GEOLOGICA CARPATHICA, APRIL 2005, 56, 2, 179189
www.geologicacarpathica.sk
Origin of amphibole megacrysts in the Pliocene-Pleistocene
basalts of the Carpathian-Pannonian region
ATTILA DEMÉNY
1
, TORSTEN W. VENNEMANN
2
, ZOLTÁN HOMONNAY
3
, ANDY MILTON
4
,
ANTAL EMBEY-ISZTIN
5
and GÉZA NAGY
1
1
Laboratory for Geochemical Research, Hungarian Academy of Sciences, Budaörsi út 45, H-1112 Budapest, Hungary;
demeny@geochem.hu
2
Institut de Minéralogie et Géochimie, Université de Lausanne, BFSH-2, CH-1015 Lausanne, Switzerland
3
Department of Nuclear Chemistry, Eötvös Loránd University, Pázmány Pétér sétány 2, H-1117 Budapest, Hungary
4
School of Ocean and Earth Science, Southampton Oceanography Centre, European Way, Empress Dock, Southampton, SO14 3ZH,
United Kingdom
5
Department of Mineralogy and Petrology, Hungarian Natural History Museum, Ludovika tér 2, H-1088 Budapest, Hungary
(Manuscript received October 28, 2003; accepted in revised form June 16, 2004)
Abstract: Major and trace element compositions, stable H and O isotope compositions and Fe
3+
contents of amphibole
megacrysts of Pliocene-Pleistocene alkaline basalts have been investigated to obtain information on the origin of mantle
fluids beneath the Carpathian-Pannonian region. The megacrysts have been regarded as igneous cumulates formed in the
mantle and brought to the surface by the basaltic magma. The studied amphiboles have oxygen isotope compositions
(5.4±0.2 , 1 σ), supporting their primary mantle origin. Even within the small δ
18
O variation observed, correlations
with major and trace elements are detected. The negative δ
18
O-MgO and the positive δ
18
O-La/Sm(N) correlations are
interpreted to have resulted from varying degrees of partial melting. The halogen (F, Cl) contents are very low (<0.1 wt. %),
however, a firm negative (F+Cl)-MgO correlation (R
2
= 0.84) can be related to the Mg-Cl avoidance in the amphibole
structure. The relationships between water contents, H isotope compositions and Fe
3+
contents of the amphibole megacrysts
revealed degassing. Selected undegassed amphibole megacrysts show a wide δD range from 80 to 20 . The low δD
value is characteristic of the normal mantle, whereas the high δD values may indicate the influence of fluids released
from subducted oceanic crust. The chemical and isotopic evidence collectively suggest that formation of the amphibole
megacrysts is related to fluid metasomatism, whereas direct melt addition is insignificant.
Key words: Carpathian-Pannonian region, mantle metasomatism, amphibole megacrysts, degassing, trace elements,
stable isotopes.
Introduction
Megacrysts hosted in alkaline basalts are frequently studied in
order to obtain information on the chemical characteristics of
the mantle source. Among the usual megacryst assemblage
(olivine, pyroxene, spinel, amphibole), amphibole is potential-
ly the most important mineral due to the amount of informa-
tion it can provide. Amphiboles are usually enriched in trace
elements compared to other megacryst minerals (e.g. Witt-
Eikschen & Harte 1994; Johnson et al. 1996), making trace el-
ement analyses relatively easy and accurate. Oxygen isotope
compositions of amphiboles have long been studied with high
precision due to the reactivity with the commonly used BrF
5
reagent in contrast to other megacryst minerals like olivine or
spinel. Being a hydrous mineral, amphiboles may also be anal-
ysed for H isotope composition, thus yielding important con-
straints on metasomatic processes in the mantle. These fea-
tures may make amphibole an excellent subject of mantle
geochemistry studies. However, there are a number of factors
that must be taken into account in interpreting the chemical
and isotopic compositions. First, megacryst amphiboles are
usually formed by metasomatic processes in the mantle, thus
their co-existence with other megacryst minerals can be ques-
tionable. Second, trace element contents may be strongly frac-
tionated due to crystallochemical variations in the amphiboles
and its parent magma (H
2
O activity, fO
2
, etc., e.g. Ionov &
Hofmann 1995; Oberti et al. 2000; Tiepolo et al. 2000), which
are not very well constrained at present. Third, H isotope com-
positions are very sensitive to late-stage processes, like degas-
sing and hydrothermal alteration. These uncertainties preclude
simple and straightforward interpretations of geochemical data
on amphiboles. Nevertheless, a combined approach using sev-
eral independent methods can lead to reliable reconstruction
of mantle geochemistry.
The Pliocene-Pleistocene alkaline basalts of the Pannonian
Basin frequently contain amphibole megacrysts, which have
been thoroughly studied for crystallochemical characteristics
and trace element compositions (Zanetti et al. 1995). In recent
studies, Downes et al. (1995) and Dobosi et al. (1998, 2003a)
reported the trace element contents and O-Sr-Nd isotope data
of the amphibole megacrysts. Their major conclusion is that
the amphiboles reflect the primitive mantle source beneath the
Pannonian Basin. The alkaline basalts and their mantle xeno-
liths have been widely investigated, making the region one of
180 DEMÉNY et al.
the best known areas of the world (see for example Kurat et al.
1991; Downes et al. 1992, 1995; Dobosi et al. 1995; Downes
& Vaselli 1995; Embey-Isztin & Dobosi 1995; Harangi et al.
