Bulletin of the Czech Geological Survey, Vol. 77, No. 2, 135–145, 2002
© Czech Geological Survey, ISSN 1210-3527
Trace elements and growth patterns in quartz: a fingerprint of the evolution
of the subvolcanic Podlesí Granite System (Krušné hory Mts., Czech Republic)
AXEL MÜLLER1 – ANDREAS KRONZ2 – KAREL BREITER3
1Natural
History Museum, Dept. Mineralogy, Cromwell Road, London SW7 5BD, United Kingdom; e-mail: [email protected]
2Geowissenschaftliches Zentrum Göttingen, Goldschmidtstr. 1, D-37077 Göttingen, Germany
3Czech Geological Survey, Geologická 6, 152 00 Praha 5, Czech Republic
A b s t r a c t . The Podlesí Granite System (PGS) in the Western Krušné hory Mts., Czech Republic, represents a suite of late-Variscan, highly fractionated rare-metal granites. Based on textural studies and cathodoluminescence five igneous quartz populations can be distinguished in the stock granite and the more evolved dyke granite hosting line rocks (layered granites). Trace element profiling by electron probe micro-analysis (EPMA) gives
evidence for three main crystallisation stages: (1) the zoned quartz phenocrysts representing the early stage of magma evolution in the middle crust,
(2) the stockscheider quartz and groundmass quartz of the stock granite reflecting the subvolcanic solidification conditions of the stock granite, and
(3) the zoned snowball quartz and comb quartz of the dyke granite crystallised from a highly evolved, residual melt. Ti and Al in quartz show a general temporal trend reflecting the evolution of the magma: decrease of Ti and increase of Al. The increase of lithophile elements (Li, Na, Al, P, K) and
of the water content in the magma, the decrease of Ti, crystallisation temperature and pressure are assumed to be predominantly responsible for the
trend. High Al in quartz (>500 ppm) may also result from high crystal growth rate. The textural varieties of the PGS in general are not a product of
metasomatic overprinting, but a product of crystallisation of different fractionated melts.
K e y w o r d s : granite, Krušné hory Mts., quartz population, magma evolution, residual melt, trace elements, zoned quartz
Introduction
Quartz crystals in igneous rocks record and archive
stages of magma evolution which can be decoded by
means of understanding the trace element signature, intragranular growth pattern, crystal habit and crystal size distribution.
This study concerns the igneous quartz populations of
the Podlesí Granite System (PGS), which is located in the
Western Krušné hory Mts., Czech Republic. The PGS represents highly fractionated rare-metal granite and provides
one of the most impressive exposures of rock assemblages
including line rocks (layered granites) developed at the interface between late-magmatic and hydrothermal processes.
Line rocks found at the roof of shallow felsic intrusions are often associated with rare-metal mineralisation
(e.g. Shannon et al. 1982, Kirkham and Sinclair 1988,
Lowenstern and Sinclair 1996, Balashov et al. 2000, Reyf
et al. 2000, Breiter 2002). The rock type is characterised
by repeated bands of coarse-grained, euhedral quartz
and/or K-feldspar separated by bands of fine-grained granite. Shannon et al. (1982) used the metallurgical term “unidirectional solidification textures” (UST) to describe this
layered texture that forms by growth inward from, and perpendicular to, the outer edge of the intrusion.
A central question regarding the genesis of such
evolved subvolcanic systems is whether they obtained
their chemical and textural signature during the magmatic
stage or during subsolidus modification of an igneous or
foreign fluid phase or a mixture of both. In recent years the
understanding of the complex granite system of Podlesí
became progressively more appreciated as studies of
whole rock, major and accessory minerals, rock and mineral physics, and melt inclusions accumulated a great body
of new data (Breiter et al. 1997, Chlupáčová and Breiter
1998, Táborská and Breiter 1998, Breiter 2001).
To further contribute to this knowledge we studied
trace elements in quartz by electron probe micro-analysis
(EPMA) and intra-granular growth patterns by cathodoluminescence (CL) to understand the textural and compositional variation of the system, and to distinguish between
magmatic and metasomatic processes. Cathodoluminescence (CL) is a sensitive method for revealing growth zoning and different populations of igneous quartz (e.g.
Schneider 1993, Watt et al. 1997, D’Lemos et al. 1997,
Müller et al. 2000). Recently, EPMA proves to be an effective in situ micro-beam technique to obtain analysis of
Ti and Al in zoned quartz, because of the high spatial resolution down to 5 µm and the capability of combining the
position of spot analyses with CL imaging (Müller and
Seltmann 1999, Müller et al. 2000, Van den Kerkhof et al.
2001, Müller et al. 2002).
Geology, petrography and whole rock chemistry
The Podlesí Granite System (PGS) is a tongue-shaped
body in the late-Variscan Eibenstock-Nejdek pluton
(Western Krušné hory Mts.; Fig. 1). The top of the granite
cupola outcrops over an area of about 0.1 km2. The internal structure of the copula has been studied through
drilling by the Czech Geological Survey (Lhotský et al.
1988, Breiter 2001) and is well exposed in an abandoned
quarry.
135
Axel Müller – Andreas Kronz – Karel Breiter
Fig. 1. a – Geological map of the Western Krušné hory/Erzgebirge with the distribution of late-Variscan granites and the location of the Podlesí Granite
System (PGS); b – Geological cross-section of the PGS from Breiter (2001).
The PGS intruded Ordovician phyllites and biotite
granite ~320 Ma ago (Förster 2001). General petrologic
and geochemical characteristics of the PGS are reported
by Breiter et al. (1997) and Breiter (2001).
The PGS is mainly formed of albite-protolithionitetopaz granite, named traditionally as stock granite (Lhotský et al. 1988, Breiter et al. 1997). Protolithionites (in
sense of Weiss et al. 1993) are Li-rich micas but poorer in
Li and richer in Fe than zinnwaldite. The uppermost 30–40
m of the stock granite is fine-grained and porphyritic,
whereas the main part of the stock consists of mediumgrained granite (Fig. 2a). The sharp contact of the granite
cupola with the phyllites is bordered by a marginal, 50 cm
thick pegmatite (stockscheider; Fig. 2b). The stock granite
is peraluminous (A/CNK 1.15–1.25), enriched in incompatible elements such as Li, Rb, Cs, Sn, Nb, W and in
phosphorus and fluorine (Breiter et al. 1997, Breiter 2001,
Table 1).
Within the depth of 40–100 m the stock granite is intercalated with sub-horizontal layers of albite-zinnwalditetopaz granite (Fig. 2c). This leucocratic, fine-grained dyke
granite contains K-feldspar, small isometric quartz crystals,
and inclusion-rich topaz that are embedded in fine-grained
albite tablets. Borders of larger K-feldspars are often replaced by quartz, albite, and topaz. The dyke granite is more
evolved than the stock granite and is relatively enriched in
Al (A/CNK 1.2–1.4), P, F, Na, Rb, Li, Nb, and Ta.
Unidirectional solidification textures (UST) are developed in the upper 50 cm of the major layer of dyke granite, which has a thickness of about 6 m. The UST consist
136
of alternating laminae of dyke granite with inhomogeneous chemical composition (Breiter 2001) and of oriented fan-like zinnwaldite with comb quartz (in sense of
Shannon et al. 1982; Fig. 2d). Locally, ongrowths of oriented K-feldspar megacrysts (6 cm) on the zinnwalditecomb quartz layers are developed forming pegmatite-like
bands. Directional growth of pegmatitic comb quartz,
zinnwaldite, and K-feldspar indicates downward solidification.
The samples investigated were mostly collected in the
quarry of Podlesí (Fig. 1b). The stockscheider was sampled at the contact with the host rock, 150 m SE from the
quarry.
Methods
Trace element abundances of Al, Ti, K, and Fe in
quartz were performed with a JEOL 8900 RL electron microprobe at the GZG Göttingen equipped with 5 WD detectors and with a CL detector using an extended
wavelength range from 200 to 900 nm.
SEM-CL images were collected from the JEOL system
using slow beam scan rates of 20 s at processing resolution
of 1024 × 860 pixels and 256 grey levels. The electron
beam voltage and current was 30 kV and 200 nA, respectively. CL imaging prior to and after EPMA analysis allowed the location of the sampling spots in relation to the
CL structures.
The following standards were used for the analysis:
Trace elements and growth patterns in quartz: a fingerprint of the evolution of the subvolcanic Podlesí Granite System (Krušné hory Mts., Czech Republic)
location
quarry
Podlesí
Sample No.
709
major elements (wt%)
SiO2
73.57
TiO2
0.04
Al2O3
14.93
Fe2O3
0.15
FeO
0.66
MnO
0.019
MgO
0.05
CaO
0.27
Li2O
0.24
Na2O
3.81
K2O
3.97
Rb2O
0.18
P2O5
0.43
CO2
0.01
F
1.35
H2O+
1.12
H2O0.10
F=O
–0.57
total
100.54
trace elements (ppm)
Ba
n.a.
