the

release of mineralized fluids. This would be further support-

ed by the magmatic origin of the fluids recorded in these

fractures, discussed earlier. It should be said that the model-

ing in Fig. 11 was done assuming no heat transfer by fluids,

which makes the problem even more complex.

Removal of the overburden above the intrusion by the fast

erosion interfered with the regional stress pattern changing the

normal faulting regime driven by W-E to NW-SE oriented ex-

tension (Table 1, Fig. 7) to the strike-slip faulting regime. The

input of magmatic stresses during the granodiorite emplace-

ment affected the regional tectonic stress/overburden removal

stress interplay, making

stress vertical. Magmatic pressure

occasionally exceeded the sum of the

plus vertical strength.

During related short time periods, this opened a subhorizontal

Svetozár-type vein system cross-cutting the anastomosing vein

system (Fig. 6ac). The

stress become vertical few times as

indicated by NE-SW oriented vertical veins cross-cutting sub-

horizontal veins and being cross-cut by younger subhorizontal

veins (Fig. 6d). Later on, during the late Badenianearly Sar-

Fig. 10. Structural scheme of the central zone of the tiavnica

stratovolcano (after tohl et al. 1990b) with paleostress configura-

tions and faults determined as active during the late Sarmatian to

early Pannonian. For further explanations see Fig. 7.

32 NEMÈOK, KONEÈNÝ

and LEXA

matian (1512.7 Ma), during the continuing denudation of the

volcano, quartz-diorite porphyry sills and dykes were em-

placed (Fig. 2e). Dykes pass to sills at the depth around the top

of the granodiorite body, where the (

) value meets the

condition for sill formation, outlined earlier. Dense dyke spac-

ing, general NE-SW trend and dyke dips around 60° indicate

stronger NW-SE oriented regional tectonic extension, also in-

dicated by paleostress calculations (Table 1). Lack of any ring

and/or radial dykes does not support the collapse origin of the

postulated caldera during the late Badenianearly Sarmatian.

Paleostress calculation also does not indicate the existence

of strong vertical

stress and subhorizontal polydirectional

extension (Table 1). On the contrary, the subsiding central

parts of the stratovolcano were controlled by faulting driven

by regional W-E to NW-SE extension (Table 1, Fig. 8) and

filled by sediment, tuffs, pyroclastic- and lava flows. The

same stress configuration gained progressively stronger ex-

tension during the Sarmatian (13.611.5 Ma) (Table 1, Fig.

9). During this period, the horst structure developed in the

central zone of the tiavnica stratovolcano, synchronous

with renewed andesite activity. The extension controlled NE-

SW trending normal faults and extensional veins migra-

tion paths of the base and precious metal mineralization flu-

ids. Normal faults, dissecting any older structures, were

Fig. 11. (a) The T-t path diagram related to the cooling of the granodiorite intrusion. Each curve is made for a rock point localized at the

specified distance from the intrusion/andesite country rock boundary. Explanation in text. (b-c) The same diagrams for the early stages of

cooling around the intrusive contact. (d) Fracturation events triggered when the residual stress developed by the fast cooling overcomes

the tensile strength of the host rock. Explanation in text.

CALCULATIONS OF TECTONIC, MAGMATIC AND RESIDUAL STRESS 33

detached along the subhorizontal faults, which frequently

utilized pre-existing subhorizontal veins. Fig. 6e shows an

example of the boudinage related to such a reactivation. The

distinct change of the stress configuration in the course of the

Pannonian (11.56.2 Ma) reactivated pre-existing normal

fault pattern as strike-slips (Fig. 10). This event is indicated

by the occurrence of the youngest, ore deposits in the strike-

slip related horse-tail structures at the southern ends of some

pre-existing normal faults.

Discussion and conclusions

The data discussed in this paper indicate distinct regional

tectonic stress changes during the Badenian (Table 1,

Fig. 7). It was the period when the sinistral transpression

characteristic for the NW boundary of the Inner Western Car-

pathians during the early Miocene changed to transtension.

Areas located further towards the hinterland experience the

clockwise rotation of the minimum and maximum horizontal

stresses. The central zone of the tiavnica stratovolcano ex-

perienced this rotation during the Badenian. This rotation in

different areas has been suggested in numerous papers (e.g.

Nemèok et al. 1989; Fodor et al. 1990; Csontos et al. 1991).