1995; Koneèný et al. 1995a,b; Szabó et al. 1995a,b; Török &
De Vivo 1995; Huraiová et al. 1996; Szabó et al. 1996; Kemp-
ton et al. 1997; Rosenbaum et al. 1997; Hurai et al. 1998;
Seghedi et al. 2004). The trace element and radiogenic isotope
studies showed that the mantle source was contaminated by
subducted crustal material (as exemplified by the calc-alkaline
volcanism that preceeded the alkaline basaltic one), especially
in the central areas (Downes et al. 1992; Embey-Isztin & Do-
bosi 1995; Szabó et al. 1996; Seghedi et al. 2004). These re-
sults suggested that the amphibole megacrysts may provide
useful information on the metasomatizing fluids and melts. H
and O isotope studies coupled with H
2
O wt. % and Fe
3+
% de-
terminations on mantle-derived rocks and their minerals from
the Carpathian-Pannonian region are very rare (Hurai et al.
1998; Demény et al. 2001, 2004). In this study we report ma-
jor and trace element data, stable isotope compositions and re-
dox characteristics from amphibole megacrysts and a horn-
blendite vein (from Szigliget, Balaton Highland, Hungary, see
also Embey-Isztin 1976) and compare the compositions to
those obtained on xenolith-hosted amphiboles from Southern
Slovakia (Hurai et al. 1998). With the combined use of these
methods we discuss the processes coincidental with amphib-
ole crystallization.
Geological background and samples
The Neogene volcanism developed after and probably in re-
sponse to a series of tectonic events, which formed the Car-
pathian Basin. The main factors controlling the Neogene evo-
lution of the Carpathian Basin and related post-extensional
volcanism are continental collision in the Central and Eastern
Alps, which induced extensional collapse of the orogenic ter-
ranes, subduction of the European plate along the Carpathians
and related calc-alkaline volcanism, updoming of the astheno-
sphere and heating, as well as thinning of the lithosphere (e.g.
Embey-Isztin et al. 1990; Szabó et al. 1992; Horváth 1993).
Adiabatic decompression in the rising mantle caused partial
melting and the melts erupted as basaltic lava flows and tuffs
(Dobosi et al. 1995; Harangi et al. 1995; Koneèný et al.
1995a,b). The alkaline basaltic volcanism span the period of
17 to 0.5 m.yr. (see Pécskay et al. 1995 and Koneèný et al.
1995a for review). Alkali basalts are remarkably fresh (Em-
bey-Isztin & Scharbert 1981; Embey-Isztin et al. 1993a,b),
moderately porphyritic and holocrystalline. The basalts com-
prise phenocryst olivine (forsterite (Fo) content of 7886), in
some cases accompanied by clinopyroxene. The matrix is
composed of plagioclase, Ti-rich clinopyroxene, olivine, tita-
nomagnetite occasionally coexisting with ilmenite, and apatite
(Embey-Isztin et al. 1993a). Sr, Nd, and Pb isotope ratios have
values between inferred DM (depleted mantle), EM II (en-
riched mantle II), and HIMU (high µ, high
238
U/
204
Pb) (Salters
et al. 1988; Embey-Isztin et al. 1993a; Embey-Isztin & Dobosi
1995; Harangi et al. 1995). The isotope compositions of the
lavas of the central part of the Carpathian-Pannonian region
plot closer to the EM II composition than those of the margin-
al areas with a higher amount of an asthenospheric compo-
nent. Low Nb/La, Ce/Pb,
143
Nd/
144
Nd,
206
Pb/
204
Pb, and high
87
Sr/
86
Sr and
207
Pb/
204
Pb ratios have been interpreted as evi-
dence for a relationship of the EM component with a metaso-
matic enrichment process of the mantle lithosphere during
Tertiary plate subduction (Rosenbaum et al. 1997).
The amphibole megacrysts of the alkaline basalts have been
interpreted as fragments of igneous cumulates precipited from
an early magma intrusion and brought to the surface by later
basaltic magmatism (Szabó & Taylor 1994; Downes et al.
1995; Huraiová et al. 1996; Dobosi et al. 2003a). The locali-
ties studied in this paper are in Burgenland (Tobaj and Güss-
ing, Austria; samples Tob, Gü), the Balaton Highland (Bala-
tonboglár, Mindszentkálla, Szigliget, Bondoróhegy, Hungary;
samples: Bb, Bog, M, Szg, Bo), the Nógrád (Hungary)-
Gemer (Slovakia) region (Maková, west of Luèenec, Slova-
kia, and Szilváskõ, east of Salgótarján, Hungary; samples:
Mas, Szil) and the Persani Mountains (Trestia Valley, Roma-
nia; samples: Trs) (see also Fig. 1). Hornblendite veins from
composite mantle xenoliths from Szigliget (samples Szg and
Sz-3024) (see also Embey-Isztin 1976) are also studied. The
amphibole megacrysts reach 13 cm in diameter. The collec-
tion criteria were the unaltered state and the freshness of the
host rocks. Out of the collection, only those megacrysts were
analysed whose fragments were clear and inclusion-free under
binocular microscope. The purity of the amphibole separates
was checked by the standard XRD technique.
Analytical methods
Mineral major element compositions were determined with
a JEOL Superprobe 733 electron microprobe at the Laboratory
for Geochemical Research of the Hungarian Academy of Sci-
ences. The conditions used were: wavelength dispersive spec-
trometers, 15 kV accelerating voltage, 40 nA beam current,
20 µm beam diameter, 6×4 s counting time. For F and Cl
analyses beam current was modified to 100 nA. Standardiza-
tion was conducted by using mineral and artificial glass stan-
Fig. 1. Geological sketch-map of the Carpathian-Pannonian region
and sample locality areas. After Dobosi et al. (1998).
