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Title: Middle Miocene volcanism in the vicinity of the MiddleHungarian zone: evidence for an inherited enriched mantlesource – a review
Authors: I. Kovacs, Cs. Szabo
PII: S0264-3707(07)00048-8DOI: doi:10.1016/j.jog.2007.06.002Reference: GEOD 815
To appear in: Journal of Geodynamics
Received date: 15-2-2007Revised date: 3-5-2007Accepted date: 21-6-2007
Please cite this article as: Kovacs, I., Szabo, Cs., Middle Miocene volcanism in thevicinity of the Middle Hungarian zone: evidence for an inherited enriched mantle source– a review, Journal of Geodynamics (2007), doi:10.1016/j.jog.2007.06.002
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Middle Miocene volcanism in the vicinity of the Middle Hungarian 1
zone: evidence for an inherited enriched mantle source – a review 2
3
Kovács, I.1,2 *, Szabó, Cs.1,3 4
5
1) Eötvös University, Department of Petrology and Geochemistry, Lithosphere Fluid Research Lab, H-6
1117, Bp. Pázmány Péter sétány 1/C., Budpest, Hungary 7
2) Australian National University, Research School of Earth Sciences, Mills Road, Blg 61, 0200, Canberra, 8
ACT, Australia 9
3) Division of Earth Environmental System, College of Natural Sciences, Pusan National University, Busan 10
609-735, Korea 11
12
*corresponding author: [email protected] 13
* Manuscript
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Abstract 14
15
Middle Miocene igneous rocks in the vicinity of the Middle Hungarian zone 16
(MHZ) show a number of subduction-related geochemical characteristics. Many of these 17
characteristics appear to be time-integrated, showing a decreasing subduction signature 18
with time. In contrast to previous models, which suggest southward-dipping subduction 19
of European lithosphere beneath the Alcapa microplate (along the Western Carpathians) 20
is responsible for the chemical characteristics seen in middle Miocene volcanics, we 21
propose that source enrichment occurred via the subduction of either the Budva-Pindos or 22
Vardar Oceans. Recent seismic studies have revealed that the proposed southward-23
dipping subduction was not developed beneath the entire Western Carpathians or, even if 24
it had, was overprinted by the collision of the European plate and the Alcapa unit at 16 25
Ma. This subduction is thought to have started 30 Ma ago, therefore the time between the 26
onset of subduction and collision cannot account for extensive source enrichment in the 27
overlying mantle wedge. It is also pertinent to note that the middle Miocene igneous 28
rocks of the MHZ in their reconstructed positions are not parallel to the supposed suture 29
expected for subduction-related arc volcanoes. 30
Our review suggests an alternative hypothesis, whereby source enrichment is 31
related to the subduction of either the Budva-Pindos or Vardar Ocean during the 32
Mesozoic-Paleogene. In this model the Alcapa microplate was transferred to its present 33
tectonic position via extrusion and rotations. Geophysical modeling and mantle xenoliths 34
provide evidence that this process occurred at the scale of the lithospheric mantle, 35
indicating that the subduction-modified lithospheric mantle was coupled to the crust. 36
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Melting in the lithospheric mantle of the Alcapa unit was triggered by the extension 37
during the formation of the Pannonian Basin. The preserved subduction-related 38
geochemical character of volcanics in intra-plate settings that are otherwise directly 39
unaffected by subduction, can be attributed to tectonic transport of metasomatised mantle 40
from a previous subduction-affected setting. This model provides an alternative approach 41
to understanding the geochemical complexity seen among intra-plate calc-alkaline 42
volcanics, where chemical characteristics can be explained without the involvement of 43
plumes. 44
45
Key words: Carpathian-Pannonian region, middle Miocene, volcanics, mantle xenoliths, 46
mantle tomography, geodynamics 47
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1. Introduction 48
49
The Tertiary and Quaternary volcanism of the Carpathian-Pannonian region 50
(CPR) has been the focus of extensive recent research because the area provides an 51
excellent opportunity to understand the interplay between tectonics and volcanism 52
(Chalot-Prat and Boullier, 1997; Chalot-Prat and Girbacea, 2000; Downes et al., 1995; 53
Harangi, 2001b; Harangi et al., 2001; Harangi et al., 2006; Harangi and Lenkey, 2007; 54
Konečny et al., 2002; Konečny and Lexa, 1999; Pécskay et al., 2006; Pécskay, 1995; 55
Póka, 1988; Póka et al., 2004; Salters et al., 1988; Seghedi et al., 1998; Seghedi et al., 56
2004a; Seghedi et al., 2005; Szabó et al., 1992; Zelenka et al., 2004; Vaselli et al., 1995). 57
These studies revealed that most of the Tertiary igneous rocks in the CPR were derived 58
from a subduction-enriched source, with both the degree of crustal contamination and 59
strength of the subduction-related geochemical signature (i.e., high LILE/HFSE ratios, 60
negative Nb, Ta and Ti anomalies; lower 87Sr/86Sr and higher 143Nd/144Nd radiogenic 61
isotopic ratios) decreasing with time. Most younger Plio-Pleistocene alkaline basaltic 62
rocks of the region, which mainly postdated the main phase of calc-alkaline volcanism 63
(with the exception of the Eastern Transylvanian basin where it was contemporaneous), 64
show very little evidence for subduction. These Plio-Pleistocene basalts show 65
geochemical affinities similar to that of an OIB-like asthenospheric source (Dobosi et al., 66
1998; Embey-Isztin et al., 1993b; Seghedi et al., 2004b). This being the case the 67
relationship between the onset of proposed subduction along the Carpathians and the 68
climax of magmatism still remains a matter of debate. Most previous studies have dealt 69
with the whole CPR and distinguished different volcanic provinces based on 70
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geochemistry and spatial distribution. In this paper we focus only on the middle Miocene 71
volcanics of the “Western segment” (see Fig. 1) which were formed during the first phase 72
of extension in the CPR (Husimans et al., 2001). This enables us to gain better spatial 73
and temporal resolution and put forward an alternative explanation for the subduction-74
related source enrichment of these magmas. In addition to temporal and spatial 75
considerations, this paper incorporates recent results and seismic tomography from the 76
region. Although integration of such information can be used to help decipher the 77
geodynamic evolution of the region, such attempts have been seen limited considerations 78
in previous studies. 79
The most recent reviews on the Tertiary volcanism of the region (Harangi, 2001b; 80
Harangi and Lenkey, 2007; Seghedi et al., 2004a; Seghedi et al., 2005) provide no 81
alternative to the present geodynamic paradigm of the Carpathian-Pannonian region. 82
This paradigm suggests that magmas were mainly produced by contemporaneous 83
subduction-related arc volcanism along the Carpathians. Existence of a classic 84
subduction margin for the entire Western Carpathians is now questioned by several 85
