JOURNAL OF
Journal of Petrology, 2015, Vol. 56, No. 9, 1743–1774
PETROLOGY
Advance Access Publication Date: 13 October 2015
Original Article
doi: 10.1093/petrology/egv052
Magmatic Evidence for Carbonate
Metasomatism in the Lithospheric Mantle
underneath the Ohře (Eger) Rift
Philipp A. Brandl1,2*, Felix S. Genske1,3,4, Christoph Beier1,
Karsten M. Haase1, Peter Sprung5 and Stefan H. Krumm1
1
GeoZentrum Nordbayern, Friedrich-Alexander-Universität Erlangen–Nürnberg, Schloßgarten 5, 91054 Erlangen,
Germany, 2Research School of Earth Sciences, The Australian National University, 142 Mills Road, Acton, ACT
2601, Australia, 3CCFS, GEMOC, Department of Earth and Planetary Sciences, Macquarie University, Sydney,
NSW 2109, Australia, 4Institut für Mineralogie, Westfälische Wilhelms-Universität Münster, Corrensstr. 24, 48149
Münster, Germany and 5Institut für Geologie und Mineralogie, Universität zu Köln, Zülpicher Strasse 49b, 50674
Köln, Germany
*Corresponding author. Telephone: þ61 (0)2 6125 4301. Fax: þ61 (0)2 6125 8253. E-mail:
[email protected]
Received September 12, 2014; Accepted August 18, 2015
ABSTRACT
Magmas erupted in intracontinental rifts typically form from melting of variable proportions of asthenospheric or lithospheric mantle sources and ascend through thick continental lithosphere. This
ascent of magma is accompanied by differentiation and assimilation processes. Understanding the
composition of rift-related intracontinental volcanism is important, particularly in densely populated active rift zones such as the Ohře (Eger) Rift in Central Europe. We have sampled and analysed nephelinites from Železná hůrka (Eisenbühl), the youngest (<300 ka) Quaternary volcano
related to the Ohře Rift where frequent earthquake swarms indicate continuing magmatic activity
in the crust. This nephelinite volcano is part of a larger eruptive centre (Mýtina Maar) representing
a single locality of recurrent volcanism in the Ohře Rift. We present a detailed petrographic, mineralogical and geochemical study (major and trace elements and Sr–Nd–Hf–O isotopes) of Železná
hůrka to further resolve the magmatic history and mantle source of the erupted melt. We find evidence for a highly complex evolution of the nephelinitic melts during their ascent to the surface.
Most importantly, mixing of melts derived from different sources and of strong chemical contrast
controls the composition of the erupted volcanic products. These diverse parental melts originate
from a highly metasomatized subcontinental lithospheric mantle (SCLM) source. We use a combined approach based on mineral, glass and whole-rock compositions to show that the mantle
underneath the western Ohře Rift is metasomatized dominantly by carbonatitic melts. The nephelinites of Železná hůrka formed by interaction between a carbonatitic melt and residual mantle peridotite, partial crystallization in the lithospheric mantle and minor assimilation of upper continental
crust. Thermobarometric estimates indicate that the stagnation levels of the youngest volcanism in
this part of the Ohře Rift were deeper than the focal depths of recent earthquake swarms, indicating
that those are not directly linked to magma ascent. Furthermore, close mineralogical and
geochemical similarities between the Železná hůrka nephelinite and fresh kimberlites may point
towards a genetic link between kimberlites, melilitites and nephelinites.
Key words: continental rift; nephelinite; carbonated peridotite; mantle metasomatism; assimilation
C The Author 2015. Published by Oxford University Press. All rights reserved. For Permissions, please e-mail: [email protected]
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INTRODUCTION
The complex geology of Central Europe is a result of its
evolution throughout the Phanerozoic, which is dominated by several cycles of plate collision, plate reorganization and rifting. Major parts of the European
lithosphere were accreted during the Variscan Orogeny
and their tectonic fabrics still dominate the structure of
the Central European lithosphere (e.g. Ziegler et al.,
2006; Schulmann et al., 2014). Orogenic collapse followed during the Permian and the lithosphere became
stabilized during the Late Cretaceous. However, subsequent major plate reorganization related to the collision
of Africa and Europe (the Alpine Orogeny), together
with proposed mantle plume activity (e.g. Granet et al.,
1995; Hoernle et al., 1995) led to the formation of the
European Cenozoic Rift System (e.g. Ziegler et al., 2006)
and widespread Tertiary–Quaternary volcanism. This
Central European Volcanic Province (CEVP) encompasses intraplate and, more commonly, rift-related volcanism; for example, in the Massif Central at the
southern tip of the Limagne graben, the Kaiserstuhl in
the Upper Rhine graben, the Vogelsberg linking the
Rhine graben and the Hessian depression, and in the
Ohře (Eger) Rift. Quaternary volcanic activity in Central
Europe is reported from the Massif Central (e.g.
Downes, 1987), the Ohře Rift (e.g. Ulrych et al., 2011,
2013) and the Eifel province, with the eruption of the
Laacher See being the last major eruption in Central
Europe, dated to 129 ka (e.g. van den Bogaard, 1995;
Nowell et al., 2006). Although active volcanism in
Central Europe seems to be currently dormant, active
degassing and seismic activity indicate continuing magmatic activity in the Ohře Rift (e.g. Weinlich et al., 1999;
Spicák & Horálek, 2001; Bräuer et al., 2005).
Detailed studies of these volcanic systems are
required to fully understand the processes driving recent
magmatism in Central Europe. Lavas erupted in continental rifts, for example, provide insights into the composition of the underlying mantle, in particular the
subcontinental lithospheric mantle (SCLM). More importantly, such melts allow constraints on the processes during melting, melt extraction and ascent through the
continental lithosphere. Assimilation of crustal material
may play a major role in their petrogenesis and the
SCLM probably provides a chemically variable but overall more enriched source reservoir compared with the
asthenospheric mantle (e.g. Foley, 1992). It is thus crucial
to identify the complex processes that affect volcanism
in continental rifts and to unravel their distinct compositional imprints on the erupted volcanic rocks. Insights
into melting and magmatic differentiation, along with
thermobarometric estimates, are important to link past
volcanism with recent signs of volcanic activity.
Nephelinites are volumetrically minor contributors to
global magmatism, but are ubiquitous in continental
intraplate and rift settings. Thus, the understanding of
the mechanisms that produce these special melt types is
of wide-ranging interest.
Journal of Petrology, 2015, Vol. 56, No. 9
Recently, several studies focused on the nature and
composition of the mantle source of alkaline volcanism
related to the Ohře Rift by studying mantle xenoliths
and cumulates entrained in the lavas (Geissler, 2005;
Geissler et al., 2007; Puziewicz et al., 2011; Ackerman
et al., 2013, 2014; Špaček et al., 2013). These studies
have shown that the lithospheric mantle underneath the
rift is highly heterogeneous and often overprinted by
metasomatic processes related to alkaline or carbonatitic melt infiltration. As a result, mantle lithologies are
highly variable, ranging from fertile, clinopyroxene-rich
lithologies [e.g. the ‘pyroxenite suite’ of Puziewicz et al.
(2011)] to refractory harzburgites and dunites (e.g.
Puziewicz et al., 2011; Ackerman et al., 2013). Accessory
minerals such as apatite, phlogopite, ilmenite or rutile
have been identified from various locations within the
Ohře Rift (e.g. Puziewicz et al., 2011; Ackerman et al.,
2013, 2014). Direct evidence for carbonatitic melts has
been recorded from carbonate minerals and melt pockets in mantle xenoliths from the Oberpfalz (NE Bavaria;
Ackerman et al., 2013; Špaček et al., 2013), Plešný Hill
(Ackerman et al., 2014) and Ksie˛ginki (Puziewicz et al.,
2011), all of which are located west (Oberpfalz) or east
(Plešný Hill and Ksie˛ginki) of the main Ohře Rift structure. The metasomatic processes leading to these modifications of the sub-rift mantle are likely to be multistage events (e.g. cryptic Fe-metasomatism followed by
thermal rejuvenation and melt infiltration) and will thus
differ between distinct volcanic centres related to the
Ohře Rift (e.g. Puziewicz et al., 2011).
However, because xenoliths may sample preferentially the cool lithosphere rather than the actual magma
source and are not entrained in every volcanic system,
we demonstrate that the petrology and geochemistry of
volcanic minerals and rocks themselves can be used to
constrain magmatic processes and the mantle source
composition. For this purpose we focused on Železná
hůrka (Eisenbühl), a young (<300 ka) volcano situated
at the eastern tip of the Ohře Rift that consists predominantly of fresh tephra. We present high-precision geochemical analyses of fresh volcanic glasses and show
that these differ significantly from the compositions of
whole-rock samples from the same locality and other
volcanoes nearby. Mineral textures and compositions
indicate a complex ascent history for the lava, including
the assimilation of crustal material. We determine the
conditions of melting and the magmatic evolution of
the erupted lavas and present further insights into volcanic processes and the nature of the mantle beneath
the Ohře Rift by combining major and trace element
data with radiogenic (Sr–Nd–Hf) and stable O isotope
data.
GEOLOGICAL SETTING
Volcanism related to the Ohře Rift extends for more
than 400 km from northeastern Bavaria through the
Czech Republic into Poland. The rift itself represents an
Journal of Petrology, 2015, Vol. 56, No. 9
1745
ENE–WSW-trending extensional structure, about 25–
30 km wide, that follows Variscan crustal lineaments between the Saxothuringian terrane in the NW, the
Moldanubian terrane in the SE and the TeplaBarrandian terrane between those two terranes
(Malkovský, 1987; Babuška & Plomerová, 2010).
In the region of the Cheb Basin, the crust is thinned to
25–28 km (e.g. Geissler, 2005; Heuer, 2006). The
main phase of rift-related volcanism is dated to about
30–15 Ma with episodic volcanism extending to 026 Ma
(Ulrych et al., 2011, and references therein). Quaternary
volcanic rocks are present at Komornı́ hůrka
(Kammerbühl) and the volcanic system of the Mýtina
Maar and Železná hůrka (Fig. 1a), both located roughly
above the locus of crustal thinning (Babuška &
Plomerová, 2010).
Komornı́ hůrka is a 726 6 59 kyr old volcano (Wagner
et al., 2002) that has erupted sodalite-bearing or nepheline–olivine-melilitite as scoria and a lava flow (Ulrych
et al., 2013). In contrast, Železná hůrka is younger and
consists of three eruptive units: volcaniclastic material
formed by a phreatomagmatic eruptive phase at its
base, overlain by highly olivine-phyric lava in the vent,
and tephra layers at the top (Fig. 1b). Recent studies
(e.g. Geissler et al., 2004, 2007; Mrlina et al., 2009) that
combined geophysical studies with geochemistry and
information from scientific drilling found the remnants
(a)
D
N
PL
CZ
12°E
14°E
Ch
AT
16°E
OPF
M
Cenozoic
volcanic
rocks
MLF
ZH
Cenozoic
sediments
(b)
S
KH
50°N
49°N
D
DHM
ft
r Ri
Ege
51°N
CZ
WBSZ
0 km 20
40
FL
appr. 5 m
nt)
a
av
L
V3
V2
V1
Upper teph
(ve
ra
L3
L1
L2
Lower
tephra
U3
U2
U2f
of a larger eruption diatreme, the Mýtina Maar, just
north of Železná hůrka, which was dated by the Ar–Ar
method at 288 6 17 ka (Mrlina et al., 2007). Additional
evidence for continuing magmatic activity close to the
volcano are active degassing of CO2, mantle-derived He
(high 3He/4He) in numerous mofette fields (e.g. at Soos
or Bublák; Weinlich, 2013) and recurrent earthquake
swarms (e.g. Fischer & Horálek, 2003). These earthquake swarms have a focal depth of around 6–11 km
and are either linked to the ascent, accumulation or
stagnation of magma in the crust (e.g. Dahm et al.,
2008) or, alternatively, may be explained by fluids ascending along pre-existing and reactivated fault planes
(e.g. Bankwitz et al., 2003). Interestingly, the 3He/4He in
the gas exhalations increased significantly from 1993 to
2005, reaching values of 63 Ra that are similar to the
average of the SCLM (Bräuer et al., 2009). This increase
was interpreted as evidence for the ascent of mantlederived melts into the lithosphere beneath the western
Ohře Rift, with deep dike intrusions in 2006–2008
(Bräuer et al., 2005, 2009). Thus, both seismic data and
the active degassing in the western Ohře Rift suggest
continuing magmatic activity at depth.
METHODS
Samples L1 to L3 were collected from a basal, brownish
phreatomagmatic tephra unit (lower tephra; Fig. 1b)
and samples V1 to V4 along a traverse directly
above sample L1 towards the west of the outcrop (vent;
Fig. 1b). Samples U1, U2 (with a fine-grained variety
U2f), and U3 were collected from the base, the middle
and upper layer of the upper tephra unit (Fig. 1b). An
additional sample (EG0661) had previously been collected from the upper tephra. For geochemical analyses
of whole-rocks, xenolith- and crystal-poor samples
were selected. Weathered surfaces were removed prior
to crushing and the rocks were rinsed in de-ionized
water. Splits of the crushed materials were further processed for glass and mineral separation and representative whole-rock pieces were cut for thin sections. A split
of sample V1 (massive but highly olivine-phyric lava)
was crushed by high-voltage pulse power fragmentation in the laboratories of Selfrag AG, Kerzers
(Switzerland), to separate olivine crystals for major and
trace element and O isotope analyses.