ORIGIN OF AMPHIBOLE MEGACRYSTS IN THE PLIOCENE-PLEISTOCENE BASALTS 181
dards (BaF
2
and scapolite for F and Cl, respectively). The raw
data was corrected using the ZAF correction program provid-
ed by JEOL. The relative errors of major element analyses are
lower than 2 % for oxides with >10 wt. %, about 57 % for
oxides with 15 wt. %, and about 30 % for halogens. The H
2
O
content of amphiboles was determined using the vacuum fu-
sion method adapted after Vennemann & ONeil (1993).
Hand-picked amphibole grains were powdered using an ag-
ate mortar and pestle under ethanol, which dries quickly at
room temperature, thus, oxidation during prolonged grinding
was avoided. The
57
Fe Mössbauer spectra of amphiboles were
recorded at room temperature in constant acceleration mode
with a Wissel spectrometer using
57
Co(Rh) source at the De-
partment of Nuclear Chemistry, Budapest. The Fe
3+
contents
were calculated from the relative spectral contributions of the
doublets of the corresponding Fe
2+
and Fe
3+
species, using a
Lorentzian least square fitting by Mösswinn
®
. In order to
avoid even the residual minor texture, the samples were re-
corded at the magic angle (54.7 deg), resulting in symmetrical
doublets in the spectra. This method is trivially applicable for
single crystals, but it is also good for polycrystals if the micro-
crystals have random distribution in the plane perpendicular to
the direction of the gamma rays, a condition which is very
easy to meet. Analysis of the spectra under such conditions, as
far as the Fe
2+
to Fe
3+
ratio is concerned, is very reliable
(±1 %). Further deconvolution of the spectra into several Fe
3+
and Fe
2+
doublets is rather ambiguous, but it does not interfere
with the Fe
2+
/Fe
3+
ratio if the fit in the statistical sense is cor-
rect. This is due to the fact that the higher velocity lines of the
Fe
2+
doublets do not overlap with the Fe
3+
doublet lines.
Laser ablation ICP-MS trace element analyses were carried
out on a VG Elemental PQ2+ ICP-MS coupled to a 4D Engi-
neering (Hannover, Germany) excimer laser system at the
Southampton Oceanography Centre. Measurements were
made using a 30 µm laser beam focused on the polished sam-
ple surface. Following a pre-ablation time of 10 seconds, data
were collected for 30 seconds. The analytical protocol fol-
lowed the blankstandardsamplestandard pattern. For the
purpose of calibration and monitoring instrument performance
during the analysis session, a polished piece of the NIST 610
glass standard containing the trace elements of interest was
used. After collection, the trace element data were corrected
for any instrumental drift, gas blank subtracted and then cali-
brated against the NIST 610 standard (using the preferred av-
erage values of Pearce et al. 1997). To correct for matrix ef-
fects between the NIST 610 standard and the various minerals
analysed, internal corrections were applied using appropriate
element abundances (e.g. Ca, Ti, etc.) determined by electron
microprobe analyses. On the NIST 610 standard, 10 repeat
measurements are reproducible to within 5 %. Accuracy of the
LAM-ICP-MS analyses was checked by comparison with
INAA data on amphibole samples. The REE concentrations
are accurate within 15 %. Classical ICP-MS techniques may
produce better accuracy, however, laser-ablation analysis has
the advantage of choosing the freshest (i.e. least altered) min-
eral surface, which may be difficult with bulk analysis. The
accuracy is reflected by Fig. 3, which shows good agreement
between trace element data obtained on similar samples but in
different laboratories (see also Dobosi et al. 2003a).
H and O isotope compositions, and water contents were
measured on hand-picked amphibole separates by convention-
al extraction and mass spectrometric techniques (Clayton &
Mayeda 1963; Vennemann & ONeil 1993; Demény 1995;
Demény et al. 1997) using Finnigan MAT delta S and 252
type mass spectrometers at the Laboratory for Geochemical
Research, Budapest and the University of Tübingen, respec-
tively. Some of the amphibole megacrysts have been analysed
using the laser-based method at the University of Tübingen.
The method used was adapted after Sharp (1990) using F
2
as
reagent, a Pt-disc sample holder, and measuring the
18
O/
16
O
on O
2
gas (Rumble & Hoering 1994; Kasemann et al. 2001).
The results are expressed in the δ-notation (δ = (R
1
/R
2
1)×1000
where R
1
and R
2
are the D/H or
18
O/
16
O ratios in the sample
and the standard, respectively) in per mil () relative to
V-SMOW. Reproducibilities are ±2 for δD and better than
±0.15 for
18
O values of the minerals analysed in duplicate.
NBS-28 quartz and UWG-2 garnet standards yielded δ
18
O
values of 9.64 and 5.83±0.11 (1 σ, n = 5, theoretical
value: 5.8 , Valley et al. 1995), respectively, during the
course of the laser-based analyses.
Results
The major element compositions of amphiboles determined
by electron microprobe analyses are listed in Table 1. Classifi-
cation of amphiboles was conducted according to Leake et al.