authors (i.e., Grad et al., 2006; Szafián and Horváth, 2006; Szafián et al., 1997; Szafián et 86
al., 1999, Tomek, 1993). There is evidence to suggest the exsistence of subduction along 87
eastern sections of the Western Carpathians - the roll-back effects of which can be used to 88
account fro the observed block-rotations and fault patterns in the Northern Pannonian 89
Basin (i.e., Bielik et al., 2004; Márton and Fodor, 2003; Sprener et al., 2002). The 90
reconstructed position of middle Miocene volcanics of the “Western segment” in their 91
reconstructed position are not parallel to the hypothesized southward subduction zone 92
along the Carpathian mountain belt, but in fact, were generated in the vicinity of the 93
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Middle Hungarian zone (MHZ, Fig. 1). Here the term “vicinity” refers to a zone 300-350 94
km wide and is centered on the MHZ. This width is not an arbitrary number, but is twice 95
the distance at which typical arc-volcanoes occur relative to the trench (i.e., Ellam and 96
Hawkesworth, 1988; Hawkesworth et al., 1994). Paleogene-early Miocene igneous rocks 97
show a very similar spatial distribution close to the MHZ, which suggests there may be a 98
link between the middle Miocene and this magmatic episode (Fig. 1, Kovács et al., 2007). 99
This observation raises the possibility that subduction-related geochemistry is not related 100
to any contemporaneous subduction, but it is an “inherited” feature from a previous 101
geodynamic setting. This “inherited” model is a known model and its overview is given 102
before the discussion. Our main aim in this paper is, therefore, to address the subduction-103
related source enrichment at the origin of the middle Miocene igneous rocks along the 104
MHZ. We consider all potential subduction zones which may account for this source 105
enrichment and examine how each subduction zone matches the temporal, and 106
geodynamic characteristics of volcanism seen in the CPR. 107
108
2. Geological background 109
110
The wider Carpathian-Pannonian region (CPR) comprises several mountain 111
chains, i.e., the Eastern Alps, the Carpathian arc and the Dinarides, which surround the 112
Intra-Carpathian Basin System. The Intra-Carpathian Basin System, the internal part of 113
which is commonly referred to as the Pannonian Basin, has a significantly extended 114
continental crust. This extended crust outcrops in several smaller internal mountains, and 115
is filled by middle to late Miocene sediments (Fig. 1). 116
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Based on the Mesozoic tectonostratigraphy and structural analysis, the internal 117
area is subdivided into two major units (Csontos, 1995; Csontos and Vörös, 2004; Haas et 118
al., 1995; Haas et al., 2000; Kovács et al., 2000): Alcapa (ALps-CArpathians-PAnnonian, 119
i.e. northern Intra-Carpathian Basin System and Western Carpathians) and Tisza-Dacia 120
(southern Intra-Carpathian Basin System; East and South Carpathians) (Fig. 1). In this 121
paper the Alcapa unit is defined as the unit bounded to the south by the Periadriatic and 122
Balaton faults (Balla, 1984; Csontos, 1995; Fodor et al., 1999; Fodor et al., 1998). The 123
Balaton fault separates Alcapa from the Mid-Hungarian zone. The major fault separating 124
the Mid-Hungarian unit from the Tisza unit to the south is termed the Mid-Hungarian (or 125
Zagreb-Zemplin) fault (Balla, 1984; Csontos and Nagymarosy, 1998; Haas et al., 2000; 126
Kovács et al., 2000; Wein, 1969) (Fig. 1). The Mid-Hungarian zone (MHZ, Fig. 1) is 127
composed of the Bükk Mts. (N Hungary) and a narrow structural belt between Lake 128
Balaton and the Mecsek Mts. (SW Hungary). Exposures seen within this zone consist of 129
low- to high-pressure metamorphic Paleozoic-Mesozoic continental margin sediments, a 130
mélange, and dispersed remains of a Jurassic ophiolite nappe (Csontos and Vörös, 2004; 131
Haas et al., 2000; Harangi et al., 1996; Wein, 1969). The Mid-Hungarian zone was 132
strongly deformed during Tertiary times (Balla, 1987; Csontos and Nagymarosy, 1998). 133
Based on numerous structural studies and tectonic reconstructions (Balla, 1984; 134
Csontos, 1995; Csontos et al., 2002; Fodor et al., 1999), the study area was formed in 135
three major steps during Tertiary times. The earliest tectonic phase took place in the 136
Paleogene, with a major right lateral shear event along the Periadriatic zone and the 137
Balaton fault. This shear event was initiated in the Late Eocene (Fodor et al., 1992), but 138
the bulk of shearing appears to have occurred during the Oligocene. As a consequence, 139
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Alcapa was subjected to major right lateral displacement (Kázmér and Kovács, 1985). 140
This event was contemporaneous with the proposed onset of subduction along the 141
Western Carpathians ~30 Ma ago (Oszczypko, 1992; Pécskay et al., 2006; Tomek, 1993; 142
Tomek and Hall, 1993). The existence of subduction along the entire Western 143
Carpathians is now doubted by many reconstructions (i.e., Mantovani et al., 2000), which 144
proposed that the subduction had a NW-SE strike parallel to the eastern portion of the 145
Western Carpathians and the Eastern Carpathians. The remaining western part of the 146
Western Carpathians was characterized by transtensional movements with a strike 147
subparallel to ENE-WSW (Fig. 1). It is unclear then, whether well-developed subduction 148
occurred along the entire Western Carpathians (see discussion for more detail). 149
The next tectonic phase was dominated by opposite rotations of the Alcapa and 150
Tisza-Dacia microplates within the CPR (Csontos et al., 2002; Márton, 1987). Based on 151
paleomagnetic data (Márton, 1987), the two microplates were detached from their 152
southern neighbor, the Dinarides, and were extruded or rotated into the Carpathian 153
embayment during the late Oligocene and early Miocene (ca. 20-18 Ma). 154
The last major tectonic phase occurred from the Middle-Miocene to subrecent, 155
and still shapes the area today (Bada and Horváth, 2001; Csontos, 1995; Csontos et al., 156
2002; Horváth, 1993; Horváth et al., 2006; Horváth and Cloetingh, 1996). This final 157
tectonic stage includes the formation of the Carpathians and opening of the Pannonian 158
basin, which occurred in two distinct phases. The first extension phase took place in the 159
early-middle Miocene (17.5-14 Ma) which affected both the internal and external part of 160
the Pannonian basin with a β factor of 1.4-1.6 (Huismans et al., 2001). This first stage of 161
basin formation was also contemporaneous with major counter-clockwise rotation of the 162
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Alcapa block (Márton and Fodor, 2003). This tectonic phase is considered to be driven 163
by the southward and westward directed subduction of the European margin beneath the 164
internal Carpathian area (Balla, 1984; Horváth, 1993; Jolivet and Faccenna, 2000; Márton 165
and Fodor, 2003). Since the “classical” development1 of subduction along the entire 166
Carpathians is now in question, there may be other driving forces for basin formation. 167
One possible driving force is eastward directed asthenospheric flow from the Alpine 168