Cs
103
Hf
2
Nb
47
Sn
62
Sr
47
Ta
19
Th
6
U
11
W
21
Zr
28
150 m SE of
quarry Podlesí
3361
pegmatite-like
band of the UST
dyke granite
pegmatite
(stockscheider)
stockgranite
Rock
Table 1. Major and trace element data of representative whole rock samples from the PGS (from Breiter 2001).
quarry
Podlesí
3413
quarry
Podlesí
3417
74.78
0.05
13.80
0.41
0.46
0.03
0.05
0.43
0.03
3.60
4.23
0.07
0.40
0.03
0.50
0.88
0.23
–0.21
99.86
72.52
0.01
15.98
0.01
0.55
0.04
0.04
0.17
0.29
4.89
3.51
0.22
0.52
0.01
1.33
0.90
0.14
-0.56
100.83
66.54
0.02
17.79
0.15
0.94
0.05
0.02
0.39
0.44
2.90
6.42
0.30
0.80
0.01
1.72
1.55
0.12
-0.72
99.80
23
56
2
58
20
12
19
8
43
24
31
9
159
2
54
35
53
20
6
13
23
25
12
198
2
79
31
87
54
12
29
55
21
n.a. – not analysed
synthetic Al2O3 (52.92 wt% Al), orthoclase, Lucerne,
Switzerland (12.18 wt% K), synthetic TiO2 (59.95 wt%
Ti), and haematite, Rio Marina, Elba (69.94 wt% Fe). Raw
analyses were converted into concentrations, after making
appropriate matrix corrections using the phi-rho-z method
of Armstrong (1991). For high precision and sensitivity, an
accelerating voltage of 20 kV, a high beam current of 80
nA, a beam diameter of 5 µm, and counting times of 15 sec
for Si, and of 300 sec for Ti, K, and Fe were chosen.
Counting times for the lower background and for the upper background were 150 sec for these elements. Due to
the strong drift of the background signal in the low energy
range near the Al Kα-line during long counting times the
total acquisition of Al x-ray counts was separated in 5 individual steps with a 60 sec peak, 30 sec lower background
and 30 sec upper background. This procedure minimises
the systematic error coming from a drift of the analysed
rays due to beam induced damage of the quartz. The higher energies of the K Kα, Ti Kα and Fe Kα lines are not
showing any significant shift of the background signals
during the counting time.
Limits of detection (LOD) were 15 ppm for Ti, 11 ppm
for K, 15 ppm for Fe, 51 ppm for Na, and 10 ppm for Al.
LOD’s were calculated from 77 background measurements with 95% confidence utilising the Student’s t-distribution (see Appendix).
Quartz populations
Based on CL and textural studies (crystal habit, crystal
size) five igneous quartz populations can be distinguished in
the PGS: phenocryst quartz (I), groundmass quartz (II), and
stockscheider quartz (III) in the stock granite and isometric
quartz (IV) and comb quartz (V) in the dyke granite.
The stock granite contains rare euhedral quartz phenocrysts (I, 1–5 mm) showing complex growth zoning contrasted with red-brown, violet, and blue CL colours (Fig.
3a). Generally, the phenocrysts have a rounded, red-brown
luminescent core. The core is overgrown by blue luminescent oscillatory growth zones. Marginally, the phenocrysts
are strongly embayed. The zoning fits the shape of the lobate embayments. Fig. 4a summarises schematically the
structural properties of the phenocryst quartz.
The phenocrysts are overgrown by anhedral finegrained groundmass quartz (II). The groundmass quartz is
free of growth zoning and shows weak red-brown CL. Secondary structures are bright halos around radioactive micro-inclusions and non-luminescent, neocrystallised
domains. Like the groundmass quartz, the stockscheider
quartz (III) is free of growth zoning, shows red-brown luminescence, and contains a number of secondary structures (Fig. 3b).
In the dyke granite and fine-grained layers of the line
rock, the quartz forms snowball-textured phenoblasts (IV)
0.2 to 1.5 mm in size (Fig. 3d; Müller and Seltmann 1999).
Generally, the snowball quartz has a red- to red-brown CL,
whereby bright grey growth zones in the SEM-CL image
correspond with zones having a brighter reddish to orange
luminescence. The snowball quartz shows continuous
growth into the matrix quartz, recognizable by the ramified,
dendritic grain boundaries, and penetrates the matrix. The
crystals contain common inclusions of the groundmass minerals, such as corroded K-feldspar, mica, metamict zircon,
apatite and albite. Fluid and melt inclusions are abundant
throughout quartz. Large numbers of albite crystals envelope the phenoblast edge indicating that the quartz pushed
the albite tablets away during growth. The phenoblasts show
oscillatory growth zoning characterised by planar bordered
growth zones with α-quartz habit. The zoning continues into the amoebic crystal margin and into the matrix quartz
without changes of the CL properties. Fig. 4b shows the
scheme of the structural properties of the snowball quartz.
Moreover, the figure illustrates the structural contrasts be-
137
Axel Müller – Andreas Kronz – Karel Breiter
Fig. 2. Hand specimen of textural granite varieties of the PGS. a – Stock granite with rare quartz phenocrysts (arrow); b – Marginal stockscheider at
the contact to the Ordovician phyllites. The stockscheider is fragmented by a fine-grained variety of the stock granite; c – Leucocratic dyke granite
containing common, small isometric quartz drops (arrow); d – Unidirectional solidification textures (UST) in the upper part of the major dyke granite
layer. Locally, oriented comb quartz crystals (arrow) and K-feldspar megacrysts are arranged along fan-like zinnwaldite laminas.
tween quartz phenocrysts (I) and snowball phenoblasts
(IV). The comb quartz (V) nucleated at the fan-like zinnwaldite layers of UST exhibits similar growth patterns and
CL colours to the snowball quartz (Fig. 3c).
A late- to post-magmatic fluid-driven overprint (e.g.
micro fracturing and greisening) causes small-scale dissolution, precipitation and re-equilibration of pre-existing
quartz (I–V) along grain boundaries, intra-granular microcracks, and around fluid inclusions resulting in neocrystallised quartz (VI). The neocrystallised quartz has the
same crystallographic orientation as the host quartz.
Trace elements in quartz
Locations of the trace element profiles determined by
EPMA are shown in Figure 5 and concentrations are compiled in Table 2.
The highest Ti content up to 88 ppm is bounded by
blue luminescent growth zones of quartz phenocrysts (I).
In contrast, the red-brown luminescent phenocryst core
has low Ti, down to 35 ppm. The association of high Ti
with blue luminescent quartz is in agreement with previous studies, which proved a relation between high Ti and
blue luminescence (Van den Kerkhof et al. 1996, Müller et
al. 2000, Müller et al. 2002). Ti distribution in the younger
groundmass quartz and stockscheider quartz is more homogeneous than in the phenocryst quartz. The average
content is 43 ppm for the groundmass quartz and 41 ppm
for the stockscheider quartz, which is similar to the composition of the phenocryst core. Snowball quartz and comb
quartz of the dyke granite has low average Ti contents of
26 ppm and 22 ppm, respectively.
In contrast to Ti, the Al content in the phenocryst quartz
is relatively constant and low in comparison with those in
other quartz populations. Al in the groundmass and
→ Fig. 3. Scanning electron microscope cathodoluminescence images of quartz populations (I–V) in the PGS. a – Zoned phenocryst (I) in the stock
granite overgrown by a thin film of groundmass quartz (II) which is free of growth zoning. A number of mica crystals (black) were included during
growth of the phenocryst and hindered the growth; b – Stockscheider quartz (III) with common secondary structures, like oriented, healed micro-cracks
and small non-luminescent halos which surround fluid inclusions; c – Zoned comb quartz (V) of the unidirectional solidification textures; d – Three
examples of zoned, isometric quartz of the dyke granite forming snowball-textured phenoblasts (IV). The CL colours of the snowball quartz are similar to those of the comb quartz.
138
Trace elements and growth patterns in quartz: a fingerprint of the evolution of the subvolcanic Podlesí Granite System (Krušné hory Mts., Czech Republic)
139
Axel Müller – Andreas Kronz – Karel Breiter
Table 2. Trace element concentrations (in ppm) of the quartz populations within the PGS. Na were always below the limit of detection, so we do not
include it in the table. I – phenocryst quartz, II – groundmass quartz, III – stockscheider quartz, IV – snowball quartz, V – comb quartz, VI – neocrystallised quartz
#
I-01
I-02
I-03
I-04
I-05
I-06
I-07
I-08
I-09
I-10
I-11
I-12
I-13
I-14
I-15
I-16
I-17
Al
279
204
332
220
291
265
277
322
283
318
320
287
338
318
345
445
354
Ti
88
59
73
73
76
77
81
88
72
61
73
35
49
44
41
45
40
K
108
57
135
136
126
78
101
103
88
64
110
84
88
65
110
159
109
Fe
51
<15
32
30
40
31
21
50
26
21
33
34
53
41
50
90
70
II-01
II-02
II-03
II-04
II-05
459
393
416
356
309
17
41
50
57
50
109
182
162
112
105
18
58
29
44
35
III-01
III-02
386
384
47
37
88
86
26
19
#
III-03
III-04
III-05
III-06
III-07
III-08
III-09
III-10
III-11
Al
476
379
524
434
243
482
330
360
383
Ti
46
50
43
43
33
37
47
38
33
K
92
68
187
91
27
152
41
27
84
Fe
28
31
44
44
18
51
<15
17
34
IV-01
IV-02
IV-03
IV-04
IV-05
IV-06
IV-07
IV-08
IV-09
IV-10
IV-11
IV-12
IV-13
IV-14
IV-15
IV-16
530
306
452
452
981
578
582
591
1289
686
896
936
819
786
343
414
27
25
<15
<15
16
30
30
40
41
<15
34
31
27
28
25
17
38
22
30
25
53
28
33
52
37
37
35
51
37
35
20
26
17
<15
<15
<15
23
<15
16
20
41
40
81
48
24
17
21
17
#
IV-17
IV-18
IV-19
IV-20
IV-21
IV-22
Al
806
1033
453
739
354
480
Ti
26
29
24
39
16
32
K
75
33
29
31
25
27
Fe
18
<15
15
<15
17
23
V-01
V-02
V-03
V-04
V-05
V-06
V-07
V-08
V-09
V-10
V-11
610
372
338
843
626
493
343
355
441
307
780
21
16
22
29
15
23
16
32
15
21
27
59
183
67
193
73
103
74
58
92
73
149
16
16
<15
22
18
18
14
<15
18
24
<15
VI-01
VI-02
VI-03
VI-04
VI-05
VI-06
VI-07
62
57
113
89
52
77
148
23
<15
<15
16
<15
<15
26
56
53
80
83
53
51
78
74
51
126
152
76
37
53
populations (I–V) yields a trend
reflecting the evolution of the
magma (Fig. 6a). The early crystallisation stage represented by
the quartz phenocrysts is characterised by high Ti and low Al
concentrations in the quartz lattice. During further evolution Ti
decreases whereas Al increases.