(1997) based on chemical compositions determined by means
of electron microprobe analyses. The proportion of Fe
3+
was
also calculated according to Schumacher (1997), though Fe
3+
/
Fe
tot
ratios were determined by Mössbauer spectroscopy (see
Table 2). Taking crystal chemical considerations into account,
Schumacher (1997) established 3 maximum and 3 minimum
criteria. These criteria allow the determination of an interval
rather than an exact value for the Fe
3+
/Fe
2+
ratio. The calcula-
tions yielded 0 as the minimum value for all of the studied am-
phiboles, thus, Table 1 shows only the upper limit Fe
3+
/Fe
2+
ratios along with ion numbers calculated for 23 oxygens. All
of the studied amphiboles plot in the calcic amphibole field of
Leake et al. (1997). The lines in Fig. 2 show the variations of
the ion number calculations between the lower and upper lim-
its of Fe
3+
/Fe
2+
ratios calculated for the individual amphibole
samples according to Schumacher (1997). The ion numbers
were also calculated using the Fe
3+
/Fe
2+
ratios determined by
Mössbauer spectroscopy. These latter data plot on the lines
obtained by using the method of Schumacher (1997), except
for two samples (M-2001 and M-2002, see Tables 1 and 2),
where the calculated upper limit is lower than the ratio ob-
tained by Mössbauer spectroscopy, thus as a consequence the
calculated substitution at the C crystal position is lower than
5. The large range between lower and upper Fe
3+
/Fe
2+
ratio
limits (reaching 0.97) make the classification based on calcu-
lated Fe
3+
/Fe
2+
ratio ambiguous (see also Fig. 2). Thus, a firm
classification can be made using the Mössbauer spectroscopy
data. Fig. 2 shows that the studied amphiboles are kaersutites
and magnesio-hastingsites at Maková, pargasites at Tobaj
and Güssing (Burgenland) and magnesio-hastingsites at the
other localities. As shown by Table 1, the most significant
182 DEMÉNY et al.
chemical differences are in the MgO, FeO and TiO
2
contents.
Areal variations among sampling localities have not been ob-
served. The Cl and F contents are rather low ranging between
0.02 and 0.06 wt. % and 0.01 and 0.03 wt. %, respectively.
Although the halogen contents are close to detection limits, a
firm correlation with MgO content occurs. The F and Cl anal-
yses were conducted using separate standardization proce-
dure, thus any measurement-related artifact can be excluded.
The mg# numbers tend to decrease with increasing F+Cl con-
tent (Table 1, the correlation calculation yields R
2
= 0.84). In-
Table 1: Chemical compositions (major elements in weight %, trace elements in ppm) of amphibole megacrysts from alkaline basalts of the
Carpathian-Pannonian region. Electron microprobe data are averages of two analyses selected from 45 analyses on the basis of reliablity
(e.g. oxide total). Fe
3+
/Fe
total
are maximum ratios calculated following Schumacher (1997). Trace element compositions are averages of
810 laser spot analyses.
Tob-1
Tob-2
Gü-1
M-2001
M-2002
MAS-1
MAS-2
MAS-3
Szil-1
Trs-2007
SiO
2
40.47
41.43
40.54
39.29
39.66
39.74
39.76
40.71
38.94
39.37
TiO
2
3.27
3.35
3.75
3.67
3.70
5.43
4.29
5.10
3.99
4.28
Al
2
O
3
14.54
14.45
14.10
14.48
14.67
15.57
14.81
15.68
15.35
15.42
FeO
7.84
8.80
7.92
13.93
14.22
9.38
9.72
10.51
12.17
13.24
MnO
0.05
0.08
0.10
0.15
0.15
0.09
0.10
0.12
0.15
0.17
MgO
15.70
15.0
14.76
11.17
10.93
14.32
14.06
13.33
12.35
11.45
CaO
10.31
10.15
10.85
9.88
10.28
10.05
10.14
10.28
10.51
10.05
Na
2
O
2.37
2.47
2.32
2.69
2.54
2.94
2.68
2.97
2.58
2.70
K
2
O
2.00
1.94
1.96
1.92
2.05
0.82
1.21
0.89
1.50
1.64
Cl
0.02
0.03
0.02
0.04
0.04
0.02
0.02
0.02
0.04
0.06
F
0.01
0.01
0.03
0.03
0.03
0.02
0.01
0.02
0.01
0.02
SUM
96.58
97.73
96.35
97.24
98.25
98.36
96.80
99.60
97.59
98.37
Fe
3+
/Fe
total
0.966
0.787
0.431
0.402
0.312
0.816
0.735
0.000
0.503
0.432
Si
5.841
5.930
5.942
5.834
5.856
5.650
5.768
5.851
5.704
5.747
Al(iv)
2.159
2.070
2.058
2.166
2.144
2.350
2.232
2.149
2.296
2.253
Al(vi)
0.315
0.367
0.378
0.369
0.409
0.259
0.300
0.507
0.354
0.401
Ti
0.355
0.361
0.413
0.410
0.411
0.581
0.468
0.551
0.440
0.470
Fe
3+
0.914
0.829
0.419
0.695
0.548
0.910
0.866
0.000
0.751
0.698
Mg(c)
3.378
3.209
3.225
2.473
2.406
3.035
3.040
2.856
2.697
2.492
Fe
2+
(c)