collision belt analogous to that seen in the Himalaya collision belt (Flower et al., 2001). 169
The collision between Alcapa and the European platform is thought to have started 16 Ma 170
ago, with collision becoming younger towards the southeast. During this time the MHZ 171
was subjected to extensive deformation between the rotating Alcapa and Tisza-Dacia 172
blocks (Csontos and Nagymarosy, 1998) and was almost in perpendicular position to the 173
remaining subduction front along the Carpathians. This phase was followed by a second 174
extension stage in the late Miocene (11-8 Ma) which is related to the gravitational 175
collapse of the asthenospheric dome. This phase affected only the central part of the CPR 176
and the mantle lithosphere with a δ factor of 6-8 (Husimans et al., 2001). At 6 Ma ago 177
there was a smaller rotation of the Alcapa block due to the change in the rotation of the 178
Adriatic indenter (Márton and Fodor, 2003). 179
180
3. Temporal and spatial distribution of middle Miocene igneous rocks 181
182
1 By the term “classical subduction” we refer to subduction of an oceanic slab, where the subducted slab reaches considerable depth (>200 km), the Benioff-zone is well-developed and the slab is capable of causing extensive metasomatism in the overlying mantle wedge (i.e., Peacock, 1991; van Keken, 2003; Schmidt and Poli, 1998)
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The Tertiary igneous rocks that cover a considerable part of the CPR have been 183
classified in different ways by different authors (see Seghedi et al., 2004a; Seghedi et al., 184
2005; Harangi, 2001; Harangi and Lenkey, 2007 for more detail). In this review we try to 185
adopt a spatial classification that combines these previous attempts, but does not discuss 186
volcanites of the Apuseni Mts (“Central segment”). Here we distinguish a “Western” and 187
“Eastern segment”. The “Eastern segment” covers igneous rocks parallel to the Eastern 188
Carpathians with NNW-SSE strike, from the Vihorlat Mts. to the East Transylvanian 189
basin (Fig. 1). These are mainly middle-Miocene-subrecent calc-alkaline rocks which 190
have been interpreted as subduction related arc volcanites (Seghedi et al., 2004a; Seghedi 191
et al., 2005). In the Persani Mts. (East Transylvanian basin, ETB) there are 192
contemporaneous Plio-Pleistocene calc-alkaline and mafic alkaline rocks whose genesis 193
is still matter of debate (cf., Chalot-Prat and Boullier, 1997; Chalot-Prat and Girbacea, 194
2000; Mason et al., 1998; Seghedi et al., 2004). 195
The “Western segment” is much more complex than the “Eastern”, showing wider 196
geochemical, spatial and temporal variability. In this study the “Western segment” 197
includes igneous rocks of the Alcapa, MHZ and northern portion of the Tisza units, from 198
the Tokaj Mts. to the Styrian basin which line up along a NE-SW directed strike that is 199
almost perpendicular to the “Eastern segment” (Fig. 1). The genesis of middle Miocene 200
igneous rocks in the vicinity of the MHZ is mainly attributed to the first phase of 201
extension (induced by subduction roll-back), that caused melting in the previously 202
metasomatised lithosphere (i.e., Harangi, 2001b; Harangi and Lenkey, 2007). 203
The middle Miocene mainly calc-alkaline, silicic and very rarely ultrapotassic 204
volcanites in the “Western segment” either built up mountain ranges as a series of 205
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volcanic edifices (that have subsequently been strongly disintegrated and eroded by now) 206
or are buried under the thick sedimentary sequence of the Great and Little Hungarian 207
plain nearly parallel to the MHZ (Fig. 1). The most important parts of the mountain 208
ranges (from west to east) are as follows: Visegrád Mts., Börzsöny Mts., Central 209
Slovakian Volcanic Field, Cserhát-Mátra Mts. Bükk Mts and Tokaj-Slanac Mts (Fig 1.). 210
In addition to volcanites observed in these ranges, Zelenka et al. (2004) reported large 211
andesitic and dacitic strato-volcanoes buried below southern Transdanubia, Danube-Tisza 212
Interfluve and Great Hungarian plain (Fig. 1). Similarly, Hajnal et al. (2004) identified a 213
thick middle Miocene igneous body buried beneath the Nyírség district via a low-214
frequency 3-D seismic survey. This area is on the boundary between the Alcapa and 215
Tisia microplates and sits on the MHZ. Partially buried volcanic rocks have been also 216
reported in the Styrian basin (SB) and Little Hungarian Plain (LHP), and are dominantly 217
potassic, intermediate to acidic middle and late Miocene volcanic rocks (Harangi et al., 218
1995; Harangi, 2001a). Pamić et al. (1995) identified a thick middle Miocene volcanic 219
succession buried in the South Pannonian Basin (including the Drava Basin) at the 220
southern side of the MHZ. Among the observed igneous rock types erupted along the 221
MHZ, there is a general younging tendency towards NE (Pécskay et al., 1995; 2006). 222
Alkaline mafic volcanism in the “Western segment” (Styrian basin (SB), Little 223
Hungarian Plain (LHP), Bakony-Balaton Highland (BBH) and Nógrád-Gömör (NG)) 224
post-dated the middle Miocene, predominantly calc-alkaline magmatism. Alkaline mafic 225
volcanism started at 11 Ma ago and was active till the Plio-Pleistocene (Pécskay et al. 226
1995, 2006). 227
228
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4. Petrology and geochemistry 229
230
The middle Miocene volcanic rocks of the “Western segment” which are 231
considered in this study are summarized in Table 1. They show wide petrologic variety, 232
ranging from basalts to rhyolite. Projected in the TAS diagram (Le Bas et al., 1986) most 233
compositions correspond to basaltic andesite and andesite fields (Fig. 2). Rare Earth 234
Elements (REE) display an enriched, convex-upward pattern relative to chondritic values 235
(Fig. 3a). General trace element patterns show a positive anomaly in Large Ion 236
Lithophile Elements (LILE) and Pb, with a negative anomaly in the High Field Strenght 237
Elements (HFSE). The enrichment of LILE and LREE over HFSE (especially Nb) is a 238
typical feature of subduction-related volcanic rocks (Ellam and Hawkesworth, 1988) 239
(Fig. 3b). Paleogene-Early Miocene igneous rocks along the MHZ are similar to Middle-240
Miocene rocks in terms of trace element patterns, displaying the same anomalies and 241
comparable concentrations (Fig. 3c and 3d). 242
Trace element discrimination diagrams also show that the geochemical 243
characteristics of more primitive middle Miocene volcanics are similar to subduction-244
related magmas. The Ba/Zr vs. Nb/Zr diagram of Pereplov et al. (2005) distinguishes 245
IAB, MORB and WPB sources (Fig. 4a). Ba/Zr ratio is indicative of fluid contribution 246
from a subducted slab, whereas Nb/Zr ratio is sensitive to the effect of slab-derived melts. 247
Middle Miocene volcanics show a transitional character, where the most primitive 248
magmas with SiO2<56 wt%2 lie between the IAB and MORB and WPB fields - 249
displaying an almost constant Nb/Zr ratio but a large variation in Ba/Zr ratio. More 250
2 These rocks represent the most primitive calc-alkaline rock that are likely to be unaffected by any contamination and fractionation.