The first and last stage of this rotation is also identified by

Mao et al. (1996) in the Rozália Mine.

The authors did not have sufficiently complete stratigraphy

and sophisticated paleostress methods for a complete identi-

fication. This rotation caused the dextral reactivation of the

pre-existing sinistral strike-slip in the area to the N of the

tiavnica stratovolcano (Kováè & Hok 1993) during the

Badenian. Similar rotation, only relatively dated, was indi-

cated by Sasvári & Schmidt (1994) in the Rozália Mine in

Banská Hodrua.

The discussed clockwise

stress rotation can be dated to

the middle-late Badenian by comparison with regional stress

studies (e.g. Nemèok & Lexa 1990; Nemèok et al. 1993). The

NW-SE oriented extension had to be already active during the

late Badenian as indicated by the age of basal transgressive fa-

cies in the surrounding Turiec, iar Depressions and the

Kremnica Graben, which were opened by this extension (e.g.

Gaparik 1980, 1985; Gaparik et al. 1974; Koneèný et al.

1983; Lexa et al. 1979, 1982). The same extension continued

during the Sarmatian, as shown by the northern parts of the

Turiec Depression where the Sarmatian sequence of the

downthrown hanging wall thickens towards the normal fault,

indicating its synsedimentary activity (Nemèok & Lexa 1990).

These authors describe the same evidence from the Horná Ni-

tra Depression. Sarmatian synsedimentary activity of the nor-

mal fault along the western margin of the Turiec Depression is

indicated by coarse clastic horizons, abruptly pinching out in

direction towards the basin (Gaparik 1985; Gaparik et al.

1974). The end of subsidence in these basins can be implied

from the redeposited kaoline clay horizons in upper Pannonian

and Pontian sediments of the Turiec Depression (Kraus 1986).

This tectonic event was also recognized in the Rozália Mine

by Mao et al. (1996). The authors also distinguished a strike-

slip event, which is not dated, so it is difficult to discuss its un-

expected regime.

Another interesting problem is the evolution of the stress

configuration comprising NW-SE oriented extension. This

became progressively more oblate in character during the

Sarmatian (Table 1). This indicates a progressively stronger

regional tectonic extension. If this stress configuration is

driven by plate movements, the timing and orientation of

these extensional events should be in accordance with the

timing and orientation of thrust movements recorded in the

Outer East-Carpathian accretionary wedge. They are in ac-

cordance, because this is the time when the thrusting in the

Outer Western Carpathians ceased (e.g. Buday 1965; Jur-

ková 1971; Vass et al. 1983; Ksi¹¿kiewicz 1960) and re-

mained active only in the Outer Eastern Carpathians during

the Sarmatian and late Sarmatianearly Pannonian (e.g. Via-

lov 1974; Saulea 1969, Jiøíèek 1979), which should increase

the NW-SE extension in the hinterland.

It is interesting to note that no structural evidence has been

found for collapse along the caldera fault. There are no circu-

lar and radial dyke intersections with the surface, which would

indicate a prolate stress with vertical maximum compression

and polydirectional subhorizontal extension. The caldera fault

was specially studied in the field and none of the locations in

its vicinity provide Late Badenianearly Sarmatian paleostress

indicating either gravity collapse or vertical tension (Table 1,

Fig. 7), unlike, for example the stresses determined at Sierra

Negra volcano, Galapagos (Reynolds et al. 1995). Gravity

collapse would have been proved by a stress ratio R equal to

0.9, that is nearly uniaxial compression and perpendicular

polydirectional extension.

During the early Badenianearlier late Badenian (16.515

Ma), the minimum horizontal stress

rotated clockwise

roughly 90° from the NE-SW to NW-SE orientation. Denuda-

tion of the stratovolcano started earlier in this time period and

continued contemporaneously with the emplacement of the

granodiorite body. The rapid cooling of the upper parts of the

granodiorite intrusion accumulated residual stress which

opened the anastomosing vein pattern (Fig. 6ae). The inter-

play between the erosional overburden removal from above

the intrusion and the regional stress changed the normal fault-

ing regime to the strike-slip faulting one, both driven by W-E

to NW-SE oriented extension (Table 1, Fig. 7). Magmatic

stresses during the granodiorite emplacement interfered with

both regional tectonic stress and overburden removal stress,

making

stress vertical and controlling the development of

subhorizontal veins in the areas where the sum of

plus host

rock strength was exceeded (Fig. 6ac).