The post-magmatic neocrystallised quartz in healed structures does not follow the
magmatic trend.
The plot of Al versus K
shows two different trends (Fig.
6b). The Al/K ratio in the magmatic quartz of the stock granite
(I–III) is about 1 : 0.3, whereby
Fig. 4. Schematic growth pattern of quartz phenocrysts (a) and snowball-textured phenoblasts (b).
the K concentration is up to 187
ppm. In the snowball quartz (IV)
the Al/K ratio amounts 1 : 0.05.
stockscheider quartz is also homogeneously distributed but The K of the snowball quartz is generally below 50 ppm.
it has somewhat higher average concentrations of 387 The comb quartz shows Al/K ratios between both trends.
Phenocryst quartz and snowball quartz show a positive
and 349 ppm, respectively. Snowball quartz as well as
comb quartz shows a wide variation of Al from 306 ppm correlation between the Ti/Fe and the K/Fe concentration
up to 1289 ppm. The contents of Al and Ti show no rela- ratio (Fig. 6c). The ratios of the neocrystallised quartz also
tion to the zoning in the snowball quartz, in contrast to display this correlation and, therefore, it can be concluded
the phenocryst quartz, where the zoning is related to the that this correlation is not associated with magmatic evoTi concentration. The non-luminescent quartz of healing lution. The Ti/Fe and the K/Fe ratios of groundmass
structures (neocrystallised quartz VI) is depleted in Al quartz, stockscheider quartz, and comb quartz plots more
scattered due to certain spotty high concentrations of K in
and Ti.
The plot of Al and Ti contents of the igneous quartz these quartz generations.
140
Trace elements and growth patterns in quartz: a fingerprint of the evolution of the subvolcanic Podlesí Granite System (Krušné hory Mts., Czech Republic)
Fig. 5. Location of trace element profiles measured by EPMA. a – Phenocryst quartz overgrown by groundmass quartz. The profile starts in the redbrown luminescent, unzoned groundmass quartz (A), crosses the blue luminescent zones of the phenocryst and ends in the red-brown luminescent phenocryst core (B). The growth zoning of the phenocryst adapts to the shape of lobate embayments, right of the image; b – Stockscheider quartz with
common non-luminescent, patchy halos of secondary quartz (black) and healed micro-cracks; c – Snowball quartz; d – Comb quartz.
Discussion
Interpretations of the observed trends of decreasing Ti
and increasing Al in quartz during igneous evolution are
difficult because the petrogenetic significance of trace elements in quartz is poorly known. First, Stuttner and
Leininger (1972) point out that quartz from different magmatic sources has different Ti contents. Recently, Gurbanov et al. (1999) and Larsen et al. (2000) show that the
trace element distribution in igneous quartz may follow
the petrogenetic history of the quartz-forming melt. Clearly, the trend is related to magmatic evolution, but which
parameters control the incorporation of Ti and Al into the
quartz lattice?
The significant decrease of Ti and increase of Al in
quartz is in contrast to the nearly constant Ti and Al content of the whole rock (see Table 1). Of course, there is a
relative increase of Al and decrease of Ti in the more
evolved dyke granite in comparison to the stock granite.
However, the change of the Ti and Al content in quartz is
much more drastic.
The parameters that control Al and Ti incorporation are
different for both elements, because they behave in a different way. Al and Ti are the most common trace elements
in quartz, which substitute for Si. In contrast to Ti4+, which
is independent of the activity of charge compensators, Al3+
needs charge compensators, such as Li+, H+, K+, or Na+ to
balance the substitution. These monovalent ions enter interstitial positions in channels running parallel to the
c-axis (e.g. Bambauer 1961, Dennen et al. 1970, Lehmann
and Bambauer 1973, Weil 1984). The berlinite substitution
(2Si4+ = Al3+ + P5+) known in feldspars (Simpson 1977)
has not been found in natural quartz.
High Ti concentrations are typical for quartz formed at
temperatures >500 °C and at high pressure, such as in
granulites (Blankenburg et al. 1994, Van den Kerkhof and
Müller 1999). Thomas (1992) calculated the maximum
crystallisation depth of quartz phenocrysts in tin granites
of the Krušné hory/Erzgebirge of about 21 km, based on
microthermometric studies of silicate melt inclusions.
Therefore, high crystallisation temperature and pressure
during the early stage of the magma crystallisation are responsible for the high average Ti in the quartz phenocrysts.
The variation of Ti between single growth zones may result
141
Axel Müller – Andreas Kronz – Karel Breiter
250
100
b
a
90
200
80
60
K (ppm)
Ti (ppm)
70
50
40
150
100
30
50
20
10
0
0
0
200
400
600
800
1000
1200
1400
0
200
400
600
800
1000
1200
1400
Al (ppm)
Al (ppm)
12
phenocryst quartz I
stockscheider quartz III
comb quartz V
c
10
Fig. 6. a – Plot of Al versus Ti concentrations of quartz populations
(I–VI) present in the PGS. Note, that the neocrystallised quartz (VI) does
not follow the trend of the igneous quartz populations; b – Plot of Al versus K concentrations of quartz; c – Plot of Ti/Fe versus K/Fe concentration ratio.
8
K/Fe
groundmass quartz II
snowball quartz IV
neocrystallised quartz VI
6
4
2
0
0
1
2
3
4
Ti/Fe
from crystal fluctuation and ascent, whereby the crystal is
exposed to different temperatures, pressures, as well as to
variation of the Ti content in the melt.
Both the groundmass quartz (II) and the stockscheider
quartz (III) have similar, moderate Ti and Al concentrations indicating that both populations crystallised contemporaneously from the same magma.
In the following we discuss the causes of the high Al
in the snowball quartz (IV) and comb quartz (V) which
occur in the dyke granite. Both quartz populations have
similar trace element signature. Furthermore, similar CL
properties and growth patterns support the assumption that
both populations crystallised from the same fluid-enriched
magma pulse. This is in contrast to earlier studies by Breiter et al. (1997) who assumes that the comb quartz of the
pegmatite like bands of the line rock crystallised later than
the dyke granite host rock and form a different fluid-enriched melt. The crystal size and crystal habit contrast between snowball quartz and comb quartz may be explained
by different nucleation processes. In contrast to Lowenstern and Sinclair’s (1996) model for comb quartz formation, both the comb quartz and the snowball quartz grew
from melt. The melt is preserved as melt inclusions in both
quartz populations (see also Breiter et al. 1997).
The snowball texture is typical for highly evolved al-
142
kali feldspar and topaz-bearing granites worldwide and its
origin is controversial. It is considered to be either metasomatic (e.g. Beus et al. 1962, Sonyushkin et al. 1991), or
magmatic in origin (Kovalenko 1977, Poutiainen and
Scherbakova 1998, Müller and Seltmann 1999).
According to Holten et al. (1997) the observed zonal
texture of the snowball quartz and comb quartz mostly reflects external fluctuations in an open system (e.g., degassing) and is not a product of self-organisation.
Zinnwaldite layers partially envelope the edge of comb
quartz and of K-feldspar megacrysts indicating that these
large crystals put off the zinnwaldite layers during growth
in a viscous environment (see Fig. 2d). The trigonal habit of
the zoning, the frequently enclosed groundmass albites and
the occurrence of melt inclusions indicate a growth of the
snowball quartz of the PGS in a nearly non-convecting and
fluid-saturated crystal mush at a temperature of <600 °C (at
<1 kbar). To reach such low temperatures, particularly at
the shallow depth at which the PGS is emplaced fluxing
agents like Li, B, F and water must be added to the magma.
High F (2.2 wt%), high P2O5 (2.6 wt%), and about 7.5±0.8
eq.-wt% water determined from melt inclusions in the dyke
granite (Breiter et al. 1997) should have dramatically suppressed the solidus temperature (e.g. Manning 1981). The
extrapolated solidus temperature (Thomas et al.1996) is
610 ± 26 °C for the melt of the dyke granite, which is in
agreement with our observations.