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silica-rich samples, however, are closer to the IAB field. Paleogene-early Miocene 251
igneous rocks along the MHZ (Kovács et al., 2007) are very similar to their middle 252
Miocene counterparts, with a somewhat higher Ba/Zr ratio. Typical fields for different 253
OIB compositions (HIMU, EMI and EMII) are also displayed, showing that middle 254
Miocene and Paleogene-early Miocene rock of the MHZ fall close to the EMI 255
component. 256
A plot of Th/Y vs. Nb/Y gives constraints on different mantle source components 257
(Fig. 4b). The Nb/Y ratio can be used to differentiate among mantle sources, whereas 258
Th/Y can be used to monitor the transfer of slab-derived components to the mantle wedge 259
(Elliot et al., 1997; Peate et al., 1997; Seghedi et al., 2004a). Only middle Miocene 260
magmas with <56 wt% SiO2 were used in this plot. Their relatively high Th/Y ratio is 261
indicative of considerable sediment influx, whereas variation in Nb/Y refers to 262
heterogeneous mantle source. From the OIB components only EMI and EMII show 263
similar Th/Y values, while their Nb/Y ratios are higher. Paleogene-early Miocene rocks 264
of the MHZ display larger variation in both ratios implying significant source variation 265
and even larger sediment input. 266
Nd-Sr isotopic systematics show an enriched character, where 143Nd/144Nd and 267
87Sr/86Sr ratios are in the range of 0.5122-0.5125 and 0.706-0.714, respectively (Fig. 4c). 268
More mafic rocks (with <56 wt% SiO2) are very similar to the EMI component. 269
Paleogene-early Miocene rocks again show a very similar pattern to middle Miocene 270
volcanites (Fig. 4c). 271
Younger, mainly post-Miocene alkaline basalts in the “Western segment” 272
showing common OIB character, still have slight subduction overprint on their 273
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geochemistry (Dobosi et al., 1998, Seghedi et al., 2004a). Boron concentrations are 274
anomalously high in alkaline mafic magmas of the Bakony-Balaton Highland (BBH, Fig. 275
1), which is indicative of subduction contamination (Gméling et al., 2007). 276
277
5. Xenoliths 278
279
Over the last decades our knowledge of the lithosphere beneath the CPR has been 280
greatly improved by studies of xenoliths hosted in the Plio-Pleistocene alkaline basalts. 281
Here we attempt to demonstrate that xenoliths provide very useful information in 282
geodynamic exploration. Xenolith-bearing host basalts occur at the edge of the CPR 283
(Styrian basin (SB), Nógrád-Gömör (NG) and Eastern Transylvanian basin (ETB)) and in 284
its central part (Little Hungarian Plain (LHP), Bakony-Balaton Highland (BBH)) (Fig. 1). 285
These xenoliths in alkali basalts are of lithospheric and in some cases asthenospheric 286
origin (Szabó et al., 2004; Falus, 2004). Because eruption of the host alkali basalts 287
postdate the major tectonic events of the CPR, their xenoliths record effects preceding 288
tectonic events - which include deformation (Falus et al., 2000; Falus, 2004; Falus et al., 289
2007; Hidas et al., 2007), metasomatism and melt extraction (i.e., Bali et al., 2006; 290
Chalot-Prat and Boullier, 1997). The upper mantle xenoliths are mostly spinel peridotites 291
(i.e., Downes et al., 1992; Embey-Isztin et al., 2001; Szabó et al., 2004; Vaselli et al., 292
1995), accompanied by subordinate pyroxenites and lower crustal granulites in the CPR 293
(Dobosi et al., 2003; Embey-Isztin et al., 2003; Embey-Isztin et al., 1990; Kovács and 294
Szabó, 2005; Kovács et al., 2004; Török et al., 2005; Zajacz et al., 2007). The peridotites 295
represent residual mantle material showing textural and geochemical evidence for a 296
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complex history of melting and recrystallization, irrespective of location within the 297
region (e.g., Falus, 2004; Szabó et al., 2004; Szabó et al., 1995). This deformation has 298
been attributed to a combination of extension and upwelling of asthenosphere in the 299
middle Miocene (Szabó et al., 2004 and references therein). The mantle xenoliths 300
contain hydrous phases (pargasitic and kaersutitic amphiboles and rarely phlogopite) as 301
evidence for modal metasomatism, in all occurrences of the region. 302
For the purpose of this study only xenoliths from the BBH and NG will be 303
considered in detail. This is because these localities are in close proximity to the middle 304
Miocene volcanics of the “Western segment” with which we are concerned. The LHP is 305
also close, but the available xenolith suite is limited (Falus et al., 2007). In the BBH, 306
xenolith lithologies include extremely rare orthopyroxene-rich olivine websterites and a 307
quartz-bearing orthopyroxene-rich websterite (Bali, 2004; Bali et al., 2006). From 308
petrographic and geochemical evidence, these orthopyroxene-rich mantle rocks are 309
probably the products of interaction between subduction-related melts (i.e., adakite, 310
boninite) and mantle lithologies, as pointed out by McInnes et al. (2001) and Santos et al. 311
(2002) for xenoliths from Papua New Guinea and tectonic peridotites from Spain, 312
respectively. The original (i.e., slab-derived) melts are highly reactive in the mantle. The 313
websterites, are therefore, though to be produced in fore-arc or arc setting - because melts 314
formed in this environment cannot migrate far from their source without being 315
completely consumed (Rapp et al., 1999). Downes et al. (1992) and Rosenbaum et al. 316
(1997) have already proposed subduction-related characteristics of the BBH mantle 317
xenoliths based on their radiogenic isotope signatures. 318
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Granulite xenoliths have also been found at this locality, with Dobosi et al. (2003) 319
and Embey-Isztin et al. (2003) hypothesizing a subduction-accretion origin for these 320
lower crustal rocks. These authors went on to suggest that the protolith from which these 321
xenoliths derived was a subducted oceanic slab which got stuck at lower crustal depth. 322
Previously, Embey-Isztin et al. (1990) suggested that a considerable volume of basaltic 323
magmas were trapped and crystallized close to the MOHO beneath the BBH, causing 324
massive underplating in the region. 325
Szabó et al. (1996) presented volatile-rich silicate melt inclusions in peridotite 326
xenoliths as evidence for the existence of a metasomatic agent with subduction-related 327
origin. Granulite and pyroxenite xenoliths are also present in great abundance, with 328
lower-crustal granulites interpreted as being either the remnant of an older underplated 329
magmatic body or a subducted oceanic slab (Kovács and Szabó, 2005). These rocks 330
show indication of incipient melting around former garnets. It was also proposed that the 331
incipient melting of these granulites may have played an important role in the formation 332
of the Tertiary calc-alkaline volcanic rocks of the NG representing the “lower-crustal 333
component” (Harangi, 2001a). Kovács et al. (2004) and Zajacz et al. (2007) reported 334
clinopyroxenite xenoliths with silicate melt inclusions and concluded that they are the 335
fragments of underplated basaltic rocks trapped at the MOHO or in the uppermost mantle 336
beneath the NG. 337
338