During the late Badenianearly Sarmatian (1512.7 Ma),

N-S to NNE-SSW striking faults were reactivated as dextral

strike-slip faults, NE-SW ones as normal faults by the stress

comprising NW-SE extension (Fig. 8). Normal faults and

extensional bridges along the strike-slip faults were the

places of the ore deposition.

The same stress configuration gained progressively stron-

ger extension during the Sarmatian (13.611.5 Ma) (Ta-

ble 1). During this period, the horst structure developed in

the central zone of the tiavnica stratovolcano. Normal

faulting along N-S to NE-SW trending faults was dominant

in the central zone of the tiavnica stratovolcano (Fig. 12)

and was the place of the ore deposition.

34 NEMÈOK, KONEÈNÝ

and LEXA

A distinct change of the stress configuration happened dur-

ing the Pannonian (11.56.2 Ma) (Table 1, Fig. 10). Pre-ex-

isting NE-SW normal faults were reactivated as sinistral

strike-slip faults (Fig. 10). Their extensional bridges were

the place of the ore deposition.

Acknowledgements: The follow-up work of MN, after the

end of the Slovak Geol. Survey Project, has been carried un-

der the financial support of the Lise Meitner Fund, Austria,

later by the Alexander von Humboldt Fund. Authors are

grateful for rock mechanics tests made by L. Sterba and S.

Urban. This paper benefited from the help with field work

and discussions with J. Lexa, D. Onaèila, J.-P. Petit, J. tohl,

V. Koneèný, R. Gazdik and number of other people. The ob-

servations shown in Fig. 6 were made during a short field

trip with J. Lexa and J. tohl. We thank L. Fodor, K. Schul-

mann, J. Cosgrove and anonymous reviewers whose com-

ments improved the paper.

References

Anderson J.L. & Smith D.R. 1995: The effects of temperature and

on the Al-in-hornblende barometer. Amer. Mineralogist

80, 245254.

Angelier J. 1989: From orientation to magnitudes in paleostress de-

terminations using fault slip data. J. Struct. Geol. 11, 3750.

Bergerat F. 1989: From pull-apart to the rifting process: the defor-

mation of the Panonian basin. Tectonophysics 157, 271280.

Blundy D.J. & Holland T.J.B. 1990: Calcic plagioclase equilibria

and new amphibole-plagioclase geotermometer. Contr. Min-

eral. Petrology 104, 208224.

Bogen N.L. & Seeber L. 1986: Neotectonics of rotating blocks

within the San Jacinto fault zone, southern California (abs).

EOS 67, 1200.

Buday T. 1965: Die Tektogenese und der Bau der Neogenen Beck-

en der Westkarpaten. In: Buday T., Cícha I. & Sene J. (Eds.):

Miozän der Westkarpaten. GÚD, Bratislava, 169250.

Cermak V. & Rybach L. 1982: Thermal conductivity and specific

heat of minerals and rocks. In: Angenheister G. (Ed.):

Landolt-Bornstein Numerical Data and Functional Relation-

ships in Sci. and Technol., New Ser., Group V 16. Springer-

Verlag, Berlin.

Cloetingh S.A.P.L., Zoetemeijer R. & van Wees J.D. 1995: Tecton-

ics I. Tectonics and basin formation in convergent settings;

thermo-mechanical evolution of the lithosphere and basin

evolution in compressive tectonic regimes. Short Course, Vr-

ije University, Amsterdam.

Cosgrove J.W. 1995. The expression of hydraulic fracturing in

rocks and sediments. In: Ameen M.S. (Ed.): Fractography:

fracture topography as a tool in fracture mechanics and

stress analysis. Geol. Soc. Spec. Publ. 92, 187196.

Csontos L., Tari G., Bergerat F. & Fodor L. 1991: Evolution of the

stress field in the Carpatho-Pannonian area during the Neo-

gene. Tectonophysics 199, 7392.

Csontos L., Nagymarosy A., Horváth F. & Kováè M. 1992: Tertia-

ry evolution of the intracarpathian area: a model. Tectono-

physics 208, 221241.

Fodor L., Marko F. & Nemèok M. 1990: Evolution microtecto-

nique et paléochamps de contraintes du Bassin de Vienne.

Geod. Acta 4, 147158.

Freund R. 1970: Rotation of strike-slip faults in Sistan, southeast

Iran. J. Geol. 78, 188200.