U1
V4
Fig. 1. (a) Map of the western Ohře Rift at the structural boundary between the Variscan Saxothuringian and Moldanubian
terranes in the Czech–German border region. OPF, Oberpfalz;
KH, Komornı́ hůrka; ZH, Železná hůrka; DHM, Doupovské hory
Mountains; M, Mitterteich Basin; Ch, Cheb Basin; S, Sokolov
Basin. Major structural features: MLF, Mariánske Láznĕ Fault;
WBSZ, West Bohemian Shear Zone; FL, Franconian Line. (b)
Schematic cross-section of the outcrop at Železná hůrka, showing its lithological structure and sample locations. Lx samples
represent the lower tephra unit; Vx samples are from the vent;
Ux samples are from the upper tephra unit.
Major elements
Major element analyses of glasses and minerals were
performed on a JEOL JXA-8200 electron microprobe at
the GeoZentrum Nordbayern, Friedrich-AlexanderUniversität Erlangen–Nürnberg. Glasses were analysed
using an acceleration voltage of 15 kV, a beam current
of 15 nA and a defocused beam of 10 mm diameter.
Further details of the analytical conditions have been
given by Brandl et al. (2012). Major element compositions of minerals (olivine, clinopyroxene, spinel,
phlogopite and minerals of crustal xenoliths) were
1746
determined using 20 kV acceleration voltage, 20 nA
beam current and a focused beam.
Major element analyses of whole-rock samples (and
the trace elements Ba, Cr, Ga, Nb, Ni, Rb, Sr, V, Y, Zn
and Zr) were carried out by X-ray fluorescence (XRF) on
a Spectro XEPOS plus at the GeoZentrum Nordbayern.
Further details of the analytical technique have been
given by Freund et al. (2013).
Trace elements
The trace element analyses of whole-rock powders (solution inductively coupled plasma mass spectrometry;
ICP-MS) were performed at the Geochemical Analysis
Unit (GAU) at Macquarie University, Sydney. The same
analytical protocol was followed as presented by
Genske et al. (2012). Approximately 100 mg of sample
powder was dissolved using a mix of HNO3–HF–HCl
acids for digestion. To fully dissolve Fe–Ti oxides, a
mixture of HCl and HClO4 was also used throughout the
digestion procedure. The analytical data for the samples and international rock standards were obtained on
an Agilent 7500 c/s quadrupole ICP-MS system. All analytical results (whole-rocks and glasses), including reproducibility, are reported in Table 1.
Laser ablation (LA)-ICP-MS analyses of volcanic
glasses were carried out at the GeoZentrum
Nordbayern on a New Wave Research UP193FX laser
ablation system coupled to an Agilent 7500i quadrupole
ICP-MS system. We averaged our results over at least
four spots from each sample. Laser ablation was carried
out on 25 mm spots with 072 GW cm–2 laser energy and
36 J cm–2 energy density, measuring the background
for 25 s and the sample for 30 s. Lithium, Si and Mn
were analysed for 10 ms on the maximum peak and
other elements for 25 ms (30 ms for Ta), resulting in a
total of 10082 s per mass scan. Silica concentrations
determined by electron microprobe were used for internal calibration and standard glass NIST 612 was
used for external calibration (Pearce et al., 1997).
Accuracy and reproducibility were checked using secondary standards NIST 614 and BCR-2 g. Precision and
accuracy are generally better than 10%, with slightly
higher values for Cr, Nd and Hf (1404, 1193 and
1133% RSD, respectively). Laser ablation conditions for
the trace element analyses of minerals (olivine, clinopyroxene, phlogopite, quartz, spinel) were similar to the
conditions for glasses described above, but using a
spot size of 50 mm (for small minerals or crystal rims of
25 mm; see Table 2 for further details) with 066 GW cm–2
laser energy and 33 J cm–2. Results of representative
mineral analyses are reported in Table 2 (complete data
for mineral analyses can be found in Supplementary
Data Table S1; supplementary data are available for
downloading at http://www.petrology.oxfordjournals.
org).
Detailed comparison and evaluation of trace element
data determined on rock standards by both LA-ICP-MS
and solution ICP-MS reveal that selected elements and
Journal of Petrology, 2015, Vol. 56, No. 9
corresponding element ratios deviate to slightly higher
values than reported in the literature. In particular, the
high field strength element (HFSE) Zr is determined to
be up to 8% higher in BHVO-2 than for high-precision
isotope dilution data presented on the same USGS
standard by Pfänder et al. (2007). However, we note that
in comparison with preferred GeoReM data, our data
agree within one standard deviation (i.e. <5%), which is
the commonly achieved precision for the techniques
employed here (Table 1). Nevertheless, taking the maximum error into account would still result in relatively
high Zr/Sm and low Zr/Nb for the Železná hůrka lavas
compared with the Ohře Rift (see discussion below).
Radiogenic isotope analyses
Strontium and Nd isotopes in whole-rocks and glasses
were analysed at the GAU, Macquarie University,
Sydney. The analytical routine applied is the same as
described by Genske et al. (2012). The isotopic analyses
of Sr and Nd were conducted using thermal ionization
mass spectrometry (TIMS) employing a Thermo
Finnigan Triton system. Measured 87Sr/86Sr ratios for
BHVO-2 obtained during this study are listed in Table 1.
The standard NIST SRM 987 was analysed (n ¼ 17) to
verify the accuracy of the measurements during the
period of sample analysis. The long-term reproducibility of NIST SRM 987 is 87Sr/86Sr ¼ 0710250
(2SD ¼ 0000034).
Ratios
were
normalized
to
86
Sr/88Sr ¼ 01194 to correct for mass fractionation. For
the Nd analyses, reference materials BHVO-2 and JMC
321 were analysed, yielding 143Nd/144Nd ratios close to
published values (Table 1). The external precision was
determined using JMC 321 (n ¼ 15), which yielded
143
Nd/144Nd ¼ 0511115 (2SD ¼ 0000047). Ratios were
normalized to 146Nd/144Nd ¼ 07219 to correct for mass
fractionation.
Determinations of Hf isotope compositions of wholerocks and combined determinations of Hf isotope compositions and Lu and Hf concentrations of glasses by
isotope dilution and spike stripping using a mixed
176
Lu–180Hf tracer were conducted at the WWU
Münster, Germany. Sample preparation, mass spectrometry on a Neptune Plus multicollector (MC)-ICP-MS
system at the Universität Münster, and estimation of
measurement and spike-stripping uncertainties followed the method of Sprung et al. (2010, 2013), which
makes use of a modified Ln-spec Hf purification scheme
after Münker et al. (2001). To minimize uncertainties
associated with possible isobaric interferences from
spiked Lu, a final Lu removal step using AG 50 W-x8
(1 ml) was added in which Hf was immediately eluted
upon loading in 05 M HCl–005 M HF, leaving Lu adsorbed on the resin. Mass bias was internally corrected
for using the exponential law described by Russell
et al. (1978) and normalizing to 179Hf/177Hf ¼ 07325.
All 176Hf/177Hf values are given relative to
176
Hf/177Hf ¼ 028216 for Ames Hf, which is isotopically
indistinguishable from JMC-475 (Scherer et al., 2000).
Journal of Petrology, 2015, Vol. 56, No. 9
1747
Table 1: Major and trace element and isotope analyses of samples from Železná hůrka and international reference material
L1
lower
tephra
Glass
EMPA
SiO2 (wt %)
TiO2 (wt %)
Al2O3 (wt %)
FeOt (wt %)
Fe2O3t (wt %)
MnO (wt %)
MgO (wt %)
CaO (wt %)
Na2O (wt %)
K2O (wt %)
P2O5 (wt %)
S (ppm)
Cl (ppm)
LOI (wt %)
Total (wt %)
ppm
Li
Sc
V
Cr
Mn
Co
Ni
Cu
Zn
Ga
Rb
Sr
Y
Zr
Nb
Sn
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
W
Pb
Th
U
87
Sr/86Sr
143
Nd/144Nd
176
Hf/177Hf
d18O
dupl.
L2
lower
tephra
Glass
EMPA
L3
lower
tephra
Glass
EMPA
4068
319
1434
1068
4027
293
1458
1070
3988
308
1478
1048
025
554
1523
460
375
129
1294
3167
027
480
1456
525
406
133
1375
3426
025
503
1464
457
403
138
1210
3318
10012
9936
9868
LAICP-MS
LAICP-MS
130
221
560
680
2244
637
538
101
214
974
1528
285
275
161
298
141
1241
989
186
203
768
132
387
941
114
662
106
260
0338
218
0286
561
880
234
631
121
361
0703891
0512812
0283002
625
613
123
651
471
2028
385
1048
117
158
116
1769
307
255
193
183
141
1515
117
218
231
860
140
410
978
126
671
116
286
0352
220
0296
426
104
252
748
147
447
0703631
0512758
0283008
567
565
544
569
V1
vent
lava
WR
XRF
V2
vent
lava
WR
XRF
V3
vent
lava
WR
XRF
V4
vent
lava
WR
XRF
U1
upper
tephra
WR
XRF
3978
298
1137
3994
297
1145
4035
299
1158
4048
291
1131
4155
282
1161
1269
0215
1263
1305
329
199
0774
1270
0214
1250
1286
346
229
0763
1220
0216
1254
1290
308
206
0791
1242
0211
1268
1280
332
212
0744
1223
0205
1212
1255
268
235
0735
075
9952
041
9955
084
9954
053
9952
068
9952
sol.
ICP-MS
sol.
ICP-MS
sol.
ICP-MS
709
333
308
583
689
328
311
521
667
287
317
565
562
214
881
102
192
602
903
271
283
121
556
222
664
99
189
628
779
269
281
120
562
213
778
101
186
512
736
256
280
122
0813
779
735
138
162
600
106
306
838
114
526
0918
227
0731
757
697
133
155
587
103
295
806
111
508
0897
219
0604
823
721
136
160
591
105
302
811
111
509
0891
217
165
0231
566
587
162
0226
564
568
160
0220
562
573
317
917
232
0703455
0512861
0282997
142
862
234
0703479
051286
0283003
206
842
241
0703591
0512862
0283000
XRF
XRF
294
549
286
518
229
208
103
196
653
894
275
264
101
972
164
801
966
27
269
997
806
872
210
800
U2
upper
tephra
Glass
EMPA
3974
308
1417
989
024
537
1408
443
369
120
1124
2994
9642
600
760
0703404
0512855
557
(continued)
1748
Journal of Petrology, 2015, Vol. 56, No. 9
Table 1: Continued
U2f
upper
tephra
Glass
EMPA
SiO2 (wt %)
TiO2 (wt %)
Al2O3 (wt %)
FeOt (wt %)
Fe2O3t (wt %)
MnO (wt %)
MgO (wt %)
CaO (wt %)
Na2O (wt %)
K2O (wt %)
P2O5 (wt %)
S (ppm)
Cl (ppm)
LOI (wt %)
Total (wt %)
ppm
Li
Sc
V
Cr
Mn
Co
Ni
Cu
Zn
Ga
Rb
Sr
Y
Zr
Nb
Sn
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
W
Pb
Th
U
87
Sr/86Sr
143
Nd/144Nd
176
Hf/177Hf
d18O
dupl.
4052
310
1475
1060
026
513
1470
455
404
138
1067
3359
9956
U3
upper
tephra
WR
XRF
EG-0661
upper
tephra
Glass
EMPA
4070
288
1135
1262
0211
1252
1298
266
241
0744
679
935
469
229
2017
395
1247
727
154
110
1743
331
281
190
173
132
1491
119
214
231
877
149
411
110
139
731
121
302
0402
238
0349
516
107
241
692
155
432
0283009
577
BE-N
BR
EMPA
XRF
XRF
3863
267
1003
3851
266
1003
1293
0202
1296
1401
316
140
106
1314
0200
1312
1362
299
137
104
244
9948
281
9948
4051
311
1477
1048
5053
179
1383
1163
027
502
1452
452
401
135
1227
3351
021
699
1118
267
019
021
1427
305
9911
9962
042
9949
LAICP-MS
VG-2
n¼6
XRF
302
545
225
989
164
67
1013
247
265
100
870
490
460
NIST-614
BCR-2G
BHVO-2
Absol. dev.
to GeoReM
LAICP-MS
LAICP-MS
LAICP-MS
sol.