6. Geophysical constraints 339
340
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Seismic tomographic studies in the CPR show that the remnants of subducted 341
slab(s) cannot be adequately identified. Piromallo and Morelli (2003), Spakman (1990) 342
and Wortel and Spakman (2000) defined a cool slab lying subhorizontally beneath the 343
CPR at a depth of 600 km. In the east, the Vrancea zone was thoroughly examined by 344
Oncescu et al. (1984), where a cool and steeply dipping body is found. This seismogenic 345
body was thought to be a remnant of the westward-subducted European slab. Detailed 346
seismic tomography of Wortmann et al. (2001), however, showed a very steep (but 347
slightly eastward-dipping) localized body at the same location. Geophysical constraints 348
are even more controversial for the Western Carpathians, where the existence of a 349
southward dipping subduction is ambiguous. Bielik et al. (2004) and Sroda et al. (2006) 350
found features as southward dipping reflectors, thickened crust for the central and 351
western part of the Western Carpathians deploying geophysical modeling (i.e., seimic and 352
gravitational measurements) which is compatible with the idea of a southward dipping 353
subduction of European plate beneath the Alcapa unit. In comparison to Bielik et al. 354
(2004) and Sroda et al. (2006), Tomek (1993) and Tomek and Hall (1993) suggested that 355
either the European slab did not subduct beneath the western CPR, or that slab is now 356
completely detached, disintegrated and sunk into a considerable depth. Grad et al (2006) 357
came to a very similar conclusion by studying the Carpathian transect (CEL05) of the 358
Celebration 2000 project. The CEL05 transect is NNE-SWS directed and runs through 359
the point where the W-E strike of the Western Carpathians turns into the NNW-SSE 360
strike of the Eastern Carpathians. Grad et al. (2006) found northward-dipping reflectors 361
beneath the Western Carpathians at mantle depth, which may exclude the existence of a 362
well-developed southward-directed subduction of the European lithosphere beneath 363
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Alcapa. In their interpretation, even if southward-dipping subduction existed, it would 364
have to be a relatively old feature overprinted by a younger most probably northward 365
directed thrusting of the Alcapa unit. Furthermore, 2D and 3D gravitational modeling 366
showed that a deep crustal root did not develop beneath the Western Carpathians which 367
would be expected for a “classical” convergent margin (Szafián et al., 1997; Szafián et 368
al., 1999; Szafián and Horváth, 2006). These later authors suggested alternatively that 369
this is due to the fact that the Western Carpathians are characterized by transtensional-370
transpressional plate motions with little or no collision at least since the middle Miocene. 371
The nearest unambiguous trace of a subducted slab outside the CPR has been 372
identified along the eastern shoreline of the Adriatic Sea and at its northern continuation 373
beneath the Po Plain (Koulakov et al., 2002; Lippitsch et al., 2003; Piromallo and 374
Morelli, 2003). It is a north-northeastward-dipping slab that can be detected down to 600 375
km. This slab is most probably the remnant of the former Budva-Pindos Ocean (Csontos 376
and Vörös, 2004). 377
378
7. An overview of the “inherited” enriched source theory 379
380
It is a common geological situation that magmas show subduction-related 381
geochemical signatures in places otherwise unaffected by any recent or contemporaneous 382
subduction (i.e., Basin and Range, (USA), [Fitton et al. (1991); Hawkesworth et al., 383
(1995); Harry and Leeman (1995)]; Southern Sonora, (Mexico), [Till et al. (2007)]; 384
Anatolia, (Turkey), [Pearce et al. (1990); Keskin et al. (1998); Aldanmaz et al. (2000)]; 385
Eastern Rift (Morocco), [El Bakkali et al. (1998)]; This is because the subduction-related 386
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signature from a previous geodynamic setting can be preserved in the mantle as 387
metasomatic alteration(s), which can be reactivated via melting under various 388
circumstances. 389
The subduction-related signature commonly refers to rocks which show relative 390
enrichment in LILEs (Ba, Sr), alkalis and LREEs and depletion in HFSEs (Nb, Ta, Ti, Zr 391
and Hf) (Saunders et al., 1991). This geochemical signature can originate from two main 392
sources: the subducted slab flux and the hybridized mantle flux. The former one is due to 393
the dehydration and melting of the subducted oceanic slab and interaction of these 394
melts/fluids with the overlying mantle wedge. The melting of the slab occurs in the range 395
of 90-120 km (Drumont and Defant, 1990), whereas the dehydration is a more or less 396
continuous process from the compaction at relatively low pressures until the break-down 397
of the most stable hydrous phases at deep-mantle depth (Ringwood, 1991; Frost, 2006; 398
Schmidt and Poli, 1998). During these processes LILEs and LREEs go preferentially into 399
the fluid phase, whereas HFSEs are retained in the actually stable titanium phase (titanite, 400
rutile) (Saunders et al., 1991 and references therein). This is generally what leads to the 401
formation of the “subduction” signal in mantle wedges. The second geochemical source, 402
the hybridized mantle flux, is the result of melting in the previously metasomatised 403
portion of the mantle wedge (i.e., “hybridized mantle”) above subducting slabs. This 404
remelting produces a similar subduction-related geochemical signature that can differ 405
somewhat depending both on P-T-fO2 conditions and solid phases present in the 406
residuum. It is for this reason that intra-plate magmas with subduction signature usually 407
show larger variation in geochemistry than arc magmas (Foley et Wheller, 1990). An 408
additional mechanism for producing subduction signature is via delamination and melting 409
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of over-thickened eclogite-granulite facies lower crust. In this case melts/fluids released 410
from this delaminated lower-crust resemble those of subducted oceanic crust and 411
metasomatise the asthenosphere filling up the gap behind the detached lower crust (i.e., 412
Lustrino and Wilson, 2007; Lustrino et al., 2007) 413
The mantle wedge consists of the original MORB-like subarc mantle with 414
metasomatism-related magmatic veins and hydrous minerals, thought to be due to 415
metasomatism occurring above a subducting slab. These veins are usually of tonalitic-416
dacitic character, which represents melting of the subducted slabs (Rapp et al., 1999; 417
Ringwood, 1991). At the contact of these dikes one often finds pyroxenite and websterite 418
reaction zones (i.e., Bali et al., 2007a and reference therein). There can also be hydrous 419
phases in both the mantle wedge and reaction zones (pargasite, phlogopite), whose modal 420
proportion can be up to 10% (Niida and Green, 1999; Wallace and Green, 1991). This 421
means that an extensively metasomatised mantle portion can contain up to 0.4 wt % of 422
H2O (Hawkesworth and Gallagher, 1993). The presence of hydrous phases suggests that 423
melting temperature can be significantly lowered relative to that of MORB mantle (i.e., 424