The occurrence of ghost K-feldspar phenocrysts and mica in the dyke granite with a similar element signature and
CL characteristics to the K-feldspar and mica in the stock
granite (Breiter 2001) provides evidence that quartz phenocrysts were also originally present in the dyke granite, but
Trace elements and growth patterns in quartz: a fingerprint of the evolution of the subvolcanic Podlesí Granite System (Krušné hory Mts., Czech Republic)
they were completely cannibalised. P-enriched rims of Kfeldspar and Li-enriched rims of mica found only in the
dyke granite document a distinctly more evolved environment enriched in phosphorus and fluorine. Other features
that significantly support subsolidus processes are (1) the
disappearance of “common” accessory minerals in the
dyke granites (Förster 2001) and (2) the resetting of the
oxygen isotope thermometers to lower submagmatic temperatures (Žák et al. 2001). The latter is a result of the activity of high-δ18O fluids, whereas low-δ18O fluids are not
involved in this process.
The increase of Al in the snowball quartz and comb
quartz may by related to the evolution of magma composition, e.g., to the increasing content of lithophile elements, particularly Li+, K+, and Na+, which are the main
charge compensating ions of substitutional Al3+. Beside
these elements, hydroxyl groups and adsorbed H2O act as
charge compensators for [2AlO4/M+]- defects, where M+
is a combination of Li, K, and Na ions (Bambauer et al.
1963, Maschmeyer and Lehmann 1983, Kronenberg et al.
1986, Stenina 1995). Therefore, the uptake of Al into the
quartz lattice may be partially controlled by the presence
of Li+, K+, and Na+ and water in the melt. Substitutional
P5+ may also function as a charge compensator of substitutional Al3+ as is known from feldspar (Simpson 1977).
Therefore, the high P5+ content of up to 2 wt% P2O5 (Breiter 2001) in the line rock may also influence the Al3+ content in the quartz lattice. The high Al in the snowball
quartz and in the comb quartz may, therefore, reflect the
enrichment of lithophile elements and water in the magma
of the dyke granite.
However, spotty high Al in the snowball quartz and
comb quartz (>500 ppm) suggest that Al is not only structurally incorporated, but that it also occurs as impurity
clusters or inclusions (Flicstein and Schieber 1974,
Blankenburg et al. 1994, Brouard et al. 1995, Götze et al.
1999, Mullis and Ramseyer 1999, Müller et al. 2002).
Caused by the larger ion radius of Al3+ (r = 0.47 Å) in
comparison to Si4+ (r = 0.34 Å) such high amounts of
structurally incorporated Al3+ should cause an extreme deformation and weakening of the quartz lattice. Structurally incorporated Al in igneous quartz typically ranges up to
400 ppm (Watt et al. 1997, Morgan et al. 1998, Müller et
al. 2000). Surprisingly, analyses of the snowball quartz
with very high Al have low K (Fig. 6b) although K functions as charge compensator for Al. However, the spotty
high Al suggests that Al is accumulated within submicrometer sized inclusions.
In general, high impurity densities in quartz correlate
with fast growth rates in synthetic crystals (Brice 1985).
Such growth rates (up to >3 mm/day) are 50 times faster
than the highest known growth rates for skeletal magmatic high-quartz (Swanson and Fenn 1986). Temperature and
pressure alone do not appear to have a significant affect on
Al substitution in quartz (Scotford 1975, Pavlishin et al.
1978, Stavrov et al. 1979). The Al uptake may be more
sensitive to variations in growth rate. The characteristic
zoning pattern, the frequently enclosed groundmass minerals and the dendritic crystal margins point to rapid crystal growth during multiple degassing pulses.
Another source of Al in the snowball quartz and comb
quartz may be impurity clusters that originate from replaced and dissolved albite crystals during the blastic
growth of quartz.
Low trace element concentrations in neocrystallised
quartz point to a systematic recovery of defect centres and
to trace element redistribution in the quartz lattice, whereby impurities are probably released from quartz during
solid-state neocrystallisation (Müller et al. 2002). Also the
very low CL intensity indicates a low defect concentration
in this neocrystallised quartz.
Conclusion
The examination of quartz populations in the PGS
gives evidence for three main stages of crystallisation:
(1) quartz phenocrysts representing the early stage of
magma evolution, (2) stockscheider quartz and groundmass quartz of the stock granite reflecting the solidification conditions of the stock granite, and (3) the snowball
quartz and comb quartz of the dyke granite that crystallised from a highly evolved, residual melt.
We observe a general trend of the trace element distribution in the igneous quartz populations (I–V) reflecting the evolution of the magma: a decrease of Ti and an
increase of Al.
The increase of the lithophile elements (Li, Na, Al, P,
K) and of the water content in the magma and the decrease of Ti, crystallisation temperature and pressure are
assumed to be predominantly responsible for the trend.
Extremely high Al in quartz may also result from high
crystal growth rate. The Ti-enriched phenocrysts (I) originated in greater depth represent the early less evolved
magmatic stage. Groundmass quartz (II) and stockscheider quartz (III) are the result of crystallisation from the
stock granite in the intrusion level.
The Al-rich snowball quartz (IV) and comb quartz
(V) of the UST crystallised at low temperatures
(<600 °C) and pressures (<1 kbar) from a highly evolved,
fluid-enriched magma. A spotty high Al concentration in
both quartz populations is mainly interpreted as submicrometer sized inclusions incorporated during fast crystal growth. The zoning pattern of the snowball quartz and
comb quartz is explained by periodic degassing during
melt cooling in the granite roof. Highly reactive fluids
exsolved in the late magmatic stage from the melt cause
the resorption and, partially, the complete solution of
quartz and feldspar phenocrysts of the early crystallisation stage. Summarising, the dyke granite in general is
not a product of metasomatic overprinting of the stock
granite, but the product of crystallisation of a different,
more fractionated melt.
143
Axel Müller – Andreas Kronz – Karel Breiter
A c k n o w l e d g e m e n t s . We thank A. M. van den Kerkhof, H.-J.
Behr, and R. Seltmann for discussion and critical comments. We are also
indebted to R. B. Larsen and W. Siebel for reviews and fruitful discussions. This work was funded by the Deutsche Forschungsgemeinschaft
(MU 1717/2-1) and by the German Bundesministerium für Bildung und
Forschung (CZE 00/011).
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Appendix
Limits of detection (LOD) were calculated from 77 background measurements with 95% confidence utilising the Student’s t-distribution according to the following equation (Table 3):
ILOD = tz (P;f) σBG
where ILOD = intensity of the LOD [cts]; σBG = standard deviation of the average background; tz (P;f) = the Student’s t for binomial limitation determined by the confidence P and the degrees of freedom f = n – 2, where n is the number of background measurements. The Student’s t with 95%
confidence amounts 1.99 for 77 background measurements.
For comparison also the common method for calculating LOD, given by:
ILOD = 3 √BG
for an single measurement was estimated.
BG are the total accumulated counts of the background signal. (“LOD by single measurement”). Both methods give comparable results, but we
prefer the random sampling using Student’s t-test. The latter is more sensitive due to systematic errors and the occurrence of drift-phenomena.
The concentration of LOD CLOD is calculated using the factor of the linear calibration of total intensity vs. concentration of an element (see Table 3):
CLOD [ppm] = factor * ILOD
(Due to the constant composition of the matrix SiO2, there would be no change in the ZAF-factors, hence a linear calculation is reliable.)
Table 3. Limits of detection (LOD). The average background (BG) is the mean of the total acquired counts for each element using 150 seconds counting time on each side of the peak (20 kV and 80 nA).
element counting time peak each BG average BG
[s]
[s]
[cts]
Al
300
150
42580
K
300
150
14800
Ti
300
150
33060
Fe
300
150
36040
factor
(lin. calib.)
0.016
0.030
0.028
0.029
LOD by n=77 measurements
ILOD = 1.99 *σBG [cts] CLOD [ppm]
615
10
365
11
540
15
520
15
LOD by single measurement
ILOD = 3 *√BG [cts]
CLOD [ppm]
619
10
365
11
545
15
570
17
145
Bulletin of the Czech Geological Survey, Vol. 77, No. 2, 147–155, 2002
© Czech Geological Survey, ISSN 1210-3527
The compositions of rock-forming and accessory minerals from the Gemeric granites
(Hnilec area, Gemeric Superunit, Western Carpathians)
IGOR BROSKA1 – MICHAL KUBIŠ1 – C. TERRY WILLIAMS2 – PATRIK KONEČNÝ3
1Geological
Institute, Slovak Academy of Sciences, Dúbravská cesta 9, 842 26 Bratislava, Slovakia; e-mail: [email protected]
of Mineralogy, The Natural History Museum, Cromwell Road, London, SW7 5BD, United Kingdom
3Geological Survey of Slovak Republic, Mlynská dolina 1, 817 04 Bratislava, Slovakia
2Department
A b s t r a c t . The Hnilec granites represent a suite of specialized S-type granite massifs in the Western Carpathians, with a primary enrichment in
elements B, F, Sn, Nb, Ta and W being reflected in a suite of characteristic minerals (e.g. tourmaline, cassiterite, and Nb-Ta phases). The highly evolved
nature of these granites results from a lowering of the viscosity of the primary melt, which is due to the high content of the volatiles (B, F), and an increased activity of P. Phosphorus is incorporated into primary apatite as well as in the K-feldspar, whereas the Na-feldspar (albite) is P-bearing only
in the exsolved perthitic part of the K-feldspar. The majority of the perthitic albite crystals contain abundant exsolved apatite grains that in general, do
not exceed 3–5 µm in size and due to the apatite exsolution, albite is low in P. Compositionally, secondary apatite in albite differs from the primary
one in having low concentrations of Mn, Fe and REE. The LREE are predominantly hosted in apatite, and to a lesser extent monazite, whereas the
HREE are mainly distributed in zircon, xenotime and apatite. The morphology of zircon crystals is similar in all the granitic rock types investigated,
i.e. in the deep-seated coarse-grained granites, upper fine-grained granites and the greisenized cupola. Cathodoluminesce images of zircon crystals
show a strong primary magmatic zonation and some secondary alteration features. These aspects of zircon morphology and composition suggest that
zircon crystallized mainly in the deep magma chamber and was incorporated within the ascending differentiated melt. Some of the zircons can be classified as G1 and L subtypes which represents a late magmatic population of zircon. The contribution of the early magmatic zircon to the bulk rock
HREE distribution is considered to be significant. The co-magmatic features of the granites are evident from extreme differentiation leading to the zonality of the plutons and alteration of the granite, i.e. albitization and greisenization, by fluids in the cupola.