Freund R. 1971: The Hope fault, a strike-slip fault in New

Zealand. N. Z. Geol. Surv. Bull. 86, 149.

Freund R. 1974: Kinematics of transform and transcurrent faults.

Tectonophysics 21, 93134.

Gaparik J. 1980: Geological evaluation of southern part of

Turèianska kotlina Depression. Report, archive GÚD Bra-

tislava (in Slovak).

Gaparik J. 1985: Basic features of geological structure of Horno-

nitrianska and Turèianska kotlina depressions. In: Samuel O.

& Franko O. (Eds.): Sprievodca ku XXV. celotátnemu zjazdu

slovenskej geologickej spoloènosti. GÚD, Bratislava, 5356

(in Slovak).

Gaparik J. et al. 1974: Structural borehole GH-1 (Horná

tubòa). Reg. Geológia Západ. Karpát, 3 (in Slovak).

Gutdeutsch R. & Aric K. 1988: Seismicity and neotectonics of the

East Alpine-Carpathian and Pannonian area. In: Royden L.H.

& Horváth F. (Eds.): The Pannonian Basin: A Study of Basin

Evolution. AAPG Memoir 45, 183194.

Háber M., Jeleò S., Mao L. & Kovalenker V.A. 1997: Modelling

of mineral-forming processes at the Banská tiavnica epither-

mal deposit (Western Carpathians, Slovak Republic). Pro-

ceedings of the 9

IAGOD Symposium, Beijing, China,

August 1218 1994, E. Schweizerbartsche Verlagsbuchhand-

lung, Stuttgart, 21.

Hammarstrom J.M. & Zen E. 1986: Aluminium in hornblende: an

empirical igneous geobarometer. Amer. Mineralogist 71,

12971313.

Hardcastle K.C. 1989: Possible paleostress tensor configurations

derived from fault-slip data in eastern Vermont and western

New Hampshire. Tectonics 8, 265284.

Hardcastle K.C. & Hills L.S. 1991: BRUTE3 and SELECT. Quick-

basic 4 programs for determination of stress tensor configura-

tions and separation of heterogeneous populations of

fault-slip data. Comput. & Geosci. 17, 2343.

Hollister L.S., Grisson G.C., Peters E.K., Stowell H.H. & Sisson

V.D. 1987: Confirmation of the empirical correlation of Al in

hornblende with pressure of solidification of calc-alkaline

plutons. Amer. Mineralogist 72, 231239.

Ibrmajer J. et al. (Ed.) 1989: Geophysical picture of ÈSSR. UUG

Praha.

Jaeger J.C. & Cook N.G.W. 1976: Fundamentals of rock mechan-

ics. J. Wiley & Sons, New York.

Jiøíèek R. 1979. Tectogenetic development of the Carpathian arc

in the Oligocene and Neogene. In: Mahe¾ M. (Ed.): Tectonic

Profiles of the West Carpathians. GÚD, Bratislava, 203214

(in Czech, English summary).

Johnson M.C. & Rutherford M.J. 1988: Experimental calibration

of an aluminium in hornblende geobarometer applicable to

calc-alkaline rocks. EOS 69, 1511.

Jurková A. 1971: Die Entwicklung der Badener vortiefe im Raum

der Mährischen Peorte und Gebiet von Ostrava (eine Rekon-

struktion auf Grund des Studiums des Begraben Paläozoikum

und dessen Neoiden deckgebirges). Geol. Práce, Spr. 57,

155160 (in Slovak, German résumé).

Kantor J. & Ïurkovièová J. 1985: Genetical characteristics of se-

lected mineralizations in the Western Carpathians. Report, ar-

chive GÚD, Bratislava (in Slovak).

Kantor J., Ïurkovièová J., Eliá K., Repèok I., Ferenèiková E.,

Hasková A., Kvarová A., Rúèka I. & Sladková M. 1988: Iso-

topic research of metallogenic processes. Part I: RudnoBre-

hyPukanec area. Report, archive GÚD, Bratislava (in

Slovak).

Koneèný V. & Lexa J. 1984: Regional geological maps of Slova-

kia: Geological map of the Central Slovakia Neogene Volca-

nic Field. Scale 1:100,000. GÚD, Bratislava.

Koneèný V., Lexa J. & Planderová E. 1983: Stratigraphy of the

CALCULATIONS OF TECTONIC, MAGMATIC AND RESIDUAL STRESS 35

Central Slovakia, Volcanic field. Západ. Karpaty, Sér. Geol.