Green and Lieberman, 1976; Green and Fallon, 2005). 425
These metasomatic alterations can be preserved both in the subcontinental 426
lithospheric mantle and asthenosphere. The SCLM is defined as the uppermost layer of 427
the mantle which is not involved in convection; it is therefore capable of preserving its 428
distinct geochemical signature for a long time. The SCLM also has a better potential of 429
transferring its geochemical heterogeneities during plate tectonic processes (i.e., Harry 430
and Leeman, 1995). There is evidence for the existence of even 1 Ga old subduction 431
enriched mantle source of kimberlites under western Australia (McCulloch et al., 1983) 432
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and similar Proterozoic-Mezozoic source of calc-alkaline rocks in the Basin and Range 433
province (USA) (Fitton et al., 1991; Hawkesworth et al., 1995). In contrast, the 434
asthenosphere is convective which makes it less likely to preserve its geochemical 435
heterogeneities for a long period of time. This is due to its continuous deformation 436
(convection) and higher temperature relative to SCLM. 437
There are basically three ways to trigger melting in the subduction modified 438
mantle after the cessation of subduction: 439
1) Continental extension arguably represents the simplest way in which the SCLM 440
may cross the solidus. Melting in this scenario is caused by decompression associated 441
with extension. Hawkesworth and Gallhager (1993) studied the relationship among 442
initial lithosphereic thickness, stretching factor and temperature, in order to find out if 443
melting can occur within the SCLM or asthenosphere or in both (assuming that 0.4 wt% 444
of H2O is present). The authors found that melting preferentially occurs in the SCLM if 445
the degree of extension is small, while the SCLM is thick. 446
2) Plumes can trigger melting too because they disturb the thermal structure and 447
raise the temperature above the solidus. 448
3 Slab delamination can place the lower part of the SCLM beneath thick 449
lithosphere, due to its negative buoyancy. As the delaminated slab sinks, its temperature 450
exceeds the solidus (as it is demonstrated in Anatolia, (Turkey) by Keskin et al. (1998) 451
and Pearce et al. (1990)) 452
It has also been realized that subduction-related signal usually vanishes with time 453
after the cessation of subduction or commencement of volcanism in intra plate settings. 454
This is clearly reflected in decreasing LILE/HFSE ratios and increasing HFSE 455
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concentrations with time (i.e., Almandaz et al., 2000; El Bakkali et al. 1998; Fitton et al. 456
1991; Hawkesworth et al., 1995). This implies that the original enrichment is used up by 457
multiple melting events until the state where the mantle becomes depleted and, 458
henceforth, loses its potential to yield further melt batches. Such phenomena was 459
demonstrated by Till et al. (2007) for the southern Sonora volcanites (Mexico), where 460
volcanism following the cessation of subduction (after a short quiescence) displayed 461
decreasing subduction signal and eventually terminated after a 4 Ma of activity. Another 462
noteworthy observation is that OIB-like mafic alkaline volcanism often follows the 463
paroxysm of volcanism in the mature phase of extension. This can be explained by 464
decreasing degree of melting in the deeper enriched asthenosphere, rather than in the 465
lithosphere (SCLM) (El Bakkali et al., 1998; Fittin et al., 1991). It was suggested by 466
Hawkesworth and Gallagher (1993) that plume and extension related volcanism of 467
enriched lithosphere can be distuinguished by their volume and temporal evolution. This 468
is because extension-related volcanism is less voluminous and subduction signature turns 469
into OIB character with time. By comparison, plume volcanism is quiet extensive and 470
generally starts with OIB character. 471
472
7. Discussion 473
474
7.1. Subduction-related geochemistry, crustal contamination and underplating 475
476
Studies of the petrology and geochemistry of the middle Miocene volcanic rocks 477
of the “Western segment” have revealed a clear subduction character (Fig 3 and 4). 478
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Their mantle source appears to be heterogeneous, showing evidence for predominantly 479
enriched mantle components, with EMI dominating over EMII component. The trace 480
element and isotopic compositions of EMI have been interpreted in terms of mixing of 481
various components such as: pelagic and/or terrigenous sediments, sub-continental 482
lithospheric mantle, ancient oceanic plateau, oceanic crust and lower continental crust 483
(Lustrino and Dallai, 2003). The geochemical character of the middle Miocene 484
volcanites in the “Western segment” was interpreted as being due to subduction-related 485
diapiric uprise of a large-volume of partial melts from the metasomatised lithospheric 486
mantle and asthenosphere. These partial melts were thought to have caused underplating 487
and crustal anatexis (i.e., Harangi, 2001b; Harangi and Lenkey, 2007; Seghedi et al., 488
2005). Crustal contamination of the volcanites is indicated by 87Sr/86Sr vs. 206Pb/204Pb 489
isotope sytematics (Harangi, 2001b) and the clear correlation between 87Sr/86Sr and 490
Th/La ratios. Such isotopic and geochemical characteristics imply that crustal 491
contamination was more significant than aqueous fluids from a subducted slab during the 492
genesis of middle Miocene volcanites (Harangi and Lenkey, 2007). The mixing between 493
crustal and mantle-derived melts led to the formation of middle Miocene volcanites 494
(Harangi, 2001b; Harangi and Lenkey, 2007; Seghedi et al., 2004b; Seghedi et al., 2005). 495
Oxygen isotopic compositions of the middle Miocene volcanites also refer to either 496
source contamination or mixing of mantle-derived magmas with lower crustal rocks 497
(Harangi et al., 2001; Harangi et al., 2006; Seghedi et al., 2004a). Melting in the 498
subduction-enriched mantle, the melting temperature of which was lowered by the 499
presence of volatiles (Green, 1973a; Green, 1973b), was triggered by the first phase of 500
extension in the middle Miocene (17-14 Ma ago). Although the tectonic processes (i.e., 501
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roll-back) driving the extension are well constrained (i.e., Horváth, 1993), the subduction 502
which may have caused the enrichment in the mantle still remains matter of intense 503
debate. For the Central Slovakian Volcanic Field (Fig.1) it is thought that delamination 504
of part of the lithospheric mantle played a further role in shaping the temporal 505
distribution and geochemistry of volcanism (Seghedi et al., 1998; Seghedi et al., 2004a). 506
As discussed earlier, underplating is well-documented beneath both BBH and NG. 507
Such underplating may be responsible for the granulite-facies metamorphism in the lower 508
crust (both localities are in the vicinity of the middle Miocene volcanites of the “Western 509
segment”, Fig. 1). Furthermore, underplating could have supplied heat for lower crustal 510
melting and played a role in the generation of middle Miocene intermediate and silicic 511