0.032
0.225
0.552
1.035
1.208
0.205
0.313
1.086
0.740
0.918
Mn(c)
0.006
0.010
0.012
0.019
0.019
0.011
0.012
0.000
0.019
0.021
Mg(b)
0.000
0.000
0.000
0.000
0.000
0.000
0.000
0.000
0.000
0.000
Fe
2+
(b)
0.000
0.000
0.000
0.000
0.000
0.000
0.000
0.177
0.000
0.000
Mn(b)
0.000
0.000
0.000
0.000
0.000
0.000
0.000
0.015
0.000
0.000
Ca(b)
1.594
1.557
1.704
1.572
1.626
1.531
1.576
1.583
1.650
1.572
Na(b)
0.406
0.443
0.296
0.428
0.374
0.469
0.424
0.225
0.350
0.428
Na(a)
0.258
0.242
0.363
0.347
0.354
0.341
0.330
0.602
0.382
0.336
K
0.368
0.354
0.367
0.364
0.386
0.149
0.224
0.163
0.280
0.305
ÓKation
15.626
15.596
15.730
15.710
15.740
15.490
15.554
15.765
15.663
15.642
(Na+K)a
0.63
0.60
0.73
0.71
0.74
0.49
0.55
0.77
0.66
0.64
Mg#
0.99
0.93
0.85
0.71
0.67
0.94
0.91
0.69
0.79
0.73
Rb
14
16
13
3
5
Sr
296
417
415
478
402
Y
8
9
10
15
14
Zr
30
52
54
31
34
Nb
13
16
16
12
16
Cs
0.05
0.04
0.09
0.04
0.10
Ba
235
302
309
160
273
La
3.0
4.6
4.7
2.2
3.2
Ce
10
15
14
8
11
Pr
1.7
2.5
2.4
1.5
2.1
Nd
10
14
15
10
13
Sm
2.9
4.1
4.0
3.6
4.2
Eu
1.0
1.3
1.5
1.5
1.7
Gd
2.7
3.6
4.0
3.1
3.2
Tb
0.35
0.46
0.48
0.60
0.71
Dy
1.7
2.2
2.7
3.6
3.4
Ho
0.30
0.35
0.33
0.55
0.57
Er
0.67
0.84
1.02
1.57
1.36
Tm
0.07
0.09
0.09
0.18
0.16
Yb
0.51
0.75
0.71
1.22
1.11
Lu
0.06
0.07
0.06
0.14
0.11
Hf
1.4
2.2
2.0
1.2
1.6
Ta
0.68
0.80
0.92
0.54
0.74
Pb
0.13
0.30
0.33
0.23
0.21
Th
0.07
0.07
0.04
0.00
0.00
U
0.03
0.06
0.02
0.05
0.12
terestingly, the halogen contents have no relationship with the
H
2
O content of the amphiboles (Table 1), thus, they are not
subject to simple OH
F+Cl substitution.
Trace element compositions are listed in Table 1. The sam-
ples show only small, but firm variations (Fig. 3). Chondrite-
normalized REE and primitive mantle-normalized trace ele-
ment diagrams (Fig. 4) show that the megacrysts studied are
very similar to those analysed by Dobosi et al. (2003a) from
the Balaton Highland Volcanic Field (western part of the Pan-
nonian Basin). Even within the small variations observed, sys-
ORIGIN OF AMPHIBOLE MEGACRYSTS IN THE PLIOCENE-PLEISTOCENE BASALTS 183
Fig. 2. Classification diagrams of Leake et al. (1997) for calcic
amphiboles. The lines represent intervals obtained by Fe
3+
/Fe
tot
calculation using the method of Schumacher (1997); diamonds are
the compositions calculated on the basis of Mössbauer spectrosco-
py data (listed in Table 2).
oxygen isotope compositions are negatively correlated with
MgO and positively with La/Sm(N) ratios (Fig. 6).
The hydrogen isotope compositions show a wide variation
from 107 to 15 (Table 2). The δD values are plotted
as a function of H
2
O contents of amphiboles (Fig. 7). The δD
and δ
18
O data on xenolith-hosted amphiboles from the
Pliocene to Pleistocene maar volcanoes and tuff cones of
Southern Slovakia (Hurai et al. 1998) plot in the normal man-
tle ranges. Although Fe
3+
contents are given only on the basis
of electron microprobe analyses for two amphibole samples
analysed for δD and δ
18
O, they can be used as a rough esti-
mate of the oxidation state. Thus, the data of Hurai et al.
(1998) will be used for comparison for the discussion of δD-
δ
18
O-H
2
O wt. %-Fe
3+
correlations (see also Figs. 7, 8, 10 and
11). A slight positive correlation appears with low δD values
Fig. 3. Trace element contents normalized to the primitive mantle
(Hofmann 1988) in amphibole megacrysts from the Carpathian-Pan-
nonian region. Shaded field: amphibole megacrysts of Dobosi et al.
(2003a).
Fig. 4. Chondrite-normalized (Anders & Grevesse 1989) REE-
contents in amphibole megacrysts from the Carpathian-Pannonian
region. Shaded field: amphibole megacrysts of Dobosi et al.
(2003a).
tematic relationships with major element and stable isotopic
compositions are detected (Figs. 5, 7). Fig. 5 shows that there
is a negative correlation between mg# and La/Sm ratios in the
amphibole megacrysts of the Pliocene-Pleistocene basalts.
The Tob-1 megacrysts have peculiar chemical and oxygen iso-
tope characteristics (see Figs. 5, 6, 9), but being only an outli-
er, this difference will not be discussed.
The stable isotope compositions also show some variations.
The δ
18
O values (Table 2) of amphibole megacrysts scatter
around the primary mantle compositions (5.5 , Mattey et al.