K e y w o r d s : granite, two-mica granite, rock-forming minerals, accessory minerals, chemical composition, zircon, boron, fluorine, niobium,
phosphorus, tantalum, tin, greisenization, albitization,Western Carpathians
Introduction
The results of tin prospecting of the Hnilec granite during the 1970’s discovered the existence of a high temperature mineralization within these granites that had intruded
into the low-grade metasediments and metavolcanics. Underground mining activity has revealed the presence of
zonation within the Hnilec granite body, and provided an
opportunity to study its differentiation and alteration
processes, including the greisens, that had occurred in the
apical parts of the granite.
The S-type character of Hnilec granites is demonstrated by several features including the high initial 87Sr/86Sr
ratio of 0.7119 (Cambel et al. 1990). The REE normalized
pattern of the granites has a significant negative Eu anomaly (Eu/Eu* ≈ 0.02), and the Rb content ranges from 170
to 870 ppm depending on the differentiation level (Broska
and Uher 2001). A high content of Fe and Al in biotite is
typical along with increasing contents of Sn in biotite and
low contents of MgO (3–4 wt%) and TiO2 (2.5–3 wt%).
On the basis a study of garnets, Faryad and Dianiška
(1989) suggested the Hnilec granites originated at a depth
>21 km and pressure ~850 MPa and that magma solidification occurred at a depth between 5–7 km. Zircon dating
indicates a Permian/Triassic age for the Hnilec granites
(243 ±18 Ma; Poller et al. 2002). A Permian age was also
obtained from electron microprobe dates of monazite (276
±13 Ma; Finger and Broska 1999), as well as an older
Rb/Sr data age (282 ±2 Ma; Cambel et al. 1990), although
the contact aureole of Súľová granite (Hnilec area) has a
140 Ma date obtained by 40Ar/39Ar dating (Vozárová et al.
2000).
The Hnilec granites are characterized by their tourmaline mineralization. Tourmaline contents generally increase towards the top of the granite body. One borehole in
the Hnilec granite body showed that whereas only 8 g/t of
tourmaline is present at a depth of 900 m, 700 g/t occurs
at 600 m, and 4200 g/t at 300 m (Rub et al. 1977). In contrast, zircon, monazite, anatase and apatite decrease in
abundance with depth. The distribution of tourmaline is
probably associated with its polystage development: the
older magmatic tourmaline is relatively unstable, especially in the two-mica granites, and is substituted by younger,
X-site deficient tourmaline, of schorl-foitite composition
(Broska et al. 1998, 1999).
The western Hnilec granite occurrence called “Súľová”
(Fig. 1) is a relatively large body approximately 2 x 1 km
in surface area, and with a strong vertical zonal structure.
The lower part is comprised of two-mica granite, the middle section consists of a medium-grained Ms granite, and
the highest part is fine-grained greisenized granite. The
greisen occurs in the apical endocontact parts and form
100–200 m long bodies, 30 m in height (locally even
more) in vertical cross section (Fig. 2). Greisenization produced a zonal arrangement of various types of metasomatites: 1) Quartz zone (0.2–1.0 m), 2) a zone of
147
Igor Broska – Michal Kubiš – C. Terry Williams – Patrik Konečný
N
Hnilec B
A
C
G
F
D
Košice
H
E
Betliar
1
2
3
4
5
0
Rožňava
5
6
7
8
10 15 km
9
Fig. 1. Geological sketch map of the Spiš-Gemer granite occurrences.
Explanation: 1 – Neogene to Quaternary sedimentary rocks, 2 – Sediments of the Inner Carpathian Paleogene: sandstones, claystones, 3 – Silicicum: limestones, dolomites (Lower Triassic-Upper Jurassic), 4 –
Meliaticum: limestones, dark shales, radiolarites, glaucophanites (Upper Permian-Upper Jurassic), 5 –
Turnaicum: limestones, dark shales (Upper Permian-Middle Jurassic), 6 – Veporicum: paragneisses, granites, migmatites (Paleozoic), 7 – Gemericum: phyllites, metasandstones, metavolcanics (Paleozoicum), 8
– Gemeric granites, 9 – Thrust faults, A - Súľová body, B - Delava body, C - Surovec body, D - Podsúľová
body – hidden granite, E - Betliar body, F - Hummel body, G - Zlatá Idka body, H - Poproč body. The
Súľová, Delava and Surovec bodies form the Hnilec granite.
quartz-mica greisen, 3) a zone of albitized fine-grained
granite together with lenses of ore greisens, 4) a zone of
microclinized medium-grained granite (Grecula 1985).
Tin is the main ore element in the gradual facial zones and
it increases in concentration towards the contact. The disseminated Sn-W-(Li-Nb-Ta) mineralization is concentrated mainly in the greisenized cupola, quartz-cassiterite
veins, as well as in the albite-Li-mica-topaz granites
(Drnzík 1982).
The main scientific goal of this study is to provide a
A`
N
B
Methodology
The samples were investigated from both thin sections and
liberated accessory mineral concentrates. The heavy mineral
fractions were obtained using
standard mineral separation procedures, i.e. a combination of
rock crushing, sieving, preliminary concentration using a Wilfley table, and final heavy
liquid and magnetic separation. The zircon morphology was
studied using a binocular microscope. The cathodoluminiscence images of zircons were obtained on a Cameca SX100
(ŠGÚDŠ Bratislava) electron microprobe, as were analyses
of feldspar, mica and zircon employing a variety of natural
minerals and synthetic compounds as standards. Analytical
parameters were an accelerating voltage of 15 kV, beam
current of 20 nA, and beam diameter of 5 µm. For analyses
of feldspar, a peak count time of 20 sec on each element was
A
840
new perspective to aspects of the
development of Hnilec granites,
based mainly on the compositions of the feldspar and REE
accessory minerals, and zircon
morphology of the principal
rock types of this intrusion, and
compared with mineralogical
and geological reviews from this
region. The samples were collected from the mine dumps, and
represent the typical rock types
present, including greisens.
820
C
B`
A
Hnilec
700
C`
740
600
A`
760
B
720
B`
780
800
C`
C
0
100 m
Fine-grained granite
Greisens and quartz veins
Fine-grained greisenized granite
Middle-grained Ms granite
Chlorite-sericite phyllites and quartzite
Overfaults
Prospecting gallery
Fig. 2. Cross-sections of the Hnilec granite body according to mining work (Drnzík et al. 1982). Sample GK-8 was taken from the medium-grained, and
GK-9 from fine-grained granite. Sample GK-10 represents the greisenized part of fine-grained granite. The granites from the map are undistinguished.
148
The compositions of rock-forming and accessory minerals from the Gemeric granites (Hnilec area, Gemeric Superunit, Western Carpathians)
2
Kfs
1
Ab
3
Ap
Ap
4
500 µm
500 µm
A
5
Kfs
B
8
Kfs
Ap
Ab
7
6
Kfs
Ab
Ms
500 µm
C
500 µm
D
Fig. 3. BSE images of feldspar from the coarse-grained and fine-grained granites. Numbers indicate the analysis positions which are presented in Table
1. A, B – GK-8; C, D – GK-9. The numerous small bright spots in the albite (Ab) grains (Fig 1A, D) are from tiny secondary apatite crystals.
employed. For the determination of low levels of P in
feldspar, a 60 sec count time was used and measured on the
P Kα line utilising a high reflectivity PET crystal. Apatite
was analysed using a Cameca SX50 electron microprobe at
the Natural History Museum, London. Operating conditions
were 15 kV, 25 nA beam current and, depending on the apatite grain size, a 1–5 µm diameter beam.
Results
Petrological and mineralogical description
The samples studied of the principal granitic rocks
from the Hnilec profile showed variable degrees of visible
mineralization. The Hnilec two-mica granite from the
lowermost part is a medium-grained rock consisting of hypidiomorphic, sericitized plagioclase (31 vol%), perthitic
K-feldspar (24 vol%), quartz (37 vol%), muscovite (5
vol%), and minor biotite (1 vol%). The main accessory
minerals are tourmaline, zircon, apatite, monazite, xeno-
time, rutile and fluorite. The upper emplaced mediumgrained Ms granite consists of plagioclase (33 vol%), Kfeldspar (25 vol%), quartz (34 vol%) and muscovite (8
vol%). Accessory minerals are represented mainly by tourmaline, zircon, apatite, monazite, xenotime, fluorite and
rare cassiterite and Nb,Ta-minerals. The fine-grained granite has the lowest amount of feldspar, and highest quartz,
comprising plagioclase (26.7 vol%), K-feldspar (12.5
vol%), quartz (44.7 vol%) and muscovite (16.6 vol%).