magmas. It is plausible that the large underplated bodies may represent basaltic melts 512
trapped at MOHO levels during the first extension phase in the middle Miocene. The 513
exact timing of these processes is, however, not yet known. The relationship between 514
melt underplating and tectonics is, therefore, somewhat unclear. 515
516
7.2. Potential subductions being responsible for source enrichment 517
518
The most appropriate candidate responsible for metasomatism beneath the CPR is 519
the former Penninic-Magura Ocean along the Carpathians (Fig. 5, Downes et al., 1992; 520
Downes et al., 1995; Szabó et al., 1992). There may have been intervening, smaller 521
continental blocks within the Penninic-Magura Ocean (e.g. Czorsztyn ridge). This was a 522
Mesozoic ocean of possible Jurassic age, was partially closed in the late Cretaceous 523
(Birkenmajer, 1986) and/or Paleogene-Miocene (Kovać et al., 1994; Meulenkamp et al., 524
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1996). Evidence for subduction of this ocean can be found in a wide accretionary prism 525
composed of turbidites, i.e., the external Carpathians. The oceanic nature of the 526
basement of these “Flysch nappes” is doubted by several authors (Winkler and Slaczka, 527
1992). The facts that both this basement disappeared entirely or easily gave place for the 528
overriding Alcapa and Tisza-Dacia microplates during their Paleogene-Miocene 529
movements, suggest that it could have an oceanic basement. The final collision occurred 530
diachronously along the Carpathian arc: early Miocene in the west, middle Miocene in 531
the north-northeast, late Miocene in the east of the arc (Nemcok et al., 1998). Although 532
subduction of the external part of the Magura Ocean might have started in the late Eocene 533
(Oszczypko, 1992), the downgoing slab could barely have reached magma-generating 534
depths by the middle Miocene. This is also supported by Gerya et al. (2002) who stated 535
that at least 24 Ma is necessary for creating extensive metasomatism of a mantle wedge 536
above an active subduction zone. Similarly, Arcay et al. (2006) and Arcay et al. (2005) 537
showed that at least 15 Ma is needed to create substantial hydration and erosion in the 538
mantle wedge (considering various convergent rates). There is also a spatial problem at 539
the time of intense middle Miocene igneous activity, as the subducted slab was too far 540
north (~300 km) from Alcapa (in its reconstructed position) to account for the intense 541
mantle metasomatism (Fig. 5; Kovács et al., 2007). The existence of well-developed 542
subduction beneath the entire Western Carpathians is now also doubted by several 543
geophysical studies as discussed above (i.e., Grad et al., 2006; Szafián and Horváth, 544
2006; Szafián et al., 1997; Szafián et al., 1999; Wortel and Spakman, 2000). An 545
alternative hypothesis has been proposed, where the lithospheric structure beneath the 546
(the western portion of) Western Carpathians can be explained either by transcurrent 547
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motion of nappes along the orogen without major convergence of the subcrustal 548
lithosphere (Szafián et al., 1997) or a collisional model implying a “crocodile mouth”-549
like structure (Grad et al., 2006). There are multiple models that suggest that subduction 550
along the Western Carpathian arc explains inadequately the metasomatised nature of the 551
mantle beneath the “Western segment”. 552
An alternative possibility for the mantle metasomatism at the origin of these 553
magmas is that it was caused by an earlier, pre-Miocene subduction outside the present 554
CPR. Kovács et al. (2007) showed that there is a geochemical and geodynamic link 555
between the Paleogene and Early Miocene subduction related volcanics of the wider 556
CPR, including the eastern segment of the Alps and the Dinarides. Kovács et al. (2007) 557
also proposed that the Paleogene-early Miocene igneous rocks of the Alpine-Carpathian-558
Pannonian-Dinaric region formed a uniform volcanic arc mainly along the trench of the 559
Budva-Pindos Ocean (Fig. 5). Mantle tomography revealed the remnant of this former 560
subducted slab along the eastern Adriatic shoreline, which is the only well documented 561
slab remnant in the wider CPR (Koulakov et al., 2002; Lippitsch et al., 2003). Middle 562
Miocene rocks in the “Western segment” largely overlap with the occurrence of the 563
Paleogene-early Miocene rocks of the MHZ and share common geochemical characters 564
(Fig. 3 and 4). Therefore, we conclude that the Budva-Pindos (and Vardar) subductions 565
are likely the wanted candidates for explaining source enrichment of middle Miocene 566
volcanites. 567
568
7.3. A model for the “inherited” subduction-related character 569
570
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Kovács et al. (2007) hypothesized that the enriched lithospheric mantle of the 571
Alcapa and Tisza blocks could have been transferred from the former supra-subduction 572
environment to the CPR during extrusion and block rotations in the late Oligocene-early 573
Miocene. This inherited enrichment in the source implies that the uppermost mantle 574
moves together with the crust during major block rotations, and should preserve its 575
metasomatic components from a previous geodynamic setting. Accordingly, Falus 576
(2004) and Hidas et al. (2007) suggested that the Alcapa unit and possibly the Tisza block 577
(Fig. 1) moved together with a portion of their lithospheric mantle on the basis of 578
deformation analysis of mantle xenoliths. They found that shallower (“older”) sections of 579
the upper mantle beneath the BBH suffered multiple deformation events, whereas the 580
deeper sections (“younger”) show only one strong deformation event most probably 581
related to basin formation. In this scenario the shallower part represents lithospheric 582
mantle which could have been displaced with the crust during the extrusion of Alcapa 583
and Tisza. Although it is still a matter of debate whether the extrusion happened at 584
crustal or lithospheric scale in the CPR, the available data suggest that this process likely 585
happened at lithospheric scale. If so, the question arises at what depth did the decoupling 586
happen in the lithospheric units close to the Budva-Pindos subduction? At this depth the 587
rheology is controlled by amphibole stability (Niida and Green, 1999; Wallace and 588
Green, 1991) because amphibole decomposition leads to a rheological low-point in the 589
mantle where the decoupling is most likely to occur. This depth corresponds to ~ 90 km, 590
which is excellent agreement with the 100 km depth of lithosphere-asthenosphere 591
boundary under Phanerozoic continental areas (Plomerova et al., 2002). This also 592
matches calculation of Falus et al. (2000) who suggested on the basis of mantle xenoliths 593
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from the Styrian basin (SB), that the maximum thickness of the lithosphere was 90-120 594
km before the extension events. Rheological considerations and modeling of 595
Willingshofer and Cloetingh (2003) and Willingshofer et al. (1999) also confirm that 596
models show better agreement with field observation in the Eastern Alps if coupling is 597