ICP-MS
123
751
435
489
1851
373
109
959
146
191
159
0999
117
149
0792
0912
355
223
981
351
439
169
1526
392
124
146
157
415
327
305
256
0653
0677
119
241
102
1561
277
249
171
194
136
1325
103
195
207
773
127
378
914
118
632
105
263
0352
216
0290
498
904
215
704
1275
393
0874
443
0737
0755
0796
168
0682
315
0688
0769
0761
0727
0718
0693
0813
0639
0667
0699
0713
0680
0741
0677
0746
0781
0852
234
0702
0764
437
124
132
102
205
953
389
284
180
189
127
465
533
131
151
0420
721
240
767
0829
0103
133
158
381
551
248
625
200
635
0960
533
101
259
0003
169
0571
0563
0162
0323
0182
0068
0114
0040
0024
0032
0054
198
028
426
108
0023
0002
0103
0057
492
323
337
170
119
211
111
640
242
505
631
274
635
181
618
0956
634
126
353
0515
347
0512
458
0773
0526
103
577
167
156
128
043
0703459
0512978
0283099
0045
0064
0025
60000012*
60000015*
60000004*
548
557
(continued)
Journal of Petrology, 2015, Vol. 56, No. 9
1749
Table 1: Continued
BIR-1
Absol. dev.
to GeoReM
SiO2 (wt %)
TiO2 (wt %)
Al2O3 (wt %)
FeOt (wt %)
Fe2O3t (wt %)
MnO (wt %)
MgO (wt %)
CaO (wt %)
Na2O (wt %)
K2O (wt %)
P2O5 (wt %)
S (ppm)
Cl (ppm)
LOI (wt %)
Total (wt %)
Oxygen isotope analyses
sol. ICP-MS
ppm
Li
Sc
V
Cr
Mn
Co
Ni
Cu
Zn
Ga
Rb
Sr
Y
Zr
Nb
Sn
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
W
Pb
Th
U
87
Sr/86Sr
143
Nd/144Nd
176
Hf/177Hf
d18O
dupl.
*2SD relative to BHVO-2.
International rock standard BHVO-2 was analysed for
its Hf isotope composition alongside the unknowns,
yielding a 176Hf/177Hf composition of 0283099
(2SD ¼ 0000004).
312
450
317
0078
198
247
525
179
120
674
150
0191
105
168
153
0547
0475
127
128
465
0331
0009
418
121
133
0003
0005
652
0617
190
0387
233
110
0487
189
0353
251
0576
172
0002
0622
0002
0016
0017
0046
0021
0043
0017
0007
0004
0016
0065
159
0239
0566
0036
0055
0011
0016
0001
322
0039
0011
0115
0007
0001
Oxygen isotope analyses were performed on fresh and
visibly inclusion-free olivines and glasses. Single grains
were handpicked and cleaned and then coarsely
crushed in a steel mortar to obtain multiple splits. Small
chips were embedded for electron microprobe analysis
(EMPA) and LA-ICP-MS analyses and the remaining material (2–3 mg) was used to determine the O isotope
composition. Oxygen isotope analyses were performed
using the 25 W-Synrad CO2-laser fluorination line at the
GeoZentrum Nordbayern following the methods
described by Haase et al. (2011) and Genske et al.
(2013). The long-term reproducibility of the UWG-2
garnet standard obtained during the course of this
study is 584 6 007% (1SD, n ¼ 29).
RESULTS
Samples from the lower tephra are of lapilli size, with
brownish-weathered surfaces but fresh and glassy interiors. Blocky lavas in the western part of the outcrop
are situated in the position of the former vent and appear unaltered and fresh. These lavas contain abundant
xenoliths and up to centimetre-sized crystals of olivine,
clinopyroxene and phlogopite. The upper tephra
reaches a thickness of about 5–7 m at the eastern side
of the outcrop and is uniformly dark. The crystal assemblages in this unit are similar to those found in the vent
lavas. Xenoliths in the upper tephra are of variable size
ranging from millimetre-scale to more than half a metre
and show reddish, oxidized surfaces. Crustal xenoliths
are common in all the lithological units sampled.
Petrography and mineral composition
We selected representative samples for petrographic
description by optical examination under a binocular
microscope (Fig. 2). Samples L1 to L3 were collected
from the lower brown tephra (Fig. 1b) and have a glassy
but phenocryst-rich groundmass. Samples V1 to V4
(massive lava flow in the vent) and U1 to U3 (upper
tephra) have a holo- to cryptocrystalline matrix and
about 30 vol. % vesicles (Fig. 2a). The mineral assemblages of these samples are similar and include about
20 vol. % olivine, 10–15 vol. % clinopyroxene and accessory spinel and haüyne. Within their volcanic matrix the
lavas host felsic crustal xenoliths (Fig. 2b) and ultramafic cumulates (clinopyroxene and phlogopite and/or
olivine) (Fig. 2a–c).
Olivine
Olivine is by far the most abundant mineral in the
Železná hůrka lavas. Crystals of olivine range in size
from a few micrometres to 10 mm. Single crystals are
1750
Journal of Petrology, 2015, Vol. 56, No. 9
Table 2: Representative major and trace element analyses of mineral phases and one interstitial glass from Železná hůrka
Mineral:
Sample:
Spot size:*
wt %
SiO2
TiO2
Al2O3
Cr2O3
FeOt
MgO
MnO
CaO
Na2O
K2O
NiO
P2O5
SO2
Total
ppm
Li
Sc
V
Cr
Mn
Co
Ni
Cu
Zn
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Pb
Th
U
d18O
Min. form.
Si ¼
Ti ¼
Al ¼
Cr ¼
Fe2þ ¼
Fe3þ ¼
Mg ¼
Mn2þ ¼
Ca ¼
Na ¼
K¼
Ni
P¼
Cat. ¼
Fo/Mg#
Wollast.
Enstatite
Ferrosilite
Olivine
V4-17
core
4056
b.d.l.
006
005
965
5014
012
014
001
b.d.l.
035
001
Olivine
V4-17
rim
4010
005
005
001
1323
4658
033
079
b.d.l.
001
005
002
Olivine
U1-Ol11
4052
001
003
004
1056
4997
016
018
b.d.l.
b.d.l.
024
Cpx
V1-12 I
5018
113
674
027
539
1527
013
2101
092
b.d.l.
002
002
Cpx
V1-12 VI
Phl
V1-12 I
4887
117
750
038
542
1456
011
2039
099
b.d.l.
b.d.l.
003
3827
421
1729
032
706
2006
004
003
070
924
008
001
10109
10122
10170
10107
9942
9730
209
310
592
540
1188
146
3597
276
762
b.d.l.
b.d.l.
0023
0059
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
448
634
529
340
2122
163
1018
189
982
b.d.l.
0115
0076
0289
0020
b.d.l.
0129
0010
0038
0003
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
0005
0028
b.d.l.
0051
0006
0028
b.d.l.
0082
b.d.l.
b.d.l.
203
356
567
325
1439
159
2190
293
796
b.d.l.
b.d.l.
0034
0034
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
0010
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
0055
b.d.l.
b.d.l.
524
100
775
355
3022
1053
336
168
176
219
0024
777
888
417
0436
b.d.l.
0098
300
113
205
109
317
0953
277
0398
227
0357
0878
0113
0641
0073
231
0125
0172
0063
0012
0999
773
375
2992
1112
342
169
182
220
0062
758
881
418
0452
b.d.l.
0514
277
108
189
106
302
102
286
0374
214
0367
0807
011
065
0085
217
0146
0129
0055
0012
142
917
424
3071
274
788
668
186
462
332
115
0044
457
888
165
2731
0007
0016
b.d.l.
b.d.l.
b.d.l.
0046
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
0117
0687
0431
b.d.l.
b.d.l.
099
000
000
000
020
000
182
000
000
000
000
001
301
903
099
000
000
000
027
000
171
001
002
000
000
000
301
863
098
000
000
000
021
000
181
000
000
000
000
000
302
894
182
003
029
001
016
000
083
000
082
006
000
000
403
835
452
457
91
181
003
033
001
017
000
080
000
081
007
000
000
403
827
454
451
94
529
044
282
004
082
000
413
001
000
019
163
001
1564
835
Haüyne
V1 H5
3199
006
2644
001
060
058
n.d.
816
1123
787
n.d.
025
1031
9750
Spinel
V4-5 I
009
132
2263
4438
1895
1597
040
b.d.l.
b.d.l.
b.d.l.
b.d.l.
017
Glass
V1 glass
25 mm
6896
071
643
b.d.l.
350
139
012
219
615
1049
002
012
10391
10006
119
268
930
n.d.
1271
184
1264
716
437
0025
0016
b.d.l.
0803
0446
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
0034
0025
0202
b.d.l.
b.d.l.
274
358
585
479
728
826
412
533
774
153
216
841
339
241
163
431
190
298
315
120
181
0846
191
0245
184
0288
0708
0100
0759
0065
0637
0968
283
174
0514
The full data table is given in Supplementary Data Table S1. b.d.l., below detection limit. n.d., not determined.
*If not indicated otherwise, size of LA-ICP-MS spot is 50 mm.
Journal of Petrology, 2015, Vol. 56, No. 9
1751
(a)
(b)
UCC xeno
UCC
xeno
cpx+
phlog
ol
qtz
qtz
2 mm
ZH-V1
ol
cpx+
phl
2 mm
ZH-V1
(d)
(c)
cpx
phl
qtz
ol
Tiaug
di
UCC xeno
500 µm
ZH-V1
500 µm
ZH-L3
Fig. 2. Representative thin-section photographs. (a) Sample V1 plane-polarized light and (b) cross-polarized light. The clinopyroxenes–phlogopite cumulate is shown enlarged in (c) (plane-polarized light). (d) Olivine showing a dissolution texture and adjacent
crustal xenolith fragments; cross-polarized light.
generally idiomorphic to hypidiomorphic. Larger crystals are normally zoned (e.g. Fig. 3a) with well-defined
cores and overgrowth rims and sometimes dissolution
textures at their grain boundaries (Fig. 2d). Small clinopyroxene phenocrysts often crystallize around these
olivine grains (Fig. 3e). Smaller groundmass crystals
lack chemical zoning. In terms of composition, olivine
displays a bimodal distribution in forsterite content between crystal cores (Fo90) and overgrowth rims and
groundmass crystals (Fo86; inset Fig. 4). These two
groups are also distinct in their Ni (Fig. 4a), Ca (Fig. 4c)
and Mn (not shown) concentrations. Cores have up to
3000 ppm Ni, between 800 and 1500 ppm Mn and 900–
1300 ppm Ca. The olivine rims and smaller olivine crystals from the matrix have lower Ni contents of between
400 and 1300 ppm, but higher Mn (1300–2500 ppm) and
Ca concentrations (1100–4000 ppm; some up to
8300 ppm). Representative mineral analyses and data
for olivine separated from sample V2 (Ol-1 to Ol-18) are
presented in Table 2.
The olivine separates were also analysed for their O
isotope composition. Grains with compositional heterogeneity (i.e. zoning) where the variability in forsterite
content exceeds 605 were excluded from the O isotope
analysis. The isotopic composition of O in these olivines
ranges from d18O ¼ þ50 to þ57% V-SMOW, with the
full range of variability present in homogeneous highforsterite olivines (Fo887–902). Three separated olivine
grains have lower forsterite contents (Fo836) than
those reported from thin-section EMPA work, but have
a corresponding d18O of about þ52%. Applying a typical fractionation factor between olivine and silicate
melt of 04% (e.g. Eiler, 2001), we find that melts in
equilibrium with the olivines should have d18O values of
þ54 to þ61%, in good agreement with the O isotope
data obtained for the glass samples (see below).
1752
Journal of Petrology, 2015, Vol. 56, No. 9
(a)
(b)
(c)
(d)
qtz
gl
di
spongy
reaction
zone
Ti-aug
(e)
sp
cpx
laths
(f)
Ti-aug
ol
phl
ol
Fig. 3. Electron microprobe backscattered electron (BSE) images of key petrological features of the Železná hůrka lavas. (a) Olivine
crystals with forsterite-rich cores (Fo 90) and slightly more Fe-rich margins (Fo 86). (b) Zoned glomerophyric clinopyroxene. (c)
Upper crustal xenolith composed of quartz (dark grey), potassic feldspar (grey) and accessory muscovite. (d) Contact between a
single (disaggregated) quartz crystal and a cumulate composed of intergrown phlogopite (not shown) and clinopyroxene (Ti-augite
with diopsidic overgrowth). The presence of mingled interstitial glass (dotted area) and the spongy reaction zone in the clinopyroxene cumulate should be noted. (e) Rounded phlogopite hosted in olivine. The clinopyroxene laths oriented along the edge of the
olivine crystal should be noted. (f) Cumulate of clinopyroxene (Ti-augite with diopsidic overgrowth rim) and olivine.
Clinopyroxene
Clinopyroxenes generally occur as microphenocrysts,
as glomerocrysts that show zonation (Fig. 3b) or around
olivine grains (Fig. 3e), or as crystal cumulates intergrown with phlogopite (e.g. in sample V1, Figs 2a–c and
3d) or olivine (e.g. sample U3, Fig. 3f). One cumulate of
clinopyroxene (size of single crystals about 2–3 mm)
and phlogopite found in sample V1 is about 7–8 mm in
diameter, but cumulates and megacrysts in volcanic
bombs of the nearby Mýtina Maar can reach several
centimetres in size (Geissler, 2005; Geissler et al., 2007).
Spongy reaction textures are especially visible within
and around the rims of the Ti-augite rich clinopyroxenes (see spongy ‘fractures’ in Fig. 3d). Diopside occurs
as zones around Ti-augite and shows largely idiomorphic overgrowth textures (Fig. 3d).