1994; Chazot et al. 1997) with an average of 5.45±0.23 (1 σ,
n = 15). These δ
18
O values are very similar to those reported by
Dobosi et al. (1998, 2003a) for amphibole megacrysts from the
alkaline basalts (5.27±0.15 , 1 σ, n = 8). Although the varia-
tions in the δ
18
O values are close to the analytical precision, the
184 DEMÉNY et al.
Sample
@
18
O
@D
H
2
O
Fe
3+
Tob-1
5.8
29
1.33
28.28
Tob-2
5.2
19
1.20
Tob-3
5.5
25
1.28
28.33
Gü-1
5.4
24
1.25
30.68
Gü-2
5.2
31
1.24
31.11
Bog-2001
5.5
29
1.11
M-2001
5.9
107
0.60
56.11
M-2002
5.6
61
1.04
36.63
MAS-1
5.1
31
0.82
34.47
MAS-2
5.1
75
0.87
40.69
MAS-3
5.4
56
0.97
31.79
Szil-1
5.5
15
0.99
46.45
Trs-2007
5.5
49
1.15
28.14
Trs-2008
5.5
59
1.05
29.20
Trs-2018
5.4
50
1.16
31.00
Szg-3024
61
1.00
29.34
Szg
61
1.11
27.79
Bb-2
52
1.29
32.96
Bo-20
81
1.22
Table 2: Oxygen and hydrogen isotope compositions, water content
(in weight %) and Fe
3+
contents (in % as 100×Fe
3+
/Fe
total
) of amphibole
megacrysts from alkaline basalts of the Carpathian-Pannonian region.
associated with low H
2
O contents, suggesting that degassing
may be responsible for the large δD variation observed. System-
atic δD differences between sampling localities do not exist.
One of the most important parts of the information needed
for the interpretation of hydrogen isotope compositions is the
Fe
3+
content of the amphiboles (Dyar et al. 1992; Feldstein et
al. 1996). Fe
3+
contents expressed as percentages (100 Fe
3+
/
Fe
tot
) also show a large scatter from 28 to 56 % (Table 2).
When the Fe
3+
contents are plotted as a function of H
2
O %, a
slight negative correlation appears (Fig. 8), that again points
to H
2
-degassing, which causes iron oxidation in the amphibole
structure (Feldstein et al. 1996; King et al. 1999). Similarly to
the hydrogen isotope compositions, the oxidation state of am-
phibole is sensitive to late stage processes, whose effects can
be superimposed on the primary features, thus a combined
evaluation is needed to explain the observed variations. The
hornblendite veins have chemical and isotopic characteristics
very similar to those of the megacrysts (see Tables 1 and 2),
thus they will not be discussed separately.
Discussion
Origin of
δδδδδ
18
O variation
The first question about the origin of amphibole megacrysts
is their primary mantle nature, which is best reflected by the
oxygen isotope composition. Although the δ
18
O range of the
megacrysts is very close to the normal mantle composition
with a small scatter, a systematic variation appears. The nega-
tive correlation between δ
18
O values and MgO contents
(Fig. 6) may be related to several processes: (i) crystal chemi-
cal effects, (ii) crustal contamination either by assimilation
during magma transport through the crust or by source con-
tamination by subducted material, (iii) fractional crystalliza-
tion, or (iv) partial melting. Since the oxygen isotope fraction-
ation due to Mg-Fe variation does not exceed 0.2 (Zheng
1993), the effect of crystal chemistry on the observed oxygen
Fig. 6. A δ
18
O values (in relative to V-SMOW) vs. mg num-
bers. B δ
18
O values vs. La/Sm ratios normalized to the primitive
mantle (Hofmann 1988). Legend as in Fig. 5.
Fig. 5. La/Sm ratios normalized to the primitive mantle (Hofmann
1988) in amphibole megacrysts vs. their mg numbers. Filled square:
this study, grey circle: Tob-1. Shaded field: amphibole megacrysts
of Zanetti et al. (1995) and Dobosi et al. (2003a).
isotope variation (0.5 , excluding Tob-1 amphibole on the
basis of its chemical compositions, see Results) can be ex-
cluded. Crustal contamination processes may explain the δ
18
O
ORIGIN OF AMPHIBOLE MEGACRYSTS IN THE PLIOCENE-PLEISTOCENE BASALTS 185
and MgO variation. Trace element compositions, especially
certain trace element ratios that are sensitive to crustal material
addition (Ce/Pb, Nd/Pb, Ba/Th, etc.) can attest the contamina-
tion hypothesis. However, no correlation between these ratios
and the δ
18
O values could be observed (Fig. 9), thus contami-
nation is unlikely as a cause of the δ
18
O-MgO variation. Frac-
tional crystallization of the basaltic magma can also produce
the observed changes provided that high-MgO and low-δ
18
O
minerals (e.g. olivine) are crystallized. Apart from the amphib-
ole megacrysts, the basalts contain numerous pyroxene, spinel
and olivine mega- and xenocrysts. Crystallization of forsteritic
olivine and Mg-Al spinel from the basaltic magma has been
modelled using MgO % = 15 wt. % and δ
18
O = 5.1 as start-
ing compositions, theoretical mineral compositions and pub-
lished mineral-melt oxygen isotope fractionations (Mattey et
al. 1994; Chazot et al. 1997). The calculations indicate that
about 30 % of the magma should have crystallized as olivine
and spinel in order to explain the observed δ
18
O-mg# varia-
tion. Zanetti et al. (1995) observed varying correlations be-
tween mg# and elements with different compatibility, which
they interpreted as resulting from fractional crystallization
from evolved melt. However, olivine and spinel crystallization
cannot explain the relationships between oxygen isotope com-
position and some trace element ratios (e.g. La/Sm, see Fig. 6)
considering the very low mineral-melt partition coefficients
for olivine and spinel (Beattie 1994; Horn et al. 1994). Due to
the low partition coefficients, crystallization of olivine and
spinel would not affect the trace element content of the re-
sidual melt significantly. Other minerals (e.g. pyroxene and
amphiboles) with higher mineral-melt coefficients have lower
MgO contents or higher δ
18
O values, resulting in lower min-
eral-melt oxygen isotope fractionation (Mattey et al. 1994;
Chazot et al. 1997). The degree of crystallization that would
be needed to explain the observed δ
18
O variations (>99 % for
pyroxene and >85 % for amphibole considering Rayleigh
fractionation) is unreasonably high.