9, 1203 (in Slovak, English abstract and summary).

Kováè P. & Hók J. 1993: The Central Slovak fault system the field

evidence of a strike-slip. Geol. Carpathica 44, 3, 155159.

Kraus I. 1986: Kaolins and kaolinite clays of the Western Car-

pathians: geologicdepositional conditions, mineralogical

composition, genesis and age. DrSc Thesis, University

Comenianae, Bratislava (in Slovak).

Ksi¹¿kiewicz M. 1960: Zarys paleografii polskich Karpat fliszow-

ych. Práce Inst. Geol. 33, 209231.

Lexa J. et al. 1979: Evaluation of the structural borehole LX-5

(Kremnica). Report, archive GÚD, Bratislava (in Slovak).

Lexa J. et al. 1982: Explanations to the sheets 36132 (H. tubna)

and 36134 (Kremnica). Report, archive GÚD, Bratislava

(in Slovak).

Lexa J. & Koneèný V. 1974: The Carpathian volcanic arc: A dis-

cussion. Acta. Geol. Acad. Sci. Hung. 18, 279293.

Lexa J., Koneèný V., Kalièiak M. & Hojstrièová V. 1993: Distribu-

tion of volcanites of Carpatho-Pannonian region in time

space. In: Rakús M. & Vozár J. (Eds.): Geodynamical model

and deep structure of the Western Carpathians. GÚD, Brat-

islava, 5769 (in Slovak).

Lexa J., tohl J. & Koneèný V. 1999: Banská tiavnica ore dis-

trict: relationship among metallogenetic processes and the

geological evolution of a stratovolcano. Mineralium Deposita

34, 639654.

Linzer H.G., Ratschbacher L. & Frisch W. 1995: Transpressional

collision structures in the upper crust: the fold-thrust belt of

the Northern Calcareous Alps. Tectonophysics 242, 4161.

Mao L., Sasvári T., Bebej J., Kraus I., Schmidt R. & Kalinaj M.

1996: Structurally-controlled vein-hosted mezothermal gold-

quartz and epithermal precious and base metal mineralization

in the Banská Hodrua ore field, Central Slovakia neovolca-

nites. Miner. Slovaca 28, 455490 (in Slovak, English ab-

stract and summary).

Nemèok M., Hók J., Kováè P., Marko F., Madarás J. & Bezák V.

1993: Tertiary tectonics in the Western Carpathians. In:

Rakús M. & Vozár J. (Eds.): Geodynamical model and deep

structure of the Western Carpathians. GÚD, Bratislava,

263267 (in Slovak).

Nemèok M. & Lexa J. 1990: Evolution of the basin and range

structure around the iar mountain range. Geol. Zbor. Geol.

Carpath. 41, 3, 229258.

Nemèok M. & Lexa O. 1995: Structural analysis of the central

zone of the tiavnica stratovolcano in relation to the develop-

ment of vein structures. Report, archive GÚD (in Slovak).

Nemèok M., Marko F., Kováè M. & Fodor L. 1989: Neogene tec-

tonic and paleostress changes in the Czechoslovak part of the

Vienna Basin. Jb. Geol. B.-A. 132, 2, 443458.

Nur A., Ron H. & Scotti O. 1986: Fault mechanics and the kine-

matics of block rotations. Geology 14, 746749.

Peacock S.M. 1990: Computer Exercises for Numerical Simulation

of Metamorphic P-T-t paths. Short Course, Arizona State Uni-

versity, Tempe, 1189.

Pécskay Z., Lexa J., Szakács A., Balogh K., Seghedi I., Koneèný

V., Kovacs M., Márton E., Kalièiak M., Széky-Fux V., Póka

T., Gyarmati P., Edelstein O., Rosu E. & Zec B. 1995: Space

and time distribution of Neogene-Quaternary volcanism in

the Carpatho-Pannonian region. In: Downes H. & Vaselli O.

(Eds.): Neogene and related magmatism in the Carpatho-

Pannonian region. Acta Vulcanol. 7, 2, 1528.

Piwinski A.J. 1975: Experimental studies of granitoid rocks near the

San Andreas fault zone in the Coast and Transverse Ranges and

Mojave Desert, California. Tectonophysics 25, 217231.