Journal of Petrology, 2015, Vol. 56, No. 9
1753
Fig. 4. (a) Forsterite vs Ni (ppm) content in olivine (grouped into analyses of cores, ‘intermediate’, rim and groundmass) analysed in
this study compared with (b) literature data [literature data also plotted in (a) as grey symbols]. The ‘intermediate’ group defines
spot analyses between the clearly defined core and overgrowth rim to test for any hidden chemical zonation. (c) and (d) show forsterite content vs CaO (wt %) concentration in olivines from this study and literature data, respectively. The inset in (c) shows a
histogram of olivine compositions analysed in this study. Blue lines indicate the evolution of olivine composition during fractional
crystallization. We used partition coefficients from Beattie (1994) for Fe (051–155), Mg (196–440) and Ca (00192–00375), and
from Seifert et al. (1988) for Ni (38–60). We chose the high end of the range for all partition coefficients except for Fe, for which we
used a partition coefficient of 132 to match an Mg–Fe exchange coefficient of 03, typical for a wide range of basaltic liquids
(Roeder & Emslie, 1970). We selected a mafic dyke from the Ohře Rift as the starting liquid composition (olivine melanephelinite
from the Spojil Dyke; Vaněčková et al., 1993), adopted to fit the composition of early crystallizing olivine. We added 37 ppm Ni
(þ92%) but halved the concentration of CaO (–50%) to match the starting composition with the composition of the most primitive
olivine crystallized. The contents of relevant elements or oxides in the starting liquid are 1035 wt % FeOt, 1678 wt % MgO, 504 wt
% CaO and 440 ppm Ni. The composition of the first olivine crystallizing in our model is marked by a blue star. The evolutionary
stages [1, 2 and 3 in (a) and (b)], as discussed in the text, should be noted. Literature data include OPF (Oberpfalz) peridotites from
Ackerman et al. (2013) and various types of olivine recovered from the Mýtina Maar and Železná hůrka (Geissler, 2005). Green
shaded field indicates the range of mantle olivine.
Cores of clinopyroxenes in cumulates are augitic and
have high Mg# (82–90), low TiO2 (<18 wt %) and Al contents (<033 a.p.f.u.; Fig. 5a) but high Cr# (up to 18) and
high Na2O (>07 wt %; Fig. 5b). However, green-core
clinopyroxenes, as reported from the Eifel (Duda &
Schmincke, 1985) and basanites from Slovakia (Dobosi
& Fodor, 1992) have not been observed. Clinopyroxene
overgrowth rims as well as microphenocrysts in the
matrix have a greater wollastonite and ferrosilite component (Fig. 5c). Moreover, overgrowth rims and
phenocrysts form trends towards compositions contrasting with those of clinopyroxene cores in cumulates
(lower Na2O and Mg#, higher TiO2 and wollastonite
contents). Aluminium substitutes for Si, leading to a
negative correlation of Al (a.p.f.u.) and Si contents. The
fine laths of clinopyroxene (of a few hundred
micrometres) show rhythmic zonation on backscattered
electron images. However, a clear evolution towards a
defined mineral composition is not observed and the
mineral composition ranges from augite with variable
contents of Ti to diopside with up to 57 wt % TiO2. In
terms of trace elements, cores show normal convex
rare earth element (REE) patterns with the bulge centred at the middle REE (MREE; Supplementary Data
(SD) Fig. S1a), similar to mantle clinopyroxenes in
xenoliths from the Oberpfalz (Ackerman et al., 2013).
Clinopyroxene grains in the clinopyroxene–phlogopite
cumulate of sample V1 show a simple zonation from
core (augite: Wo43En48Fs9) to rim (diopside:
Wo50En42Fs8) but diopside overgrowth rims and the
spongy contact zones between clinopyroxene and
phlogopite show a less weak bulge mainly owing to
1754
Journal of Petrology, 2015, Vol. 56, No. 9
0.6
(a)
Al [a.p.f.u.]
0.5
0.4
0.3
0.2
Cores
Zonation
Rims
Matrix
Cumulates
0.1
0.0
65
70
75
80
85
90
95
Mg#
1.2
(b)
Na2O [wt. %]
1.0
0.8
0.6
cumulates
(main group)
0.4
and U3; Fig. 3b). These latter phlogopites are relatively
small (less than 300–400 mm) and slightly more magnesian (Mg# 87–88) than the large phlogopites in cumulates
(Mg# 83–84). In terms of BaO and TiO2 concentrations,
phlogopites in the Železná hůrka lavas are relatively
primitive, similar to mantle phlogopites and intermediate between those found in lamproites and carbonatites
(SD Fig. S2; Dunworth & Wilson, 1998).
Additional accessory minerals include haüyne and
opaque crystals of the spinel group. These spinels
sensu lato are common and occur predominantly as inclusions in olivine or in direct contact with olivine. They
are generally Fe- and Ti-poor and display variable Cr
and Al concentrations, resulting in variable Cr#, ranging
from 37 to 58. Only very few spinels (often in contact
with clinopyroxenes and lower in Mg#) have Fe- and Tirich compositions and are solid solutions of Ti-magnetite. However, the majority of spinels belong to the
group of Mg–Al-chromites (solid solutions of spinel
sensu stricto, hercynite, chromite and Mg–Al-chromite).
Haüyne is present as small (<100 mm), idomorphic crystals,
typically
associated
with
clinopyroxene
phenocrysts.
Crustal xenoliths
0.2
0.0
65
70
75
80
85
90
95
Mg#
(c)
40
30
20
0
50
Wo
ll
as
ton
ite
60
10
Ferrosillite
Diopside
Hedenbergite
40
Augite
30
60
70
80
90
0
10
Enstatite
Fig. 5. (a) Aluminium (a.p.f.u.) and (b) Na2O (wt %) vs Mg# in
clinopyroxene. A noteworthy feature is the difference between
cumulates (Ti-augite) and diopsidic phenocrysts and overgrowth rims, also obvious in the proportional changes in mineral components (c).
enrichment of the light REE (LREE) relative to MREE and
heavy REE (HREE; SD Fig. S1b). Spongy reaction textures between Ti-augite and diopside (e.g. Fig. 3d) show
a chemical composition intermediate between the two
end-members (SD Fig. S1).
Accessory phases
Phlogopite is the main accessory mineral with an overall volume of the order of 1–2%; single crystals reach
sizes of up to 10 mm. Phlogopite crystals intergrown
with clinopyroxene (cumulates) are less rounded than
those that occur as inclusions in olivines (samples V2
Microscopically, we distinguish two xenolith groups: a
felsic variety containing quartz and feldspar and a
brownish, porous variety with an undefined mineral assemblage. In addition to quartz and potassic feldspar,
the felsic xenoliths (Fig. 3c) also contain albite and muscovite. Modal mineral contents cover a broad range
from almost pure quartz to almost pure potassic feldspar. These felsic xenoliths are interpreted as fragmented parts of the regional host-rocks, which are
dominated by phyllites, quartzites and mica shists
(Geissler et al., 2007). Some of these xenolithic rock
fragments seem to have disintegrated almost completely leading to isolated, subrounded quartz crystals
in the volcanic matrix. In sample V1, the assimilation–
melting reaction between a quartz crystal and a clinopyroxene–phlogopite cumulate has been preserved in
the form of an interstitial, mingled glass (Fig. 3d). The
mineral assemblage of the brownish xenoliths could
not be resolved with certainty but probably includes
amphibole and may represent altered parts of the phyllitic host-rocks.
Geochemistry of volcanic rocks and glasses
Major element composition
The compositions of the Železná hůrka samples are
relatively uniform (Fig. 6). However, the major element
contents of glasses from the lower and the upper tephra
differ significantly from those of the whole-rock samples. In particular, the glasses have much lower MgO
contents at a given SiO2 compared with the whole-rock
samples. Furthermore, the FeOt (Fig. 6c) contents of the
glasses are slightly lower, whereas the concentrations
of the other elements are higher than those of the
Journal of Petrology, 2015, Vol. 56, No. 9
1755
Fig. 6. Variation of MgO vs (a) TiO2, (b) Al2O3, (c) FeOt, (d) CaO, (e) Na2O and (f) K2O (all in wt %) ZH, Železná hůrka samples, from
this study and Ulrych et al. (2013). Data sources: central Ohře Rift, rift shoulder, Oberpfalz and Komornı́ hůrka from Ulrych et al.
(2013) and Haase & Renno (2008) and references therein; Mýtina Maar from Geissler et al. (2007); Ohře Rift melilitites from Ulrych
et al. (2008).
whole-rock samples (Fig. 6). Whole-rock samples straddle the boundary between basanites and foidites in a
volatile-free total alkalis (Na2O þ K2O) versus silica
(SiO2) diagram (Fig. 7a, TAS diagram; Le Bas &
Streckeisen, 1991). However, the volcanic glasses
show significantly higher concentrations of alkalis
(>8 wt %) than the whole-rock samples. Thus, glasses
from Železná hůrka can be classified as strongly
SiO2-undersaturated foidites. Melilite or nepheline have
not been observed in thin section and based on a
SiO2 þ Al2O3 (54–55 wt %) versus CaO þ Na2O þ K2O
(230–235 wt %) discriminant diagram (Le Bas, 1989) we
conclude that the lavas are part of the nephelinite
rock series, even though glass samples plot on the
boundary between the nephelinite and melilitite rock
series (Fig. 7b). We further note that the glasses contain
1756
Journal of Petrology, 2015, Vol. 56, No. 9
samples are well within the range reported for the Ohře
Rift (e.g. Haase & Renno, 2008; Ulrych et al., 2013) but
with relatively high Rb, Ba, Th, U, Nb and Ta and low
MREE to HREE resulting in an overall slightly steeper
slope in a multielement diagram (Fig. 8a) compared
with most other samples from the Ohře Rift. In contrast,
the trace element compositions of glasses extend the
range of published Ohře Rift data (Fig. 8a). Slight but
significant differences in trace element composition
exist between the glasses and whole-rocks (e.g. more
pronounced negative anomalies of Zr and Hf), resulting
in a strong contrast in some trace element ratios, such
as Nb/Zr (Fig. 8b). Generally, the whole-rock and glass
samples from Železná hůrka have highly enriched trace
element ratios relative to the primitive mantle of
Lyubetskaya & Korenaga (2007); for example, Nb/Zr is
6–13, Ta/Hf is 8–18, La/Yb is 29–36 (Fig. 8c) and Nb/
U is 15–20 higher relative to the respective ratios of
the primitive mantle.
Isotope geochemistry
Fig. 7. (a) Volatile-free total alkalis, Na2O þ K2O (wt %) vs SiO2
(wt %) after Le Bas & Streckeisen (1991). Whole-rock samples
from Železná hůrka overlap with the range of basanitic lavas
reported from the Oberpfalz, but glasses are much higher in
their alkali content at a given SiO2 content. (b) Combined vola(wt
%)
vs
tile-free
oxide
diagram,
SiO2 þ Al2O3
CaO þ Na2O þ K2O (wt %), for the discrimination between basanite, nephelinite and melilitite after Le Bas (1989). Data sources as in Fig. 6.
high concentrations of volatile elements such as S and
Cl (Table 1). The Cl/Nb ratios of about 17–20 are higher
than in mid-ocean ridge basalt (MORB) and ocean island basalt (OIB) but similar to those of other continental rift regions (e.g. Rowe et al., 2015).
In general, samples from the Ohře Rift and the CEVP
span a large range in their Nd–Sr isotopic composition,
with the Ohře Rift samples showing higher 87Sr/86Sr
at a given 143Nd/144Nd relative to other CEVP samples
(Fig. 9a). Whole-rock samples from Železná hůrka are
relatively homogeneous in their Sr–Nd–Hf isotope compositions but slight differences are evident in Sr–Nd isotope space amongst glass samples from the lower
tephra unit (Fig. 9a). Their compositions overlap in eNd–
eHf (Fig. 9b) and fall within the broader mantle array of
Vervoort et al. (1999), consistent with data from the
CEVP. Two glass samples have significantly lower
143
Nd/144Nd and higher 87Sr/86Sr compared with other
Quaternary volcanic rocks from the Ohře Rift (e.g.
Haase & Renno, 2008; Ulrych et al., 2013).
The O isotope compositions of glasses are between
d18O þ54 and þ58% V-SMOW, with the exception of
sample L1, which has a value of þ62% (þ613 and
þ625%; one duplicate analysis). The high d18O of sample L1 is associated with higher MgO, CaO, Cl/K, Ce/Pb
and 87Sr/86Sr, but lower Nb/Zr, La/Sm, K/Ti and
177
Hf/176Hf values.
DISCUSSION
Trace element geochemistry
Insights from mineral phases
Olivine antecrysts
The whole-rock samples show primitive mantle-normalized (Lyubetskaya & Korenaga, 2007) incompatible
element patterns that are enriched in highly incompatible trace elements and the LREE relative to HREE [e.g.
(La/Yb)N ¼ 285–335, (Gd/Yb)N ¼ 40–41 and (La/
Sm)N ¼ 42–50] (Fig. 8a)]. The most prominent anomaly
in the multi-element pattern is a negative Pb anomaly
and there are slightly negative K and Ti anomalies. The
HFSE Nb and Ta are slightly enriched (Fig. 8a) whereas
Zr and Hf are slightly depleted. Overall, the whole-rock
The compositions of olivine crystals reveal insights into
the plumbing system of Železná hůrka. Cores of olivine
crystals show normal mineral zonation (i.e. decrease in
forsterite content towards the rim) and plot within the
field of olivines in equilibrium with melts from ‘ordinary’ mantle peridotite rather than melts from pyroxenitic lithologies (Straub et al., 2011). However, olivine
cores are distinct from primary mantle olivines found in
mantle peridotite xenoliths of the same region (Fig. 4c;
lower Fo and Ni, higher CaO). Primary mantle olivines
Journal of Petrology, 2015, Vol. 56, No. 9
1757
Fig. 8. (a) Multi-element plot for whole-rock and glass samples normalized to the primitive mantle values of Lyubetskaya &
Korenaga (2007). Red, glass data; orange, whole-rock data. The grey field corresponds to literature data with two typical patterns
shown as grey lines. (b) Nb/Zr vs MgO (wt %) and (c) chondrite-normalized (Palme & O’Neill, 2003) (La/Sm)N vs La (ppm). Data sources as in Fig. 6.
usually have CaO contents of less than 01 wt % and
characteristic Ni concentrations of 2600–3200 ppm,
whereas magmatic olivines have CaO concentrations
>018 wt % (Stamper et al., 2014). Olivines entrained in
the Železná hůrka lavas are thus of magmatic origin.