Generation of basaltic magma due to a low degree partial
melting of a peridotitic source may result in elevated δ
18
O val-
ues in the basalt (Eiler 2001) and would also produce low
MgO content and elevated La/Sm ratio in accordance with the
relationships observed in this work. The melting process may
also produce volatile-rich magma. However, as we have seen
degassing may affect the amphiboles, thus the amphiboles
water contents may not reflect magma compositions. The
slight increase in halogen contents may also be related to the
melting process. However, the halogen contents of amphib-
oles strongly depend on crystal chemistry (Morrison 1991).
The negative F+ClMgO correlation can also be related to the
Mg-Cl avoidance, as suggested by Morrison (1991).
In summary, the oxygen isotope compositions of amphibole
megacrysts are in accordance with their mantle origin. Ac-
cording to relationships among major, minor and trace ele-
ment contents, the small δ
18
O variations observed may be ex-
plained in terms of varying degrees of partial melting.
Origin of mantle fluids
Hydrogen isotope compositions may provide additional in-
formation on the origin of the fluids responsible for amphibole
Fig. 7. δD values (in relative to V-SMOW) vs. H
2
O contents (in
wt. %) in amphibole megacrysts and hornblendite veins. See text for
degassing effects. Open circles: from Hurai et al. (1998) and J. Hoe-
fs (pers. comm.).
Fig. 8. Fe
3+
contents (in % relative to Fe
total
) vs. H
2
O contents (in
wt. %) in amphibole megacrysts. Open circles: from Hurai et al.
(1998) and J. Hoefs (pers. comm.).
Fig. 9. Primitive mantle normalized (Hofmann 1988) Ce/Pb ratios
vs. δ
18
O values (in relative to V-SMOW) in amphibole mega-
crysts. Legend as in Fig. 5.
186 DEMÉNY et al.
formation. To evaluate the obtained δD values, their primary
nature should be proved and the effects of secondary process-
es, such as weathering, alteration and degassing have to be
eliminated. Low temperature alteration can be ruled out owing
to the absence of secondary minerals (Fe-oxide-hydroxides,
chlorite, clay minerals) and elevated H
2
O content. Further,
preservation of primary mantle oxygen isotope signature and
correlations with trace element ratios argue against significant
alteration.
Degassing results in depletion in water in the amphibole and
increasing oxidation in the case of hydrogen release. The posi-
tive correlation between δD and water content indicates that
degassing may have indeed affected the studied amphiboles
(Fig. 7).
Degassing of amphibole is often associated with oxidation,
that is enrichment in Fe
3+
(Dyar et al. 1992; King et al. 1999).
This effect can be observed in the Fe
3+
-H
2
O plot (Fig. 8),
which shows a weak negative correlation between these vari-
ables. Dyar et al. (1993) have suggested another mechanism
for the water-deficient and oxidized nature of mantle kaersu-
tites. Considering the diffusion rate of hydrogen and mantle-
to-surface transport time of megacrysts, they found near-com-
plete degassing of large crystals difficult. Instead, they
proposed migration of oxidized metasomatic fluids in the
mantle. However, amphibole has a good cleavage, which de-
creases the effective grain size in term of diffusion. This is in-
dicated by host melt infiltration into cleavage planes observed
during hand-picking under binocular microscope. Thus, we
consider degassing a more likely process. Degassing may oc-
cur as dehydrogenation (H
2
-loss) and dehydration (partial
breakdown of amphibole). Dehydrogenation means preferen-
tial loss of the light H isotope, leaving the residual amphibole
enriched in deuterium (Dyar et al. 1992; Vennemann &
ONeil 1996). This process may explain the compositions of
the strongly D-enriched samples shown in Fig. 7, as the high-
δD group (>40 ) shows negative correlation in agreement
with dehydrogenation. On the other hand, H
2
O loss would
cause a negative δD shift in the amphibole (Suzuoki & Epstein
1976; Graham et al. 1984; Vennemann & ONeil 1996). This
process appears in the low-δD (<50 ) group in Fig. 7. On
the basis of the data plotted in Figs. 7, 8 and 10, the samples
likely to have suffered degassing (low H
2
O %, Fe
3+
/Fe
tot
>35 %) must be excluded from the evaluation of primary δD
characteristics of parental magma.
Fig. 11 shows the hydrogen and oxygen isotope composi-
tions of undegassed amphiboles. It is apparent that only some
samples fall close to the average upper mantle compositions
(δ
18
O = 5.5±0.2 , δD = 70±10 , Boettcher & ONeil
1980; Kyser & ONeil 1984; Mattey et al. 1994; Chazot et al.
1997; Javoy 1998). Hence, these samples may represent the
primary uncontaminated asthenospheric mantle. The low-δD
group partially overlaps the field of xenolith-hosted amphib-
oles from Southern Slovakia (Hurai et al. 1998). However, the
amphiboles with the lowest δD values have low H
2
O contents
(<0.7 wt. %, see also Fig. 7) indicating degassing, thus the
primary compositions for the xenolith-hosted amphiboles stud-
ied by Hurai et al. (1998) can be estimated at about 60 .