Póka T. 1988: Neogene and Quaternary volcanism of the Car-

pathian-Pannonian region: changes in chemical composition

and its relationship to basin formation. In: Royden L.H. &

Horváth F. (Eds.): The Pannonian Basin. A study in basin

evolution. AAPG Mem. 45, 257276.

Price N.J. & Cosgrove J.W. 1991: Analysis of Geological Struc-

tures. Cambridge University Press, 1502.

Ramsay J.G. & Huber M.I. 1983: The techniques of modern struc-

tural geology. Volume I. Academic Press, London, 1307.

Ratschbacher L., Frisch W. & Linzer H.G. 1991a: Lateral extru-

sion in the Eastern Alps, part 2: structural analysis. Tectonics

10, 2, 257271.

Ratschbacher L., Merle O., Davy P. & Cobbold P. 1991b: Lateral

extrusion in the Eastern Alps, part 1: boundary conditions and

experiments scaled for gravity. Tectonics 10, 2, 245256.

Reynolds R.W., Geist D. & Kurz M.D. 1995: Physical volcanology

and structural development of Sierra Negra volcano, Isabela

Island, Galapagos archipelago. GSA Bull. 107, 13981410.

Ron H., Freund R. & Garfunkel Z. 1984: Block rotation by strike-

slip faulting: structural and paleomagnetic evidence. J. Geo-

phys. Res. 89, 6 2566 270.

Royden L.H. & Báldi T. 1988: Early Cenozoic tectonics and paleo-

geography of the Pannonian and surrounding regions. In:

Royden L.H. & Horváth F. (Eds.): The Pannonian Basin. A

study in basin evolution. AAPG Mem. 45, 116.

Royden L.H., Horváth F. & Burchfiel B.C. 1982: Transform fault-

ing, extension and subduction in the Carpathian-Pannonian

region. Geol. Soc. Amer. Bull. 73, 717725.

Royden L.H., Horváth F., Nagymarosy A. & Stegena L. 1983: Evo-

lution of the Pannonian basin system. 2. Subsidence and ther-

mal history. Tectonics 2, 91137.

Royden L.H., Horváth F. & Rumpler J. 1993: Evolution of the Pan-

nonian basin system. 1. Tectonics. Tectonics 2, 6390.

Salters J.M., Hart S.R. & Panto G. 1988: Origin of late Cenozoic

volcanis rocks of the Carpathian arc, Hungary. In: Royden

L.H. & Horváth F. (Eds.): The Pannonian Basin. A study in

basin evolution. AAPG Mem. 45, 279292.

Sãndulescu M. 1988: Cenozoic tectonic history of the Car-

pathians. In: Royden L.H. & Horváth F. (Eds.): The Pannon-

ian Basin. A study in basin evolution. AAPG Mem. 45,

1726.

Sasvári T. & Schmidt R. 1994: Some results of the structural-geo-

logical evaluation of Svetozár vein, 16th level, Rozália mine,

Banská Hodrua. In: Proceedings of the Conference The

latest knowledge from investigation, exploration, exploitation

and treatment of precious ores in Slovakia, 1. July 1994,

Hodrua-Hámre, 5871 (in Slovak).

Saulea E. 1969: Atlas litofacial. VI. Neogen. From: Jiøíèek 1979.

Tectonical development of Carpathian arc in Oligocene and

Neogene. In: Mahe¾ M. (Ed.): Tectonic profiles in the West

Carpathians. GÚD, Bratislava, 205215 (in Slovak).

Scoti O., Nur A. & Estevez R. 1991: Distributed deformation and

block rotation in three dimensions. J. Geophys. Res. 96, 12

22512 243.

Schmidt M.W. 1992: Amphibole composition in tonalite as a

function of pressure: an experimental calibration of the Al-

in-hornblende barometer. Contr. Mineral. Petrology 110,

304310.

Sperner B., Ratschbacher L. & Ott R. 1993: Fault-striae analysis:

a TURBO PASCAL program package for graphical presen-

tation and reduced stress tensor calculation. Comput. &

Geosci. 19, 13611388.

Stegena L., Geczy B. & Horváth F. 1975: Late Cenozoic Evolu-

tion of the Pannonian Basin. Tectonophysics 26, 7191.

Suppe J. 1985: Principles of Structural Geology. Prentice-Hall,

Inc., Englewood Cliffs 1537.

efara J. et al. 1976: Geophysical research of the Central Slovak

neovolcanite basement. Report, archive of Geofond Bratislava,