However, they are in disequilibrium with the host lava,
which is evident both petrographically (common dissolution textures) and chemically (Mg# of melt much lower
than expected if assuming olivine–melt equilibrium).
These olivines are thus antecrysts and their evolution can be differentiated into three generations:
(re-)crystallization (stage 1) in equilibrium with mantle
peridotite (Fo > 89, Ni > 1700 ppm; Straub et al., 2011),
as demonstrated by the modelled crystallization trend
shown in Fig. 4 (see figure caption for model details)
and SD Table S2. Minor differences between the calculated fractionation trend and observed mineral
compositions during stage 1 (Fig. 4a) may be explained
by minor changes in the olivine–melt partition coefficient for Ni as a result of changes in liquid composition
(e.g. concomitant crystallization of magnetite), pressure
and temperature (e.g. Matzen et al., 2013). Further crystallization of olivine along the predicted crystallization
path from forsterite contents of 89 to 84 (Stage 2 in
Fig. 4a) is accompanied by a drop in Ni content from
>2000 ppm to <1000 ppm (Fig. 4a) but a minor increase
in CaO (Fig. 4c). In Stage 3, groundmass crystals and
rims of olivine crystals show a significant drop in Ni
at constant forsterite content (500 ppm Ni and below;
Fig. 4a) but a strong enrichment in CaO (up to >10 wt
%; Fig. 4c) and MnO. These low-Ni–high-CaO rims of
olivine antecrysts have not been reported from Železná
hůrka in previous studies (e.g. Geissler, 2005). A very
similar pattern in mineral evolution (although at a
1758
Journal of Petrology, 2015, Vol. 56, No. 9
Fig. 9. (a) 87Sr/86Sr vs 143Nd/144Nd for volcanic rocks from the CEVP (Hocheifel: Jung et al., 2006; Röhn: Jung et al., 2013;
Vogelsberg: Jung & Masberg, 1998) compared with data from the Ohře Rift region (Ohře Rift: Haase & Renno, 2008, and references
therein; Ulrych et al., 2013; Ohře Rift melilitites: Ulrych et al., 2008; Oberpfalz mantle xenoliths: Ackerman et al., 2013; Komornı́
hurka: Haase & Renno, 2008; Železná hůrka glasses and whole-rocks: this study). Also shown are the approximate positions of
mantle endmembers PREMA, DMM, EM1 and EM2 (Stracke, 2012), LVC (Hoernle et al., 1995) and EAR (Cebriá & Wilson, 1995; as
defined by Lustrino & Wilson, 2007). (b) (eNd)i vs (eHf)i of volcanic rocks from the CEVP [Vogelsberg (17 Ma), Rhön (24 Ma),
Hocheifel (40 Ma): Jung & Masberg, 1998; Jung et al., 2011; Pfänder et al., 2012], Železná hůrka glasses and whole-rocks (this study)
and the Udachnaya East kimberlite (Kamenetsky et al., 2009b). Fields for MORB, HIMU, EM1 and EM2 (after Pfänder et al., 2007) are
shown for comparison. The mantle array is after Vervoort et al. (1999; eHf ¼ 133 eNd þ 319) and data for CHUR are from Bouvier
et al. (2008). Decay constants for 147Sm and 176Lu are from Begemann et al. (2001) and Scherer et al. (2001), respectively.
smaller magnitude and at overall more primitive compositions) is recorded in fresh olivines of the
Udachnaya East kimberlite in Yakutia (Kamenetsky
et al., 2008). Furthermore, the composition of the olivine
antecrysts cannot be explained by crystallization from a
single parental liquid, but instead requires at least one
other liquid with a very different composition (e.g. high
Ca and Mn, low Ni).
Compositional profiles from the rim of olivine crystals towards the core (SD Fig. S3) allow tracking of the
different steps of crystal evolution (e.g. composition of
parental melt, solid-state crystal diffusion, presence or
Journal of Petrology, 2015, Vol. 56, No. 9
absence of crystal–melt equilibrium). Diffusion of calcium in olivine is significantly slower than that of Fe or
Mg, whereas Ni shows similar diffusion rates to Fe and
Mg (e.g. Jurewicz & Watson, 1988; Petry et al., 2004).
The forsterite content changes in a narrow zone between roughly 50–100 mm from the rim from about 90 in
the core to 860–865 at the rim, but the overgrowth rim
itself displays a constant forsterite content. The Ni concentration, in contrast, shows the most significant decrease, concurrent with the change in forsterite, but
also decreases across the overgrowth rim where the
forsterite content is constant (SD Fig. S3a and b). The
Ca profile (SD Fig. S3c) shows a different pattern, with
almost constant concentrations over a wide range, even
over most of the overgrowth rim and the zone of forsterite change. We therefore conclude that the high-Fo
olivine cores represent olivines crystallized in the (lithospheric) mantle, whereas their rims were diffusively
modified after entrainment in the parental melt of the
Železná hůrka nephelinite.
To constrain further the composition of the parental
magma, we empirically calculated the (theoretical) partition coefficient of Ca between olivine and hostmagma, assuming that a diffusion profile reflects an
attempt to reach a chemical equilibirium at the crystal–
melt interface. Applying the method of Libourel (1999),
we see that the partitioning for CaO between olivine
and melt, if calculated based on olivine composition
only [equation 14 of Libourel (1999)], is much lower
than if calculated from the molar fraction of CaO in
olivine and a*meltCaO [equations 12 and 13 of Libourel
(1999); note that this partition coefficient is of a hypothetical nature as olivines are in fact in disequilibrium
with the host lava]. This implies that, to explain the
high concentration of CaO in the olivine rims, a liquid
is required that contained much higher concentrations
of CaO. Based on experimental data (e.g. Libourel,
1999, fig. 5), we can estimate a minimum concentration of 20 wt % CaO in the parental liquid to explain
the high enrichment of Ca (5000 ppm or more; SD Fig.
S3c) in the olivine rims at intermediate alkali contents
(4–8 wt %). It is important to note that the total alkali
content (Na2O þ K2O) has a much stronger influence
on the partitioning of Ca into olivine than oxygen fugacity or temperature (Jurewicz & Watson, 1988;
Libourel, 1999).
Clinopyroxene cumulates and phenocrysts
Clinopyroxenes can be separated into an augitic cumulate domain (cores) and a more evolved and LREE-enriched phenocryst (and overgrowth rim) domain
(diopside). These clinopyroxene overgrowth rims and
phenocrysts formed at a later stage of magmatic differentiation compared with the olivines, as indicated by (1)
the lower Mg# (68–84) compared with olivines (86–905)
and (2) the growth of clinopyroxene phenocrysts at the
margins of pre-existing olivines (e.g. Fig. 3e). These
clinopyroxenes show an increase in wollastonite and
1759
ferrosilite components with increasing differentiation
(Fig. 5c), implying that Mg–Fe exchange follows a normal differentiation trend whereas the Ca contents do
not. Consistent with our previous observation of Ca enrichment in olivine overgrowth rims, the Ca-rich composition of clinopyroxene overgrowth rims and
phenocrysts may be seen as another hint of a highly
Ca-enriched parental melt. The Ti enrichment in clinopyroxene is equal to the enrichment observed in SW
German melilitites by Dunworth & Wilson (1998) and
has been considered to be close to the limits of Ti4þ
substitution.
Clinopyroxenes are an important carrier for trace
elements and their distinctive trace element patterns
allow us to infer magmatic processes during crystallization. We previously described the similarity in REE patterns between clinopyroxenes in the lavas of Železná
hůrka and mantle xenoliths of the Oberpfalz (SD Fig.
S1). Even more interesting are their similar trace element patterns compared with clinopyroxenes hosted as
inclusions in fresh olivines from the Udachnaya East
kimberlite (Fig. 10). Both clinopyroxene domains show
a relative enrichment in the LREE to MREE and a negative Zr anomaly. However, there are also some differences compared with the Udachnaya clinopyroxenes,
which are more enriched especially in the LREE but
more depleted in the HREE relative to Železná hůrka
(Fig. 10). This HREE depletion could result from the preferential partitioning of these elements into garnet at
pressures >25 GPa (in the garnet stability field), which
is also supported by their high Na and Cr concentrations, indicating even higher pressures of >45 GPa
(Kamenetsky et al., 2009a). The most prominent contrast, however, is the strong negative Ti anomaly in the
kimberlitic clinopyroxene inclusions, which could result
from direct substitution of Ti for Si in the olivine host
(Hermann et al., 2005) or the preferential partitioning of
Ti into ilmenite, present in the kimberlite but absent in
Železná hůrka lavas.
Accessory phases
Železná hůrka phlogopites probably formed in equilibrium with mantle peridotite, reflected in their high Mg#,
low Ba and intermediate Ti concentrations. The phlogopite inclusions in olivine may have formed by reaction
of pre-existing melt (in inclusions in olivine) during the
entrainment of olivines in a Ba-enriched residual melt
(Seifert & Kämpf, 1994) or by reaction of olivine with
CO2 to form phlogopite and carbonate melt (Mysen &
Virgo, 1980). In contrast, phlogopite in cumulates has
slightly lower Mg# and higher Ba and Ti contents. The
overall enrichement in these elements may be an effect
of shallow magmatic processes (Seifert & Kämpf, 1994).
All phlogopites from Železná hůrka are intermediate
between the distinct melilitite and leucite–nephelinite
trends of Dunworth & Wilson (1998) but similar
to phlogopite in the mantle and kimberlites (see SD
Fig. S2).
1760
Journal of Petrology, 2015, Vol. 56, No. 9
Cpx / Primitive Mantle
100
Cpx p
henoc
r. & rim
s
10
1
Cpx cores
0.1
Kimberlite cpx inclusions
(Udachnaya East)
0.01
0.001
Rb
Th
Ba
Nb
U
La
Ta
Pb
Ce
Sr
Pr
Zr
Nd
Hf
Sm Ti
Tb
Y
Er
Yb
Eu Gd Dy Ho Tm Lu
Fig. 10. Primitive mantle-normalized [values of Lyubetskaya & Korenaga (2007)] trace element patterns of clinopyroxenes from
Železná hůrka (cores of cumulates, and phenocrysts and overgrowth rims) and inclusions hosted in fresh olivines of the
Udachnaya East kimberlite (Kamenetsky et al., 2009a). The similarity in the trace element patterns should be noted. However, the
Udachnaya East kimberlite shows a more pronounced difference between the light (more enriched) and heavy (depleted relative to
the primitive mantle) REE.
Spinel sensu lato shows a broad negative correlation
between Mg# and stoichiometrically calculated Fe3þ
content, indicating a trend towards more oxidizing conditions. The corresponding mineralogical change from
Mg–Al-chromite towards Ti-magnetite is commonly
observed in carbonate-rich magmas (e.g. Jones &
Wyllie, 1985). Further evidence for highly oxidizing conditions in the melt is the presence of haüyne, a sodalitegroup mineral that contains oxidized sulphur as a major
phase (5 wt % S). The association of titanomagnetite
and haüyne with clinopyroxene phenocrysts indicates
that high oxygen fugacities were reached late during
magmatic differentiation.
Crustal assimilation
Physical evidence for interaction between the ascending
magmas and the continental crust is provided by the
presence of quartz crystals in the matrix and crustal
xenoliths. This interaction has a high potential to alter
the chemical composition of the erupted lavas and thus
to obscure the nature of the parental melt. A spectacular
snapshot of crustal assimilation is preserved in sample
V1 (Fig. 3d). Here, the composition of interstitial glass
that formed at the contact of a (crustal) quartz crystal and
a clinopyroxene–phlogopite-cumulate shows a strong
chemical contrast to the otherwise homogeneous matrix
glass. The interstitial glass is mingled (Fig. 3d) and very
SiO2-rich (c. 69 wt %; Table 2) as opposed to the strongly
silica-undersaturated (40 wt % SiO2) matrix glass
(Table 1). With respect to the low solidus temperature of
quartz in the presence of phlogopite and H2O–CO2 vapour (700–800 C at crustal depths; e.g. Wones &
Dodge, 1977; Bohlen et al., 1983) and the presence of
numerous dispersed single quartz crystals (likely to represent disintegrated crustal xenoliths) in the lavas, assimilation of crustal material may play an important role
in the petrogenesis of the Železná hůrka lavas. However,
because sample V1 was taken from a blocky lava, in
which temperatures may remain at higher levels for a
longer period of time relative to the explosively erupted
tephra, we need to further constrain the possible role of
crustal assimilation using geochemistry.
Assimilation of crustal material may effectively
change elemental concentrations and ratios as well as
isotope ratios [e.g. decreasing Ce/Pb or Nb/U in conjunction with increasing SiO2, 87Sr/86Sr or d18O of
whole-rock or glass (subsequently noted as d18OWR; e.g.