Zircon, uraninite, monazite, garnet and Nb,Ta-minerals
are the principal accessory minerals. The greisen consists
of quartz (60 vol%) and muscovite (38 vol%), and locally,
relicts of plagioclase could be identified. In the greisen,
clusters of tourmaline, relatively abundant cassiterite and
increased amounts of topaz, fluorite, arsenopyrite, molybdenite and pyrite are typical.
Compositions of some of the rock forming minerals
The feldspar compositions show some specific features
relating to the P contents both in the albite and K-feldspar
149
Igor Broska – Michal Kubiš – C. Terry Williams – Patrik Konečný
Table 1. Representative microprobe analyses (in wt% oxide) of feldspar.
Pos
Min
Sample
SiO2
Al2O3
FeO
CaO
Na2O
K2O
P2O5
TOTAL
Si4+
Al3+
Fe2+
Ca2+
Na+
K+
P5+
XAb
XAn
XOr
1
2
3
4
5
6
7
8
Ab
Kfs
Kfs
Kfs
Ab
Kfs
Ab
Kfs
GK-8/5-2 GK-8/5-6 GK-8/4-3 GK-8/4-4 GK-9/1-3 GK-9/4-1 GK-9/4-2 GK-9/4-3
69.36
64.90
61.21
65.01
69.44
65.29
65.18
65.62
19.41
18.27
21.53
18.06
19.41
18.23
18.33
18.27
0.00
0.01
0.01
0.05
0.01
0.00
0.00
0.00
0.02
0.00
0.09
0.17
0.06
0.00
0.00
0.00
11.74
0.28
0.29
0.19
11.53
0.24
0.28
0.22
0.10
16.30
15.40
16.42
0.08
16.58
16.69
16.73
0.02
0.04
0.00
0.00
0.05
0.11
0.00
0.00
100.64
99.80
98.53
99.90
100.57
100.44
100.48
100.84
calculated on the basis 8 O
3.007
3.004
2.869
3.009
3.009
3.005
3.002
3.010
0.992
0.996
1.189
0.985
0.992
0.989
0.995
0.987
0.000
0.000
0.000
0.002
0.000
0.000
0.000
0.000
0.001
0.000
0.005
0.008
0.003
0.000
0.000
0.000
0.987
0.025
0.027
0.017
0.969
0.021
0.025
0.020
0.005
0.963
0.921
0.970
0.004
0.973
0.981
0.979
0.001
0.002
0.000
0.000
0.002
0.004
0.000
0.000
0.994
0.026
0.028
0.017
0.993
0.021
0.025
0.020
0.001
0.000
0.005
0.008
0.003
0.000
0.000
0.000
0.005
0.974
0.967
0.974
0.004
0.979
0.975
0.980
Table 2. Representative microprobe analyses (in wt% oxide) of white mica.
No
Sample
SiO2
TiO2
Al2O3
Cr2O3
FeO
MnO
MgO
CaO
Na2O
K2O
F
Cl
total
O=F, Cl
TOTAL
Si
Ti
AlIV
AlVI
Cr
Fe
Mn
Mg
Ca
Na
K
F
Cl
1
GK-8
46.60
0.35
32.78
0.00
3.89
0.11
0.34
0.00
0.46
9.62
0.00
0.01
94.16
0.00
94.16
6.325
0.036
1.675
3.568
0.000
0.442
0.010
0.069
0.000
0.121
1.666
0.000
0.002
2
3
4
GK-8
GK-9
GK-9
47.16
48.37
47.09
0.14
0.14
0.08
31.70
28.12
30.12
0.01
0.05
0.15
3.04
6.53
5.78
0.07
0.23
0.21
1.71
0.24
0.06
0.04
0.00
0.00
0.16
0.07
0.20
10.29
10.92
10.80
0.44
1.87
2.26
0.00
0.01
0.01
94.76
96.55
96.76
0.19
0.79
0.95
94.58
95.76
95.81
calculated on the basis 22 O
6.386
6.664
6.487
0.014
0.015
0.008
1.614
1.336
1.513
3.445
3.231
3.377
0.000
0.000
0.000
0.344
0.752
0.666
0.007
0.022
0.020
0.345
0.049
0.012
0.006
0.000
0.000
0.042
0.019
0.053
1.778
1.919
1.898
0.185
0.785
0.949
0.000
0.002
0.002
within the principal granitic rocks of the Hnilec-Medvedí
potok locality. The P concentrations vary in the alkali
feldspars, although some general features are apparent. In
all cases, albite contains very low (or below detection) concentrations of P (e.g. Table 1 anal. No. 1, Fig. 3A). The Kfeldspars in the deepest coarse-grained leucogranite (see
profile) usually contain very low concentrations of P2O5
150
5
GK-10
48.23
0.16
30.62
0.00
5.64
0.02
0.41
0.00
0.13
10.03
0.70
0.00
96.79
0.29
96.50
6.501
0.016
1.499
3.365
0.000
0.636
0.002
0.083
0.000
0.021
1.893
0.181
0.000
6
GK-10
46.98
0.00
34.19
0.04
3.86
0.06
0.02
0.00
0.21
10.76
0.53
0.00
96.56
0.23
96.34
6.267
0.000
1.733
3.654
0.000
0.432
0.006
0.004
0.000
0.054
1.835
0.223
0.002
(0.0x wt%, Table 1 anal. No. 2,
and Fig. 3A, B. However, Kfeldspars from the apical finegrained granites, i.e. those rocks
that are more evolved, are more
enriched in P with P2O5 contents
up to 0.10 wt% (Table 1 anal. No.
6, Fig. 3C). It should be noted
however, that some individual Kfeldspar crystals contain exsolved
apatite grains, similar to Kfeldspar from the adjacent coarsegrained granites, are also free of P
(Table 1 anal. No. 7, 8, Fig. 3D).
Albite from this coarse-grained
granite has very low contents of P
(<0.0x wt%), although albite, exsolved from the K-feldspar as
perthitic phase, can occasionally
reach P2O5 contents of 0.05 wt%
(Table 1 anal. No. 5).
X-ray diffraction and microprobe data indicate that the white
mica present in the coarsegrained, fine-grained and greisenized granites is muscovite, and
it is present as both a primary and
secondary generation mineral.
The secondary muscovite develops in feldspar crystals, whereas
the primary muscovite is distributed mainly in interstitial localities. The muscovite contains
relatively high concentrations of
Fe, but is heterogeneously distributed. Within one grain, the FeOtot
content varies from 2 wt% to 9
wt% and so this “phengitic component” is not simply a reflection
of the Alpine metamorphic overprint of these granites. The F content in muscovite, including
muscovite from the greisenized
granite, is generally very low, although some F enrichment was
detected in the fine-grained granites. The white mica corresponds
to ferro-aluminoceladonite (Rieder et al. 1998) (Table 2).
Accessory apatite
The alkali feldspars in the Hnilec granitic rocks are
characterized by the presence of newly-formed apatite
crystals that are located mainly in albite. These apatite
crystals are usually very small, typically less than 3 microns in size (Fig. 3A, D). They are considered to be late
The compositions of rock-forming and accessory minerals from the Gemeric granites (Hnilec area, Gemeric Superunit, Western Carpathians)
Table 3. Representative microprobe analyses (in wt% oxide) of apatites.
Primary apatite are those from grains >50 µm in size; secondary apatite
are those ~5 µm in size, and distributed within albite crystals.
0.35
0.30
0.35
0.00
0.00
0.01
0.02
0.03
0.04
Fe (p. f. u.)
0.05
0.06
0.07
Fig. 4. Fe vs Mn binary plot of apatite from the Hnilec granites (samples
GK-8, GK-9, GZ-1, GZ-3).
Sample/chondrite
106
monazite
xenotime
105
104
103
apatite
102
granite
101
La Ce
Pr
Nd Sm Eu Gd Tb Dy Ho
Er Tm Yb Lu
Fig. 5. Chondrite-normalized REE patterns of minerals from the coarsegrained Hnilec granite. The whole rock REE pattern, plotted for comparison, is the average of 3 coarse-grained Hnilec granite samples. The
apatite pattern is the average of all primary apatite compositions from the
studied area.
Note the large range in REE abundances among these minerals, particularly the high concentration in monazite (LREE) and xenotime (HREE),
and the low but similar concentration and pattern of apatite compared
with the bulk rock.
magmatic or secondary phases. In contrast, exsolved apatites distributed within K-feldspar crystals, form much
larger grains (generally >250 microns in size), and often
with xenomorphic (irregular) shapes, but can also occur as
idiomorphic forms (Fig. 3D). The newly-formed apatite
had the effect of decreasing the P contents in both orthoclase and plagioclase feldspar, but the presence of minor
Ca in the Na-feldspar provides more suitable conditions
for apatite formation and consequently it is more common
in albite (Fig. 3A, Table 1). The crystallisation of secondary apatite was enhanced by the high F activity in the latestage corrosive fluid. The formation of pure end-member
albite and muscovite accompanied this late-stage process.