Although the megacrysts studied in this paper are character-
ized by mantle-like δ
18
O, they form two groups according to
their δD values (Figs. 10, 11). As shown in Fig. 10, the high-
D group cannot be produced from the low-δD one by dehy-
drogenation as they have identical Fe
3+
contents. Indeed, the
Fe
3+
content of 27 % seems to reflect an initial oxidation de-
gree characteristic of the mantle source. Thus, the δD values
also represent mantle compositions ranging from about 60 to
20 . The low δD limit is close to the normal mantle range,
but the high δD values are rather rare in terrestrial rocks and
are usually found in rocks, which have undergone seawater-
rock exchange (e.g. Stakes & ONeil 1982; Yui & Jeng 1990).
The constant δ
18
O values exclude contamination by seawater-
bearing sediments regarding the strong
18
O-enrichment in sed-
imentary rocks (δ
18
O~1530 , Hoefs 1987). The most like-
Fig. 10. δD values (in relative to V-SMOW) as a function of
Fe
3+
contents (in % relative to Fe
total
) in amphibole megacrysts.
Open circles: from Hurai et al. (1998). See text for degassing effects.
Fig. 11. δD vs. δ
18
O values (in relative to V-SMOW) in unde-
gassed amphibole megacrysts. Open circles: from Hurai et al.
(1998). Mantle ranges: see text.
ORIGIN OF AMPHIBOLE MEGACRYSTS IN THE PLIOCENE-PLEISTOCENE BASALTS 187
ly explanation of the D-enrichment in the mantle is subduction
of high-δD rocks (e.g. serpentinites in the lower oceanic crust)
and subsequent release of D-enriched fluids, which may have
metasomatized the upper mantle. Unfortunately no H-O iso-
tope study is available from oceanic crust relics from the East-
ern Alps and Carpathians. Serpentinites and ophicarbonates of
the Western Alps display large H isotopic variations (from
150 to 30 , Burkhard & ONeil 1988; Früh-Green et
al. 1990). The high δD values record a high-temperature inter-
action of the oceanic crust with seawater. According to the
serpentine-water H isotope fractionation (Wenner & Taylor
1973), the fluid released from the subducted serpentinite mass
would be even more enriched in deuterium (approaching
20 ) than the source rock. This fluid would enter the man-
tle causing the metasomatism and D-enrichment.
Serpentinite devolatilization may occur at mantle depths
(Scambelluri et al. 1995, 1997), carrying large amounts of
high-δD fluid into the mantle. An independent indication of
subducted oceanic crust material under the Carpathian-Pan-
nonian region is the presence of mafic granulite xenoliths with
low δ
18
O values and MORB-like Sr-Nd isotope compositions
in the alkaline basalts (Dobosi et al. 2003b). Subduction of
oceanic crust has been presumed to have taken place during
the Alpine orogenesis in the Carpathian-Pannonian region,
producing high-D metamorphic rocks (Demény et al. 1997)
with δD values very similar to those detected in this study.
It is important in the evaluation of the above model that an
areal variability has been observed in radiogenic isotope com-
positions of the Pliocene-Pleistocene alkaline basalts of the
CPR (Embey-Isztin & Dobosi 1995), which was related to an
increased amount of the subducted component in the central
part of the region. The constant oxygen isotope compositions
throughout the region indicate that the contamination was not
induced by direct melt addition from subducted crust, since
such process would have shifted the oxygen isotope composi-
tions to higher values. Our results indicate that the radiogenic
isotope variations may have been caused by metasomatism
triggered by release of fluid from the subducted slab.
Conclusions
Amphibole megacrysts and hornblendite veins from com-
posite mantle xenoliths of Pliocene-Pleistocene alkaline ba-
salts of the Carpathian-Pannonian region were studied in this
paper to investigate the origin of mantle fluids, which induced
hydrous mineral formation. Former studies have proven that
the megacrysts represent cumulate phases formed during an
early upwelling of basaltic magma, and were brought to the
surface by subsequent magma pulse. The δ
18
O values of the
amphiboles support their primary mantle origin. Correlations
with the MgO contents and La/Sm ratios are interpreted in
terms of varying degrees of partial melting and amphibole
crystallization from the evolved basaltic melt. The halogen
contents (F, Cl) of the amphiboles are very low, and their vari-
ations can either be related to the partial melting process, or to
the Mg-Cl avoidance rule.
Water and Fe
3+
contents and H isotope compositions indicate
degassing effects. Dehydrogenation is subordinate, whereas de-
hydration prevails in the studied amphibole megacrysts. Un-
degassed amphiboles show a wide δD range from normal
mantle compositions (about 60 ) to unusually high values
(20 ) attributed to metasomatism incidental to fluid release
from subducted oceanic crust. The trace element and O iso-
tope co-variations indicate negligible direct melt addition to
the mantle. Radiogenic isotope variations observed in previ-
ous studies may be related to fluid-mediated element transport
from the subducted slab.
Acknowledgments: The study was financially supported by
the Hungarian Scientific Research Fund (to A.D., OTKA
T 029078) and was conducted in the framework of a Bolyai
János Research Scholarship provided to A.D. The ICP-MS
work was carried out in an EC funded Large Scale Analytical
Facility (SOCFAC, contract no. HPRI-CT-1999-00108). H
2
O
wt. % data on Slovak amphiboles were kindly provided by
Prof. J. Hoefs. We gratefully thank J. Lexa, E. Bali and an
anonymous reviewer for constructive comments.
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