Taylor, 1980; Jung & Hoernes, 2000; Jung et al., 2013)]
and has been found to play a key role in the magmatic
evolution of several suites of the CEVP. Lavas from the
Rhön, for example, show a broad positive correlation
between d18OWR and SiO2 (Fig. 11a), interpreted as the
result of combined assimilation and fractional crystallization processes (Jung et al., 2013). Similarly, one of
our new samples from Železná hůrka shows higher
d18OWR, associated with higher 87Sr/86Sr (07039),
whereas all other samples show mantle-like d18OWR
[mantle range of Taylor (1980) and Eiler et al. (2000)].
Quantification of the influence of continental crustal
assimilation in the petrogenesis of the Železná hůrka
lavas is difficult, in particular with respect to their isotope characteristics. Further insights can be obtained
from the major element compositions of the crustal
xenocrysts and the general trace element characteristics of the continental crust. We identified quartz,
K-feldspar, albite and muscovite as the main phases in
the crustal xenoliths; amphibole may also be present.
Journal of Petrology, 2015, Vol. 56, No. 9
1761
Fig. 11. Stable oxygen isotope composition in per mil relative to V-SMOW. Grey shaded bands indicate the d18O range (in per mil
relative to V-SMOW) of normal (N)-MORB glasses (Eiler et al., 2000) and the range of d18OWR for primitive melts containing 4–5 wt
% Na2O (Eiler, 2001). (a) d18O of Železná hůrka (ZH) glasses and whole-rocks (Rhön: Jung et al., 2013; Garrotxa, NE Spain: Cébria
et al., 2000) vs SiO2 (wt %). Most of the Železná hůrka glasses (except for L1) plot in the d18OWR range of N-MORB (Eiler et al., 2000).
Rhön samples have been affected by assimilation of crustal material and subsequent fractional crystallization (Jung et al., 2013);
this is also visible in the broad correlation between d18OWR and SiO2 (trend encompassed by the dashed lines). (b) d18OWR vs CaO/
Al2O3 of whole-rocks from the Rhön (Jung et al., 2013), Garrotxa (NE Spain; Cébria et al., 2000) and Železná hůrka glasses. The average compositions of upper (UCC) and lower continental crust (LCC) are shown for comparison (Rudnick & Gao, 2003). Rhön lavas
are consistent with assimilation of continental crust whereas glasses from Železná hůrka show a positive correlation between O isotope composition and CaO/Al2O3.
Figure 12 shows the chemical composition of phenocrysts [clinopyroxene, phlogopite, haüyne and potentially amphibole; found in cumulates by Geissler et al.
(2007)], along with olivine antecrysts and crustal xenocrysts (quartz, potassic feldspar, albite, muscovite).
The first phase of melting in metapelitic crustal xenoliths (such as phyllite) involves albite-rich plagioclase
(oligoclase), controlled by the H2O and alkali release
during the breakdown of muscovite (Grapes, 1986).
This pattern is well reflected by the trend of the Ohře
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Journal of Petrology, 2015, Vol. 56, No. 9
Fig. 12. Major oxide composition of Ohře Rift lavas (data sources as in Fig. 6; this study) compared with single magmatic and
xenocrystic mineral phases and experimental melt compositions for KLB-1 [dry peridotite of Hirose & Kushiro (1993); run details in
Fig. 15], PERC-3 [carbonated peridotite of Dasgupta et al. (2007); run details in Fig. 15] and KC2 [carbonatite melt of Sweeney
(1994)]. (a) CaO and (b) Na2O vs SiO2 and (c) K2O vs Al2O3 (all in wt %). The compositions of the Quaternary volcanic rocks of the
Ohře Rift are controlled by variable proportions of the primary magmatic minerals [Cpx, clinopyroxene; Hbl, hornblende (Geissler
et al., 2007); Hy, haüyne; Phl, phlogopite; Ol, olivine] rather than by assimilation of crustal components (Ab, albite; Kfsp, K-feldspar
Ms, muscovite; Qz, quartz). It should be noted that fractional crystallization would result in trajectories away from the crystallizing
mineral compositions, whereas assimilation trajectories (grey arrows) would point towards the assimilated mineral. Albite-rich
plagioclase and potassic feldspar exhibit a range of compositions (grey and brown crosses, respectively).
Journal of Petrology, 2015, Vol. 56, No. 9
1763
120
47±10
Oh e (Eger) Rift
Oh e (Eger) Rift: melilitites
Oberpfalz
Komorní h rka
M tina Maar
ZH whole rocks
ZH matrix glasses
SWG melilitites
100
Ce/Pb
80
60
40
25±5
20
UCC
Assimilation of upper continental crust
0
0
20
40
60
80
100
120
Nb/U
Fig. 13. Nb/U vs Ce/Pb for Ohře Rift lavas (data sources as in
Fig. 6; this study) and the SW German (SWG) melilitites of
Hegner et al. (1995) and Dunworth & Wilson (1998). All the
Quaternary lavas studied here fall within the oceanic array for
Nb/U (47 6 10; Hofmann et al., 1986) but are more enriched in
Ce/Pb. The average upper continental crust (UCC) has Nb/U
and Ce/Pb of 44 and 37, respectively (Rudnick & Gao, 2003).
Assimilation trajectories are indicated.
Rift lavas in Fig. 12, extending from a parental magma
composition, probably similar to the PERC-3 composition of Dasgupta et al. (2007), towards variable mixtures of albite and muscovite (plus K-feldspar).
However, the composition of the Železná hůrka lavas is
controlled by the accumulation of olivine (whole-rock)
and fractional crystallization of clinopyroxene (glasses)
as evident from the trajectories between whole-rocks,
glasses, olivine antecrysts, clinopyroxene and other
magmatic phases [haüyne as phenocryst and phlogopite (þ hornblende; Geissler et al., 2007) as cumulate
phases; Fig. 12]. All these samples consistently fall at
the low-silica end of the range of most Ohře Rift lavas,
with only melilitites (Ulrych et al., 2008) being more
undersaturated. The predominant crustal assimilation
of albite and mica (6 K-feldspar) could also explain the
presence of residual single quartz crystals. The trends
in Fig. 12 point to a parental melt composition of the
Železná hůrka lavas with even lower SiO2 and Al2O3,
but higher CaO and total alkalis, if we assume that the
highly silica-undersaturated nephelinites of Železná
hůrka have already assimilated significant amounts of
silica-rich crustal lithologies. To produce such parental
melt compositions from a peridotite source, significant
amounts of carbonate (i.e. CO2) must be involved (see
PERC-3 and KC2 in Fig. 12 and discussion below).
In contrast to major elements, where large amounts
of assimilated material are necessary to significantly
change the whole-rock chemical composition, trace
elements (and their ratios) are more sensitive to assimilation. Trace element ratios such as Nb/U or Ce/Pb have
proven to be powerful tracers, with respective values of
47 6 10 and 25 6 5 in oceanic basalts (Hofmann et al.,
1986) and 4.4 and 3.7 in the continental crust (Rudnick &
Gao, 2003; Fig. 13). Železná hůrka lavas have Nb/U
similar to the range observed in oceanic basalts but
with Ce/Pb one to four times higher (Fig. 13).
Furthermore, O isotopes show a narrow range between
54 and 63% V-SMOW and are positively correlated
with CaO/Al2O3 in the glasses (Fig. 11b), contrary to
what is expected for the assimilation of upper continental (granitoid) crustal material as observed in the Rhön
province [Fig. 11b; tephrites, phonolites and trachytes
of Jung et al. (2013)].
Precise information on the composition of the parental melt (which is likely to have higher Ce/Pb) remains
obscure and thus assimilation of crustal material can be
neither excluded nor confirmed and quantified with certainty. However, the Železná hůrka lavas have much
higher trace element abundances than average upper
continental crust (Rudnick & Gao, 2003). The assimilation of crustal material would thus lead to a relative depletion of trace elements in the melt. Assuming a
maximum of 10 vol. % of upper continental crust being
assimilated in the lavas erupted would not change their
overall incompatible trace element patterns significantly. It is not sufficient to explain the high enrichment
in Nb, Ta, LREE to MREE and Sr along with a relative depletion in Pb, Zr and Hf (Fig. 8a). These trace element
characteristics are thus assumed to be of primary magmatic origin (melting and/or source). Before we further
constrain these, we will first try to reconstruct the
plumbing system of Železná hůrka.
Constraints on the magma plumbing system
Thermobarometry
The composition of clinopyroxene phenocrysts can be
used to determine the thermobarometric conditions of
crystallization using equations 32c, 33 and 34 of Putirka
(2008; Excel spreadheets are available for download
from the website of K. Putirka: http://www.fresnostate.
edu/csm/ees/faculty-staff/putirka.html). Successful P–T
estimates were assumed if (1) the calculated Fe–Mg exchange coefficient is within the range of experimental
observations (028 6 008; Putirka, 2008) and (2) the
clinopyroxene components [diopside–hedenbergite
and enstatite–ferrosilite, calculated using the normative
procedure of Putirka et al. (2003)] from measured mineral compositions are in agreement (6 001) with the expected clinopyroxene components calculated from the
host-rock composition. This resulted in 38 successful P–
T estimates (Fig. 14; SD Table S3) that fall into two
groups. The first group consists of Na- and Cr-rich augites that occur as cores of cumulate crystals (Fig. 3d)
and indicate pressures of around 10–13 GPa at temperatures of 1250–1300 C (Fig. 14). The second clinopyroxene group (diopsidic rims) displays lower pressures
and temperatures at around 08 GPa and 1150–1200 C,
respectively. These two groups are distinct from each
other even given the method’s relatively large uncertainty of 6015 GPa and 650 C (Putirka, 2008). We conclude that cumulates formed in the upper SCLM,
whereas phenocrysts and overgrowth rims crystallized
1764
Journal of Petrology, 2015, Vol. 56, No. 9
0
0
Focal depth
of recent
earthquake
swarms
1SD
CC
Eifel
magma
chamber
Xenolith
re-equilibrium
20
MOHO
SCLM
r me
majo
50
lting
1000
ting
2.0
800
mel
Cumulates
Xenoliths
t
pien
Geissler et al. (2007)
t
sal
ba m
ali
r
alk othe
ge
Cores
Zonation
Rims
Matrix
Cumulates
40
inci
1.5
R
en e-h
tra eat
in ing
m
en /
t?
30
1.0
Depth [km]
Pressure [GPa]
0.5
10
nephelinites; e.g. Duda & Schmincke, 1985; Jung et al.,
2006) and is controlled by a combination of intraplate
magmatic activity and lithospheric extension (propagation of pre-rift volcanism linked to the Upper Rhine
Graben; Fekiacova et al., 2007). Depths of crystallization
are similar to our estimates for the nephelinitic rocks of
Železná hůrka with respect to crystallization below the
Moho and within the lower crust (e.g. Duda &
Schmincke, 1985). However, Sachs & Hansteen (2000)
calculated the depth of a possible magma chamber at
06–07 GPa, which is at the low end of the pressures
calculated for the crystallization of clinopyroxene in the
Železná hůrka lavas. The range of pressures recorded
by clinopyroxene overgrowth rims and phenocrysts at
Železná hůrka (07–10 GPa; Fig. 14) argues against the
presence of a lower crustal magma chamber underneath this volcano but instead for continuous crystallization during melt ascent.
60
Constraints on the parental melt composition
1200
1400
Temperature [°C]
Fig. 14. Thermobarometric estimates for clinopyroxene crystallization using the method of Putirka (2008). (For full details see
the main text and SD Table S2.) The onset of clinopyroxene
crystallization is at about 12 GPa in the SCLM. However, most
clinopyroxenes indicate pressures of crystallization between
07 and 10 GPa, close to the Moho (Babuška & Plomerová,
2010) or within the continental crust. Focal depths of recent
earthquake swarms (e.g. Horálek et al., 2000; http://www.ig.cas.cz/en/structure/observatories/west-bohemia-seismic-network-webnet) and the inferred depth of a crustal magma
chamber beneath the Eifel (Sachs & Hansteen, 2000) are
shown for comparison. Geotherms, melting paths and additional P–T data are from Geissler et al. (2007).
in the lower continental crust (Fig. 14; Moho <28 km;
Babuška & Plomerová, 2010).
The temperature estimates of Geissler et al. (2007)
yield lower values for their xenolith suite (including
hornblendite, pyroxenite, wehrlite) compared with our
cpx–melt temperatures at similar pressures (Fig. 14).
Geissler et al. (2007), however, speculated also about
the possible intrusive origin of these rocks or existing
mineral disequilibria in these rocks. In contrast, a distinct suite of cumulates (e.g. ol–cpx–spl cumulates)
studied by Geissler et al. (2007) (red open diamonds in
Fig. 14) points towards the magmatic temperatures recorded in our samples. This could possibly indicate
later reheating of those cumulates after their entrainment in the Železná hůrka host nephelinite (our data);
further studies are required to resolve the complex thermobarometric history.
Volcanism in the West-Eifel and Hocheifel regions of
Germany shows close similarities to the volcanism
in the Ohře Rift region (e.g. SiO2-undersaturated, highalkali volcanic rocks such as foidites, basanites and
So far, we have demonstrated that there is strong mineralogical evidence (in terms of mineral assemblage,
composition and evolution) for a genetic link between
the Quaternary nephelinites of Železná hůrka and other
silica-undersaturated magmas, such as melilitites and
even kimberlites. Olivine antecrysts show very similar
chemical evolution, clinopyroxenes have similar trace
element patterns, phlogopites are intermediate between those found in carbonatites and lamproites, and
haüyne and titanomagnetite indicate a shallow depth of
oxidation of the melt (e.g. during magma ascent).