Apatite from the Hnilec granites forms two genetic
types, illustrated by the Mn vs. Fe binary plot (Fig. 4). The
first type is represented by the large exsolved apatite crystals distributed within the K-feldspar (e.g. Fig. 3B) and
displays a positive correlation between Mn and Fe. Additionally, this apatite type has significantly higher Mn and
Fe contents than the second genetic type - the small, newly-formed crystals in albite (e.g. Fig. 3A). This secondary
apatite, distributed in Na-feldspar, is closest to stoichio-
grain
SO3
P2O5
SiO2
ThO2
UO2
La2O3
Ce2O3
Pr2O3
Nd2O3
Sm2O3
Gd2O3
Dy2O3
Y2O3
Al2O3
CaO
SrO
FeO
MnO
MgO
Na2O
F
Cl
OH
total
O=F, Cl
TOTAL
S
P
Si
Th
U
La
Ce
Pr
Nd
Sm
Gd
Dy
Y
Al
Ca
Sr
Fe
Mn
Mg
Na
XApFAp
XApClAp
XApHAp
secondary
0.05
secondary
type
0.10
primary (?)
0.15
primary (?)
sample GK-8
0.20
primary (?)
Mn (p. f. u.)
0.25
1
0.00
41.61
0.00
0.03
0.01
0.00
0.04
0.05
0.07
0.02
0.01
0.00
0.07
0.00
52.42
0.02
0.40
3.04
0.00
0.07
3.65
0.08
0.02
101.61
1.55
100.06
0.000
2.990
0.000
0.001
0.000
0.000
0.001
0.002
0.002
0.001
0.000
0.000
0.003
0.000
4.768
0.001
0.029
0.219
0.000
0.011
2
0.03
41.61
0.04
0.03
0.10
0.01
0.03
0.08
0.05
0.03
0.02
0.01
0.03
0.03
51.03
0.03
0.91
4.16
0.00
0.01
2.92
0.04
0.32
101.51
1.24
100.27
0.002
3.011
0.004
0.001
0.002
0.000
0.001
0.002
0.002
0.001
0.001
0.000
0.002
0.003
4.672
0.002
0.065
0.301
0.000
0.001
3
0.08
42.34
0.05
0.06
0.14
0.03
0.02
0.00
0.00
0.02
0.00
0.00
0.07
0.05
53.13
0.13
0.19
2.54
0.01
0.01
2.38
0.00
0.34
101.58
1.00
100.58
0.006
3.040
0.004
0.001
0.003
0.001
0.000
0.000
0.000
0.001
0.000
0.000
0.003
0.004
4.828
0.006
0.014
0.182
0.001
0.002
4
0.00
41.91
0.34
0.01
0.09
0.00
0.05
0.21
0.09
0.01
0.06
0.00
0.00
0.09
54.16
0.00
0.06
0.32
0.02
0.06
3.06
0.04
0.28
101.40
1.30
99.54
0.000
3.003
0.029
0.000
0.002
0.000
0.001
0.006
0.003
0.000
0.002
0.000
0.000
0.009
4.912
0.000
0.004
0.023
0.002
0.009
5
0.00
42.10
0.07
0.06
0.00
0.00
0.19
0.03
0.05
0.06
0.00
0.04
0.07
0.03
55.27
0.01
0.00
0.03
0.00
0.04
3.20
0.00
0.34
101.61
1.35
100.26
0.000
3.011
0.006
0.001
0.000
0.000
0.006
0.001
0.002
0.002
0.000
0.001
0.003
0.002
5.001
0.001
0.000
0.002
0.000
0.007
0.978
0.011
0.010
0.788
0.006
0.206
0.637
0.000
0.363
0.820
0.005
0.175
0.854
0.000
0.145
metric apatite as it contains the lowest contents of Mn and
Fe, and is also depleted in REE. In contrast, the Mn-rich
apatite (Fig. 4, Table 3) is similar to apatite from “classical” S-type granites (Sha and Chappell 2000; Broska et al.
in prep.).
151
Igor Broska – Michal Kubiš – C. Terry Williams – Patrik Konečný
Table 4. Representative microprobe analyses (in wt% oxide) of monazite and xenotime from the coarsegrained granite
Point
Min
P2O5
SiO2
ThO2
UO2
La2O3
Ce2O3
Pr2O3
Nd2O3
Sm2O3
Gd2O3
Dy2O3
Er2O3
Yb2O3
Y2O3
Al2O3
CaO
MnO
SrO
PbO
Total
14.1
Mnz
29.17
0.55
4.57
0.09
11.03
30.08
3.99
13.28
3.12
1.97
0.48
n.d.
0.07
0.49
0.09
0.52
0.14
0.00
0.01
99.65
15.1
Mnz
28.04
0.94
10.87
0.15
7.476
24.02
3.22
13.17
4.44
2.88
0.80
n.d.
0.00
1.28
0.04
1.68
0.01
0.04
0.14
99.18
P
Si
Th
U
La
Ce
Pr
Nd
Sm
Gd
Dy
Er
Yb
Y
Al
Ca
Mn
Sr
Pb
0.978
0.022
0.041
0.001
0.161
0.436
0.058
0.188
0.043
0.026
0.006
n.d.
0.001
0.010
0.010
0.022
0.004
0.000
0.000
0.954
0.038
0.099
0.001
0.111
0.353
0.047
0.189
0.061
0.038
0.010
n.d.
0.000
0.027
0.004
0.072
0.000
0.001
0.002
Mo
Br
Hu
93.63
4.21
2.16
83.00
13.28
3.72
Sample GK-8
15.2
15.21
17.1
Mnz
Mnz
Xen
28.01
28.57
32.66
1.27
0.77
0.74
11.26
9.54
0.56
0.25
0.00
2.62
7.88
7.46
0.00
25.55
25.24
0.15
3.32
3.38
0.06
11.25
13.05
0.70
4.29
4.72
0.94
2.51
2.68
3.00
0.63
0.33
6.85
n.d.
n.d.
2.83
0.00
0.00
1.84
0.83
0.58
42.19
0.02
0.03
0.00
1.64
1.64
0.13
0.00
0.12
0.08
0.00
0.19
0.00
0.21
0.00
0.37
98.92
98.31
95.73
calculated on the basis 4 O
0.953
0.972
0.970
0.051
0.031
0.026
0.103
0.087
0.004
0.002
0.000
0.020
0.117
0.111
0.000
0.376
0.371
0.002
0.049
0.049
0.001
0.161
0.187
0.009
0.059
0.065
0.011
0.033
0.036
0.035
0.008
0.004
0.077
n.d.
n.d.
0.033
0.000
0.000
0.020
0.018
0.012
0.787
0.002
0.003
0.000
0.071
0.071
0.005
0.000
0.003
0.002
0.000
0.005
0.000
0.002
0.000
0.004
82.45
12.45
5.11
84.30
12.63
3.08
–
–
–
The REE distribution of the primary apatite broadly reflects the bulk REE distribution of the host granitic rock
(Fig. 5), as apatite is the principal phase hosting the REE.
There are minor contributions from monazite (LREE) and
from xenotime (HREE) (Table 4, Fig. 5), particularly in
those granites with very low La/Yb ratios (≈ 4). Zircon also contributes to the HREE budget of the whole rock.
Accessory zircon
The morphologies of zircon crystals from all the investigated granites in the Hnilec vertical profile have very
similar typological features. There are dominated by low S
and L zircon types with the mean point represented by the
S8 typological subtype (Fig. 6). The same morphological
type is present in both the coarse- and fine-grained gran-
152
ites, as well as in the greisen or
strongly greisenized granite
types. This similarity of zircon
18.1
13.1
13.2
morphology from the different
Xen
Xen
Xen
Spiš-Gemer granitic rocks sug32.84
31.85
31.21
gests that this largest population
0.89
0.66
0.84
of zircons originated from a
0.65
1.32
1.07
1.23
2.29
3.49
common magma source, proba0.04
0.00
0.00
bly as an early crystallized phase
0.17
0.28
0.00
transported by residual melts to
0.00
0.02
0.06
their final emplacement in the
0.21
0.61
0.66
1.08
1.16
0.91
granitic rocks. It is likely that
3.17
3.32
3.07
only zircons with G1 and some L
6.83
6.58
6.28
subtypes
were formed in the al2.97
n.d.
n.d.
ready
emplaced
granites (Fig. 7).
2.27
2.17
2.12
The CL images of zircon
41.50
41.38
40.45
0.00
0.00
0.00
show a distinctive oscillatory
0.14
0.20
0.17
zonation for all of the investigat0.00
0.00
0.01
ed crystals. Oscillatory zoning is
0.00
0.00
0.00
a common feature in zircons
0.22
0.45
0.41
94.21
92.30
90.76
from acid igneous rocks and is
believed to have formed during
0.979
0.975
0.972
long periods of crystallization.
0.031
0.024
0.031
This feature supports the hypoth0.005
0.011
0.009
0.010
0.018
0.029
esis that the majority of zircons
0.001
0.000
0.000
from the Hnilec granites crystal0.002
0.004
0.000
lized in the deep-seated magma
0.000
0.000
0.001
chambers (Fig. 8A, B), and is fur0.003
0.008
0.009
0.013
0.014
0.012
ther supported by the presence of
0.037
0.040
0.037
two zircon nuclei centres present
0.077
0.077
0.074
in one grain (Fig. 8A).
0.035
n.d.
n.d.