Further evidence for the major role of carbonate in the
genesis of the parental magmas comes from the major
element composition of these rocks, which makes them
less sensitive to secondary processes such as crustal
assimilation.
We use a simple Rayleigh fractionation model
involving olivine and an olivine þ clinopyroxene assemblage (Fig. 15) and compiled literature data from melting experiments of various source lithologies (e.g. dry
peridotite versus carbonated peridotite) close to our
thermobarometric estimates. For dry peridotite we used
the melt composition generated by melting KLB-1 at
15 GPa and 1350 C (Hirose & Kushiro, 1993).
Experiments on carbonated peridotite are generally performed at much higher pressures than those on dry
peridotite; thus we selected the melt composition generated by melting a peridotite carbonated with 10 wt %
CO2 (PERC-3) at 30 GPa and 1350 C [run A509 of
Dasgupta et al. (2007)]. For each of these starting primary magma compositions we then calculated fractionation paths for two scenarios: (1) simultaneous
crystallization of clinopyroxene (composition of L3-12
clinopyroxene phenocrysts) and olivine (composition of
L3-12 antecryst) in the relative proportion 4:1; (2) a twostep model involving first 20 vol. % fractionation of olivine only, followed by simultaneous crystallization of a
Journal of Petrology, 2015, Vol. 56, No. 9
1765
2.0
2
1
20% Ol
PERC-3
1)
5
1.5
l
10
:O
KLB-1 1350°C
HK-66
Mix-1G
PERC-3 1350°
PERC-3
PERC
Oh e (Eger) Rift
OR melilitites
Oberpfalz
Komorní h rka
M tina Maar
ZH whole rocks
ZH matrix glasses
this study
literature data
PERC
@1350°C
CaO/Al2O3
= 2.62
:
(4
(4:
1)
20
increasing F
x:O
l
1.0
20% Ol
Cp
CaO/Al2O3
x
Cp
50
50
3 GPa, 1500°C
0.5
50
1 GPa, 1250°C
50
0.0
0
5
10
15
20
25
MgO [wt.%]
Fig. 15. Variation of MgO (wt %) vs CaO/Al2O3. Blue diamonds, melting experiments of a carbonated peridotite (PERC and PERC-3
with 25 and 10 wt % CO2, respectively) performed by Dasgupta et al. (2007) at 30 GPa pressure and temperatures from 1300 to
1600 C. Melting trends are indicated by blue arrows. Orange diamonds, experimental melt compositions (10–30 GPa, 1250–
1500 C) from Hirose & Kushiro (1993) of a natural spinel peridotite (HK-66); green diamonds, results from melting experiments
(10–25 GPa, 1375–1500 C) on a garnet pyroxenite (Mix-1 G) by Hirschman et al. (2003). Hexagons, starting melt compositions
[blue, PERC-3 at 1350 C and 30 GPa; orange, KLB-1 spinel lherzolite at 1350 C and 15 GPa of Hirose & Kushiro (1993)] for a simple
Rayleigh fractionation model involving olivine and an olivine–clinopyroxene assemblage. Mineral/liquid partition coefficients used
for MgO, CaO and Al2O3 are mean values of the experimentally determined range for olivine (Beattie, 1994) and clinopyroxene
(Adam & Green, 2006). Tick marks correspond to 1%, 2%, 5%, 10%, 20% and 50% crystallization. Data for the Ohře Rift and
Oberpfalz are shown for comparison (see Fig. 6 for references).
combined olivine–clinopyroxene crystal assemblage
similar to the first scenario (cpx:ol ¼ 4:1).
Fractionation of plagioclase or an olivine þ plagioclase assemblage would result in a slight decrease in
CaO/Al2O3 and is not evident from compositional or
petrographic observations. Exclusive fractionation of
olivine does not change CaO/Al2O3 significantly but is
effective in explaining the relative compositional difference between glasses and whole-rocks (involving about
20 vol. % olivine accumulation). However, with the
onset of fractional crystallization of clinopyroxene, CaO/
Al2O3 changes dramatically. One of the major differences between melts of carbonate-bearing and carbonate-free mineral assemblages is the large contrast in
CaO/Al2O3. In carbonate-free peridotite, CaO/Al2O3 increases with increasing degree of partial melting but is
always less than 10 owing to the excess of Al2O3 over
CaO (Fig. 15). However, during melting of carbonatebearing peridotite and in experiments performed at
constant pressure (PERC and PERC-3 at 30 GPa;
Dasgupta et al., 2007), CaO/Al2O3 values are controlled
by the changing CaO contents of the partial melts rather
than by changes in Al2O3 concentrations. At low degrees of partial melting, CaO concentrations are very
high because of non-modal melting dominated by clinopyroxene (high CaO/Al2O3). Further melting at increasing temperatures consumes clinopyroxene (‘cpx-out’)
with the result of decreasing CaO and increasing Al2O3
(melting of garnet) content in the melt (Dasgupta et al.,
2007). The experimental evidence for various factors
exerting control on CaO/Al2O3 values explored above is
thus consistent with production of the parental magma
for Železná hůrka lavas as low-degree partial melts of
carbonated peridotite.
Mantle sources and melting
Isotopic constraints
A distinct mantle reservoir has been proposed as a
common source component of the Cenozoic mafic alkaline magmatism in Europe [i.e. ‘component A’ of Wilson
& Downes (1992); ‘Low Velocity Component’ of Hoernle
et al. (1995); ‘European Asthenospheric Reservoir’ of
Cebriá & Wilson (1995)]. However, an alternative view
explains this common reservoir by variable degrees of
mixing between at least three distinct mantle sources
(e.g. Haase & Renno, 2008). Notably, basalts from
Lower Silesia extend to more radiogenic 143Nd/144Nd
values (Blusztajn & Hart, 1989; Fig. 9a) than the proposed European mantle reservoir, close to the prevalent
mantle (PREMA) composition of Stracke (2012), thus
putting into question whether there is a unique
European mantle reservoir (Fig. 9a). For Železná hůrka,
Sr and Nd isotope compositions may also point towards three-component mixing (e.g. PREMA, EM1,
EM2), as suggested by Haase & Renno (2008), but could
also be explained by variable amounts of crustal contamination. Slightly elevated 87Sr/86Sr ratios would be
consistent with a limited amount of crustal assimilation,
as discussed above. However, the observed range in
1766
Sr–Nd isotope space could also be consistent with a
parental melt linked to carbonatite; the Železná hůrka
samples plot at the overlap between the HIMU mantle
end-member and oceanic and continental carbonatites
(Hoernle et al., 2002), again indicating a minor influence
of upper continental crust assimilation.
Our new data include the first 176Hf/177Hf data from
the Ohře Rift. Samples from Železná hůrka plot close to
the mantle array of Vervoort et al. (1999) in the overlap
of the MORB, EM2 and HIMU fields (Pfänder et al., 2012;
Fig. 9b). The kimberlites of Udachnaya East
(Kamenetsky et al., 2009b) plot at the low end of the
HIMU field in (eNd)i–(eHf)i space (Fig. 9b). This observation is consistent with constraints on the nature of the
HIMU, EM1 and EM2 mantle reservoirs by Jackson &
Dasgupta (2008). HIMU is characterized by silica deficiency (43 wt % SiO2) and high CaO/Al2O3 (11), and
is likely to have evolved from carbonation of peridotite.
In contrast, EM1 and EM2 have higher silica and lower
CaO/Al2O3 (>10) and may represent mantle peridotite
enriched with sediments and by low-degree partial melt
metasomatism, respectively (Jackson & Dasgupta,
2008).
The geochemical challenge to be addressed here is
to identify and quantify the distinct mantle reservoirs
that contribute to magma generation below Železná
hůrka, especially in terms of their asthenospheric or
lithospheric (metasomatic) origin. Melting owing to upwelling of asthenospheric mantle in response to extensional tectonics is indicated by the abundant
occurrence of basaltic lavas in the Ohře Rift where the
lithosphere was thinned to about 80 km thickness (e.g.
Babuška & Plomerová, 2010). Widespread mixing of an
isotopically depleted (asthenospheric) component and
an isotopically more enriched (lithospheric) component
may be an efficient way to explain the isotopic compositions of most CEVP samples (including Železná hůrka),
as originally proposed by Wilson & Downes (1991).
Further insights from trace element modelling
We have shown above that Sr–Nd–Hf isotope data indicate the presence of a mixed mantle source for the
CEVP magmatism. Before we further constrain the
nature of the metasomatic agents, we demonstrate
below the need for a heterogeneous and enriched mantle source by melt modelling.
Pfänder et al. (2012) developed a model in which
they demonstrated the effect of melt metasomatism on
the composition of the subcontinental lithospheric mantle. The SCLM is characterized by an enrichment in trace
elements reflecting its ability to freeze infiltrating melts
and preserve their components in the form of volatilebearing minerals (e.g. amphibole, phlogopite). This
model reproduces the geochemical variability observed
in the CEVP and Ohře Rift lavas by variable proportions
of mixing between an asthenospheric melt component
and a lithospheric melt component, each of which
may be the results of variable degrees of partial melting
Journal of Petrology, 2015, Vol. 56, No. 9
(Fig. 16). At the low degrees of melting assumed for the
metasomatizing melt, Nb can be fractionated efficiently
from other elements (Pfänder et al., 2007) leading to elevated Nb/Ta at low Zr/Nb (Fig. 16a) or high La/Yb at high
Nb/La (Fig. 16b). The results of melting spinel- or garnet-peridotite and the modelled metasomatized SCLM
(‘lithospheric melts’) are shown in Fig. 16. In this model,
the Quaternary volcanic rocks from the Ohře Rift together with the SW German melilitites show the strongest signature of a contribution from a metasomatized
mantle source.
Composition and evolution of the mantle source
Several studies of mantle xenoliths have demonstrated
that the lithospheric mantle beneath the Ohře Rift is significantly altered by metasomatic processes (e.g.
Geissler et al., 2007; Puziewicz et al., 2011; Ackerman
et al., 2013, 2014). As a result, the lithology of mantle
xenoliths is bimodal with a refractory (mainly harzburgitic) peridotite suite and a metasomatic pyroxenite
suite (e.g. Geissler et al., 2007; Puziewicz et al., 2011).
The style of metasomatism is variable and ranges from
carbonatitic melt infiltration to ‘Fe-metasomatism’ as a
result of alkaline silicate melt infiltration (Puziewicz
et al., 2011). Physical evidence for the presence of carbonatitic melts is recorded in silicate and silicate–carbonate melt pockets in xenoliths of the Oberpfalz (Zinst,
Hirschentanz and Teichelberg), hosting subhedral
phenocrysts of olivine, clinopyroxene and combinations of the two, as well as carbonate minerals and ilmenite (e.g. Ackerman et al., 2013). However, if we
assume that fertile lithologies (e.g. metasomatic clinopyroxene-rich veins) melt preferentially (e.g. Foley,
1992; Phipps Morgan & Morgan, 1999), then these
xenolith suites would probably represent fragments of
the more refractory residual mantle, rather than the actual magma source.
Our new petrological and geochemical data provide
some additional evidence for the role of carbonatitic
melt infiltration and reaction with mantle peridotite.
Several studies of the CEVP have shown O isotopes to
be a powerful tracer of mantle metasomatism (e.g.
Kempton et al., 1988). High d18O in combination with
high CaO/Al2O3 should thus provide strong evidence for
the interaction of carbonatite and peridotitic mantle.
Indeed, olivines entrained in Železná hůrka lavas have
d18O significantly higher (up to þ56% V-SMOW) than
normal mantle olivine (þ52 6 02% V-SMOW) and extend into the field of fresh olivines entrained in the
Udachnaya East kimberlite (Fig. 17). Similarly, glasses
(‘d18OWR’; Fig. 11) have a heavier O isotope signature,
even though these may have been altered to lighter values by degassing, as indicated by the high vesicularity
of the Železná hůrka scoria (up to 30 vol. %) and the
high concentration of S and Cl, two elements with a
lower volatility than H2O or CO2 (likely to be degassed).
The corresponding decrease in d18OWR could be of the
order of 1–2% (up to –04% per 10 wt % volatile loss;
Journal of Petrology, 2015, Vol. 56, No. 9
1767
24
0.5%
0.7%
(a)
0.7%
22
0.7%
1%
1%
1%
20
Nb/Ta
Sp
l-P
eri
do
tite
2%
3%
2%
2%
18
5%
Grt
-Pe
rido 3%
tite
10%
16
30%
metasomatised
Spl-Peridotite
this study
literature data
Nephelinites
Other lavas
14
12
1
2
3
4
5
6
7
8
Zr/Nb
Vogelsberg
Eifel
Rhön
90
80
(b)
1
lithospheric
melts
2
70
H-group melts
0.7
SWG melilitites
OR melilitites
L-group melts
Ohře Rift
Mýtina Maar
Komorní hůrka
ZH whole rock
ZH glasses
1
La/Yb
60
3
50
2
40
5
30
10
20
10
10
5
10
30
0
0.0
0.5
1
5
10
5
2
3
1
3 2
1.0
1.5
0.7
0.5
Grt
ic
her
Spl
p
s
0.5 theno lts
as
me
0.7
2.0
2.5
3.0
Nb/La
Fig. 16. (a) Nb/Ta vs Zr/Nb for Quaternary and older volcanic rocks of the Ohře Rift and the CEVP (data sources as in Fig. 6). Colored
dashed curves indicate melting trends for garnet peridotite and spinel peridotite and a metasomatized, refractory spinel peridotite.
[For full details see Pfänder et al. (2012).] (b) La/Yb vs Nb/La for the same samples as in Fig. 13 (data sources as in Fig. 6) with melting trends for asthenospheric melts and lithospheric melts according to Pfänder et al. (2012). L- and H-group melts (moderately and
highly trace element enriched melts, respectively) are derived from melting an amphibole- and phlogopite-bearing spinel peridotite
with an enriched composition that has been constrained from natural mantle xenoliths from the Hessian depression [see Pfänder
et al. (2012) for details]. Numbers adjacent to the model curves indicate the per cent melting.
Eiler, 2001) and initial d18O would then be similar to the
range of CEVP lavas (up to þ75% V-SMOW) reported
by Mayer et al. (2014). In combination with CaO/Al2O3
we can further show that there is a positive correlation
between heavier O isotope compositions and Ca excess
and that this trend is contrary to crustal assimilation
trends (Fig. 11b).
Additional details on the nature of the metasomatic
agent may be resolved by considering the distinct trace
element composition of the Železná hůrka lavas
1768
Journal of Petrology, 2015, Vol. 56, No. 9
(Zr/Sm)N (Fig. 18b). Whereas the change in (Ti/Eu)N
could also be attributed to assimilation of material of
the continental crust (UCC and LCC in Fig. 18b), this process is insufficient to explain the subchondritic Zr/Sm
ratios, which provides strong evidence for carbonatite
metasomatism in the SCLM (Pfänder et al., 2012).
The genetic link between kimberlites, melilitites
and nephelinites
Fig. 17. d18O of olivine vs forsterite content for Železná hůrka
compared with data from South African melilitites (Day et al.,
2014) and the Azores (Genske et al., 2013), where the oxygen
isotope composition of olivine is related to assimilation–fractional crystallization processes (AFC; assimilation of altered
oceanic crust; blue arrow). An assimilation trend expected for
continental crustal material is shown in orange [bulk continental crust Mg# ¼ 55 (Rudnick & Gao, 2003), d18O 65%], assuming a melt–olivine fractionation of 05% and a d18O of 7–14%
for granitoid crust (Eiler, 2001). It should be noted that the Forich olivines with d18O higher than the mantle array (Mattey
et al., 1994) extend into the field (dark grey) of fresh olivines
from the Udachnaya East kimberlite (Kamenetsky et al., 2008).
[especially the transitional composition of glasses between the nephelinite and melilitite rock series (Fig. 7b)]
compared with other CEVP volcanic rocks. Carbonatitic
melt metasomatism is an effective process by which to
fractionate certain trace element ratios such as Ba/Th,
K/La, Zr/Sm or Ti/Eu (e.g. Sweeney et al., 1995; Yaxley
et al., 1998). More specifically, large ion lithophile elements (LILE), LREE and HREE, and also Th, Nb, Ta and Sr
become moderately to highly enriched, and decoupled
from Ti abundances (e.g. Green & Wallace, 1988;
Yaxley et al., 1991, 1998). These characteristic enrichments are easily recognizable in multi-element patterns
(Fig. 8a) with high concentrations of the LILE, a positive
Nb and Ta anomaly, and slight negative anomalies of
Zr, Hf and Ti. As a result, the Quaternary volcanic rocks
from the Ohře Rift (along with lavas from the Oberpfalz)
plot at the high end of the CEVP range in CaO/Al2O3 vs
(La/Yb)N (Fig. 18a). This range is even further extended
by the melilitites of the Ohře Rift and those from SW
Germany. For the latter, an origin by partial melting of a
carbonated peridotite source has been proposed by
Dunworth & Wilson (1998). It has to be noted, however,
that La can also be fractionated from Yb by melting in
the stability field of garnet, as HREE are retained as
compatible elements in residual garnet. More unambigous is the fractionation of Zr from Sm, a process that
has been associated with carbonatitic melt metasomatism (e.g. Pfänder et al., 2012). Samples from the
Oberpfalz, the Massif Central, the SW German melilitites and Quaternary volcanic rocks of the Ohře Rift
show the strongest fractionation of (Ti/Eu)N relative to
Forsterite-rich, high-Ni olivine cores may be interpreted
either as the earliest crystallization products of a melt
that infiltrates the lithospheric mantle (e.g. Dunworth &
Wilson, 1998) or as direct reaction products of carbonatitic melt infiltration into the SCLM. This reaction of enstatite and dolomite to forsterite, diopside and melt
(e.g. Yaxley et al., 1991; Dalton & Wood, 1993) would be
a plausible explanation for the close similarities in mineralogy and chemical composition between the Železná
hůrka nephelinites and kimberlites. Olivines entrained
in the lava show the same compositional trends as
fresh olivines in the Udachnaya East kimberlite
(Kamenetsky et al., 2008), clinopyroxenes have similar
trace element patterns, and accessory phases such as
phlogopite and spinel sensu lato suggest a genetic link.
Further evidence for the major role of a carbonate
phase during the petrogenesis of the Železná hůrka
magmas comes from radiogenic (enriched mantle signatures) and stable isotopes (high d18O of olivines and
glasses) and major (e.g. high CaO/Al2O3, high Cl and S
concentrations) and trace elements (e.g. low Ti/Eu and
Zr/Sm, Fig. 18b). Shallow oxidation owing to pressure
release indicated by titanomagnetite and haüyne
phenocrysts in the Železná hůrka lavas is also observed
in kimberlites (Yaxley et al., 2012).
However, there are also important differences between ‘high-carbonate’ lavas (such as kimberlites) and
the nephelinites of Železná hůrka. First of all, there is
evidence for much shallower depths of melting and
melt segregation for the nephelinites and a general difference in their eruption style. Kimberlites form major
diatremes that indicate explosive eruptions, whereas
the Quaternary nephelinites in the Ohře Rift may erupt
extrusively (Komornı́ Hurka, Železná hůrka vent) or explosively when the ascending melt comes into contact
with aquifers (Železná hůrka tephra, Mýtina Maar). The
elemental and isotopic characteristics of the nephelinites are also less extreme than those of kimberlites,
which, at least in the case of the fresh Udachnaya East
kimberlite, show a HIMU-like isotope signature and a
much higher degree of silica-undersaturation (32 wt %
SiO2; Kamenetsky et al., 2009b). This is also reflected in
the abundance of clinopyroxene, which is a major
phase in nephelinites but rarely present in kimberlites
(e.g. hosted as inclusions in olivine; Kamenetsky et al.,
2009a).
However, our new data provide direct magmatic
evidence for a genetic link between kimberlites, melilitites and nephelinites as suggested previously by
Journal of Petrology, 2015, Vol. 56, No. 9
1769
80
(a)
70
60
(La/Yb)N
50
40
30
20
10
0
0.0
0.4
0.8
1.2
1.6
2.0
2.4
CaO/Al2O3
(b)
SWG melilitites
Ohře Rift
OR melilitites
1.2
PM
Oberpfalz
(Ti/Eu)N
Mýtina Maar
Komorní hůrka
ZH whole rocks
ZH glasses
0.8
0.6
Nephelinites
Other lavas
UCC
0.4
LCC
0.2
0.0
0.0
3.2
Vogelsberg
Eifel
Rhön
1.4
1.0
2.8
0.4
0.8
1.2
1.6
2.0
2.4
2.8
3.2
(Zr/Sm)N
Fig. 18. (a) Chondrite-normalized (Palme & O’Neill, 2003) La/Yb vs CaO/Al2O3. Quaternary rocks from the Ohře Rift and SW German
melilitites point towards a mantle source metasomatized by carbonatitic melts (e.g. characteristic trace element enrichment and
high CaO/Al2O3). (b) (Ti/Eu)N vs (Zr/Sm)N [normalized to Cl-values of Palme & O’Neill (2003)]. Compositions of upper (UCC) and
lower continental crust (LCC) according to Rudnick & Gao (2003) and the value for the Primitive Mantle (PM; Palme & O’Neill, 2003)
are shown for comparison. Assimilation of continental crust or partial (batch) melting of peridotite may effectively lower (Ti/Eu)N
but cannot explain the fractionation of (Zr/Sm)N. These trace element ratios point clearly towards a carbonated mantle source. Data
sources as in Fig. 13.
experimental studies (e.g. Lee & Wyllie, 1997). We suggest that the dominant controls on the parental melt
composition are the amount of carbonate involved, the
depth of melting and melt segregation, and the amount
of melt–rock reaction (assimilation of peridotite wallrock during melt ascent). Kimberlites would represent
one endmember with high amounts of carbonate, a
great depth of melt segregation (erupting in thick cratonic lithosphere) and minor interaction with the residual peridotite mantle. Rifting and lithosphere
thinning may produce preferential pathways for alkalicarbonate melts or fluids to migrate towards the surface, with the potential to infiltrate and metasomatically
enrich the SCLM (e.g. Dalton & Wood, 1993; Giuliani
1770
et al., 2012). Nephelinites could then form either directly
from the carbonatitic melt percolating and assimilating
residual peridotite or by earlier mantle metasomatism
and later reactivation of these veins, or a combination
of the two. Pfänder et al. (2012) proposed an age of
mantle metasomatism in Central Europe of 100 Ma,
which would argue for mantle metasomatism
decoupled from recent volcanism. However, there is
also evidence for the presence of crustal fluids that
could be linked to recent melt infiltration into the SCLM.
Even though the depths of earthquake swarms in the
area of Nový Kostel (linked to migrating fluids) are shallower (85–95 km depth; Horálek et al., 2000) than
our estimates of the depth of melt migration (>20 km;
Fig. 14), near-surface fluids carry a strong mantle isotope signal (e.g. high 3He/4He; Weinlich, 2013).
CONCLUDING REMARKS
Previous studies have shown that the CEVP parental
magmas formed by melting variable proportions of
metasomatized (enriched) subcontinental lithospheric
mantle and depleted asthenospheric mantle. The nephelinites of the Quaternary Železná hůrka volcano in the
Ohře Rift provide strong mineralogical and chemical
evidence for the nature of this mantle metasomatism,
involving alkaline-carbonate melts or fluids, as follows.
1. Olivine antecrysts entrained in the nephelinite lava
show chemical evidence for crystallization in the
mantle, subsequent ‘normal’ crystallization and a
later overprint reflecting solid-state diffusion; this
evolutionary pattern is similar to that observed in
fresh olivines of the Udachnaya East kimberlite.
2. The trace element patterns of clinopyroxenes also
resemble those of kimberlitic clinopyroxenes and
their crystallization conditions argue for a continuous process in the upper SCLM and lower crust rather than a long period of melt stagnation at a
distinct level in the lithosphere.
3. The compositions of accessory phlogopite are intermediate between those in lamproites and carbonatites; the presence of spinels sensu lato and haüyne
argues for shallow oxidation of the melt, similar to
observations in carbonate-rich melts.
4. Elevated O isotope ratios (relative to mantle values)
and distinct trace element enrichment (e.g. LILE,
LREE, Nb and Ta) and trace element ratios (e.g. Ti/
Eu, Zr/Sm) provide further evidence for a contribution of carbonatite to the final melt composition.
Radiogenic Sr–Nd–Hf isotope data support this view,
although their signatures may also be explained by
processes other than carbonate melt metasomatism.
Furthermore, we have shown that crustal assimilation may play a role in the petrogenesis of the Železná
hůrka nephelinite, but is insufficient to account for all
the mineralogical and chemical evidence we have
found for carbonate melt–peridotite interaction. The
Journal of Petrology, 2015, Vol. 56, No. 9
genetic link between kimberlites, melilitites and nephelinites has been previously suggested based on experimental and xenolith studies, but this study provides
direct magmatic evidence. The depth of alkali-carbonate
melt–fluid segregation, its total volume and the proportion of peridotite assimilation during melt ascent
through the mantle may control the final magma type.
ACKNOWLEDGEMENTS
We collected our samples without using mechanical
tools to avoid any damage on the protected outcrop of
Železná hůrka and we would like to encourage every
visitor to this location to help to preserve it in its current
condition. We thank H. Brätz and M. Hertel at
GeoZentrum Nordbayern and N. Pearson at GEMOC for
their analytical help. We also acknowledge the co-operation and support of A. Weh and the Selfrag AG
(Kerzers, Switzerland) for their help with high-voltage
pulse power fragmentation of olivine-phyric rocks. L.
Ackerman, S. Jung, J. Pfänder and editor M. Wilson are
acknowledged for comments that significantly improved the quality and clarity of this paper. P.A.B.
thanks G. Yaxley and O. Nebel for constructive comments on an earlier version of this paper.
FUNDING
This work was supported by a grant of the
‘Sonderfonds für wissenschaftliche Arbeiten an der
Universität Erlangen–Nürnberg’ to P.A.B. and F.S.G.
and by funding through grant WI 3675/1-1 from the
Deutsche Forschungsgemeinschaft. P.A.B. benefited
from a Feodor Lynen Research Fellowship of the
Alexander von Humboldt Foundation.
SUPPLEMENTARY DATA
Supplementary data for this paper are available at
Journal of Petrology online.
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