The zircon nuclei display
0.024
0.024
0.024
0.777
0.796
0.792
magmatic oscillatory zonation,
0.000
0.000
0.000
and have rounded forms indicat0.005
0.008
0.007
ing a complex crystallization/re0.000
0.000
0.000
sorption process had taken place
0.000
0.000
0.000
0.002
0.004
0.004
over a long period of time. After
–
–
–
some significant change in the
–
–
–
magma composition, the relative
–
–
–
growth rates of different zircon
crystal faces changed, preferentially favouring the growth of pyramidal (311) planes (Vavra
1990). It seems likely that such a complex crystal growth
history could only be achieved from the zircon residing for
long periods of time in different magma batches. The oscillatory zoning of zircons is locally overprinted, probably by
changes in fluid composition and perhaps by loss of U, Th,
Pb, that accompanied this process (Pidgeon 1992).
Microprobe analyses of zircon showed that P was below the detection limit, and that the Zr/Hf ratios were relatively constant around a mean value of ~31 (Table 5),
which is typical for magmatic zircons from anatectic rocks
(Pupin, 2000). Some variation in Y (and U) was observed,
with Y showing some degree of correlation with Hf. As P
was not present in the zircon, no epitaxial xenotime overgrowths were observed and xenotime crystallized as a separate phase (Fig. 7).
The compositions of rock-forming and accessory minerals from the Gemeric granites (Hnilec area, Gemeric Superunit, Western Carpathians)
I.A
100
200
300
400
500
600
700
800
100
13-1
Variscan
S-type
200
Zrn
300
I.T
400
Variscan
I-type
500
600
Variscan
post-org.
SS-type
Early
Alpine
Xen
Variscan
post-orog.
A-type
Late Alpine
700
13-2
800
Fig. 6. The projection of the typological mean points of zircons from
granites in the Spiš-Gemer area. Zircon typological mean points from the
Hnilec region are represented by squares; those from Betliar, Poproč and
Hummell granite bodies are represented by triangles. Full circle show the
position of the mean point of all samples. For comparison, fields for the
Variscan I and S zircon mean points, and from the post-orogenic A-type
granites are shown. Typogram and calculation of the mean points after
Pupin (1980). The mean points of the investigated samples have the following parameters: GK-8: IA = 475, IT = 287 (where - IA = agpacity index, IT = index of temperature); GK-9: IA = 445, IT = 360; GK-10: IA
= 423, IT = 330.
Fig. 7. BSE image of xenotime-zircon paragenesis. The zircon crystal is
here cut perpendicular to the the c-axis, and the typologically belongs to
G1 or L subtype, which represents the most recent zircon crystallisation
phase. The zonation observed in the xenotime is mainly due to the variations in concentrations of U, with the lighter areas of xenotime being
enriched in U.
Table 5. Representative microprobe analyses (in wt% oxide) of the zircon. Zr/Hfw calculated from atomic weights.
Discussion and concluding
remarks
No.
Sample
grain
pos
SiO2
ZrO2
HfO2
ThO2
UO2
Al2O3
Y2O3
Yb2O3
CaO
FeO
Total
11
GK-8
1
core
32.52
65.47
1.72
0.01
0.02
0.00
0.00
0.09
0.00
0.00
99.81
12
GK-8
1
rim
32.52
65.47
1.71
0.02
0.08
0.00
0.08
0.00
0.01
0.07
99.09
14
15
16
GK-8
GK-8
GK-8
2
2
3
core
rim
core
33.38
33.81
32.28
63.71
63.41
64.33
1.83
1.83
1.76
0.01
0.00
0.03
0.05
0.00
0.00
0.00
0.10
0.00
0.03
0.07
0.34
0.04
0.08
0.13
0.00
0.04
0.00
0.02
0.00
0.00
99.05
99.34
98.86
calculated on the basis 4 O
1.027
1.034
1.004
0.956
0.946
0.975
0.016
0.016
0.016
0.000
0.000
0.000
0.000
0.000
0.000
0.000
0.003
0.000
0.001
0.001
0.006
0.000
0.001
0.001
0.000
0.001
0.000
0.001
0.000
0.000
30.4
30.3
31.9
17
GK-8
3
rim
32.47
63.38
1.67
0.00
0.14
0.00
0.52
0.19
0.00
0.10
98.46
18
GK-9
1
core
33.02
63.48
1.92
0.05
0.02
0.00
0.19
0.11
0.00
0.04
98.83
19
GK-9
1
rim
30.05
64.31
1.26
0.10
0.05
0.00
0.44
0.04
0.00
0.00
96.24
Compositional features of
one of the rock-forming minerals (feldspar), coupled with morphological characteristics of an
accessory phase (zircon) combine to provide strong evidence
for a cogenetic relationship between the main granitic phases
at the Hnilec-Medvedí potok locality. The presence of the P in Si4+
1.001
1.022
1.012
1.021
0.970
alkali feldspars in the Hnilec Zr4+
0.983
0.960
0.963
0.958
1.012
0.015
0.016
0.015
0.017
0.012
granites is the direct result of the Hf4+
0.000
0.000
0.000
0.000
0.001
high peraluminosity of the melt Th4+
4+
0.000
0.001
0.001
0.000
0.000
(ACNK >1.2) when P behaves U 3+
Al
0.000
0.000
0.000
0.000
0.000
as an incompatible element.
Y3+
0.000
0.001
0.009
0.003
0.008
Phosphorus precipitates in late Yb3+
0.001
0.000
0.002
0.001
0.000
magmatic apatite as well as be- Ca2+
0.000
0.001
0.000
0.000
0.000
ing incorporated as a minor Fe2+
0.000
0.002
0.003
0.001
0.000
component in alkali feldspar. Zr/Hfw
33.4
30.5
33.0
28.8
44.5
The concentration of P in alkali
feldspar increases from the
medium- to fine-grained granites and corresponds to an in- minous granites, including tin-bearing types (Frýda and
crease in Al activity in the melt. High P contents in alkali Breiter, 1995; Breiter et al., 1997; Breiter et al., 1999,
feldspars have been reported from all the evolved peralu- 2002; Breiter, 2001). This incorporation of P in feldspar,
153
Igor Broska – Michal Kubiš – C. Terry Williams – Patrik Konečný
19
14
18
15
A
B
Fig. 8. CL images of zircon crystals. Note the oscillatory zonation within the nucleus in image A. A – GK-8, B – GK-9.
known as the berlinite substitution, results from berlinite
(ideally AlPO4) being isostructural with the Si2O4 framework component of feldspar (London et al., 1990; London, 1992; London, 1998). The influence of P on silicate
melt structure is well known, and it has the effect of depolymerising melts that results in a lowering of the melt
viscosity (Mysen et al., 1981). The M-PO4 complexes occur in the melt as a result of metal cation transfer from
non-bridging oxygens (NBO) in the silicate portion of the
melt structure, whereas the ratio of non-bridging oxygens
to tetrahedral cation (NBO/T) is decreased. Coupled with
an enrichment of volatile elements (such as H2O, F, B) in
the fluid phase, this decrease in the melt viscosity has the
effect of mobilising and concentrating elements such as
Sn, Nb and Ta in the cupolas of granite massifs, which locally show evidence for fluid-element saturation, indicated by minerals such as cassiterite, tourmaline and fluorite,
together with albitization and greisenization. This process
also operated in the cupola of the Hnilec granite body and
resulted in an evolved tin mineralization stage. The cassiterite-bearing greisen forms part of the endocontact of the
fine-grained granite intrusion as indicated by the composition and morphology of zircon being very similar to zircon
from the main granitic phases.
The similarities of the granites from Hnilec-Medvedí
potok locality and other granite bodies in the Hnilec area
indicate a close genetic relationship to granite bodies in all
of the Spiš-Gemer area. The zircon morphology in granites is very similar throughout the Spiš-Gemer region (Jakabská and Rozložník 1989, Broska and Uher 1991), and
is undoubtedly a reflection of similar sources for, and similar evolution of, these granites. The zircon morphological
and compositional characteristics have many features in
common with zircons from aluminium-rich crustal granitoids (Jakabská and Rozložník 1989).
Evidence from zircon CL images and morphology suggest that the granites could be derived from batches of
deep-seated granitic melt on the crust-mantle boundary
154
formed by a heat source of the rift related mafic melt that
occurred during the Permian-Triassic period. The high B
concentration, so typical of the Spiš-Gemer granites, can
be explained by an enrichment of their primary melts by B
introduced from volcanic emanations that operated in the
deep fault systems or rifts during the Permian period
(Broska and Uher 2001), and particularly also from B-enriched mica sources (Petrík and Kohút 1997). The SpišGemer granites in this sense may have resulted from
melting triggered by increasing heat flow from the mantle
below the continental crust, during the post-collisional extension of the Variscan orogeny. The rifting process that
opened the Meliata-Hallstadt Ocean in the Triassic could
have been a thermal source for melting of the Gemeric
granites (Broska and Uher 2001) and what is evoked also
by former assumed contact of the Gemeric terrain with the
Meliata Ocean (Plašienka et al. 1999).
A c k n o w l e d g e m e n t s . The project was supported by grant
VEGA #1143 (Slovak Acad. Sci.) and EU-HIP Programme for IB to visit the Natural History Museum, London. The authors are greatly indebted also to Francis Wall for her help with the apatite investigation by CL
method (NHM, London).
R e f e re n c e s
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155
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