stronger between the crust and mantle. 598
Thus, it is likely that the Alcapa block was extruded from the Alpine realm as a 599
lithosperic unit in the late Oligicene-early Miocene. This is also indicated by the 600
preservation of orthopyroxene-rich mantle rocks, subduction-related volatile-rich melts 601
trapped as inclusions in peridotites and the widespread modal metasomatism (i.e., the 602
presence of amphibole and phlogopite) in mantle xenoliths from the BBH and NG. This 603
suggests that the uppermost mantle must have been in or close to a supra-subduction 604
setting. Stable isotopic studies of Demény et al. (2004) and Demény et al. (2005) also 605
support this hypothesis, as some amphibole megacrysts from the mantle show clear 606
results of subduction-related metasomatism beneath the BBH and NG. A recent study of 607
xenoliths from the Styrian basin by Coltorti et al. (2007) (see Fig. 1) also shows evidence 608
for metasomatism by slab-derived melts. In addition, granulite xenoliths beneath the 609
BBH (and NG) may represent remnants of a subducted oceanic lithosphere that were 610
emplaced at lower crustal levels (Dobosi et al., 2003; Embey-Isztin et al., 2003). 611
Besides the SLCM there may have been contribution from the asthenosphere as 612
well to middle Miocene magmas. It is plausible to assume that not only the lithosphereic 613
mantle but also the less viscous, convective asthenosphere may have suffered eastward 614
directed extrusion from the Alpine realm most probably during the late Oligocene-early 615
Miocene. This flow may have happened independently from the lithosphere, and was 616
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parallel to the Alpine orogen (W-E) as it was observed for other orogens with recent 617
convergence (i.e., Himalaya; Flower et al., 2001; Meissner et al., 2002). In this scenario 618
the asthenospheric flow may have played an important role in opening up the Pannonian 619
Basin (and Western Mediterranean basins) and driving subduction roll-back along the 620
Carpathians as well. It is hypothesized; therefore, that he asthenosphere may also have 621
transferred its enriched geochemical signature from the Alpine collision belt to the 622
marginal basins. Unfortunately, seismic anisotropy measurements are not yet available 623
which may give a clearer picture about the potential existence of an eastward directed, 624
orogen parallel anisotropy in the asthenospheric mantle beneath the CPR. However, 625
there may be mantle xenoliths of possible asthenospheric origin that show signs only one 626
deformation event, primitive geochemistry and deeper origin consistent with this model 627
(Falus et al., 2000; Falus, 2004). This model explains OIB character of younger alkaline 628
basalts of the region which are mainly of clear asthenospheric origin. 629
In this study we propose a similar scenario to that of Kovács et al. (2007) for the 630
formation of the middle Miocene volcanics of the “Western segment” along the MHZ. It 631
is proposed, therefore, that the source of the middle Miocene had been metasomatised 632
along the subduction of either the Budva-Pindos or Vardar Ocean in the Late Mesozoic 633
and Paleogene and then it was transported to the CPR during the eastward extrusion and 634
rotation of Alcapa and Tisza. The available geological record is not yet appropriate to 635
resolve and distinguish the exact contribution of the two proposed oceanic realms to the 636
source enrichment. Furthermore, the exact history of both the Budva-Pindos and Vardar 637
Ocean is still controversial (i.e., Csontos and Vörös, 2004; Pamić et al., 2002; 638
Rosenbaum and Lister, 2005). Formation of middle Miocene igneous rocks was 639
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triggered by the first phase of extension between 17-14 Ma causing extensive melting in 640
the previously enriched (metasomatised) mantle (Harangi and Lenkey, 2007; Seghedi et 641
al., 2004a). The middle Miocene volcanics are aligned roughly parallel to the MHZ, 642
which suggests that this highly fractured and deformed zone may have played a 643
significant role in the genesis and tapping of the melt formed in the mantle (Csontos and 644
Nagymarosy, 1998; Kovács et al., 2007). Basaltic volcanism of this period has not yet 645
been identified on the surface; however large bodies of underplated mafic magmas have 646
been recognized beneath the BBH and NG in pyroxenite xenoliths (Embey-Isztin et al., 647
1990; Kovács et al., 2004; Zajacz et al., 2006) and by geophysical studies (Guterch et al., 648
2003). These bodies may represent the trapped basaltic melts, which caused anatexis and 649
granulite-facies metamorphism of the lower crust, eventually leading to the formation of 650
middle Miocene mainly calc-alkaline volcanics (i.e., Harangi, 2001b; Harangi and 651
Lenkey, 2007; Kovács and Szabó, 2005; Seghedi et al., 2005). 652
Seghedi et al. (2005) pointed out that the geochemical and isotopic data confirm 653
the diminution of subduction-modified mantle with time, in favor of less affected 654
asthenospheric OIB-like (EMI) mantle in the “Western segment”. Geochemistry of the 655
older Paleogene-early Miocene volcanics of the “Western segment” provide evidence for 656
a stronger enrichment in Th, LIL elements (Fig. 4) which are typical indicators of 657
sediment and fluid contribution from a subducted slab. On the other hand, the youngest 658
Plio-Pleistocene alkali basalts of the “Western segment” (Fig. 1, i.e., LHP, BBH, SB, 659
NG) display only a slight subduction signature with a strong OIB character (Dobosi et al., 660
1998; Gméling et al., 2007; Harangi, 2001b; Harangi and Lenkey, 2007; Seghedi et al., 661
2004b), whereas middle Miocene volcanites are of transitional character. This declining 662
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subduction signature with time and OIB character of the younger alkaline basalts with 663
their relatively moderate volume exclude plume(s) as potential source of volcanism as it 664
is proposed by Hawkesworth and Gallhager (1993). The OIB character of younger late 665
Miocene to subrecent volcanics, as we demonstrated above, may have been accounted for 666
by eastward-directed, and enriched asthenospheric flow from the Alpine collision belt. In 667
contrast, older middle Miocene volcanites resemble geochemical patterns of magmas 668
formed in intra-plate setting various time after cessation of subduction worldwide (see 669
overview on the inherited enriched mantle source). This implies that extensive 670
metasomatism gained in a subduction setting and, therefore, the potential for melting in 671
the mantle does not last forever, but it vanishes with the increasing degree of melt 672
extraction owing to active plate tectonic processes. In a simple analogy the mantle works 673
like a “rechargeable battery” which is filled up with “melting potential” at subduction 674
zones and than this potential is used up by subsequent episodic melting events. “Melting 675
potential” refers to the depression of the solidus relative to the “dry” mantle, which is 676
dictated by the amount of available volatiles as melt and fluid inclusions, hydrous phases 677
and H in nominally anhydrous minerals (i.e., Gaetani and Grove, 2003; Green, 1971; 678
Green and Liebermann, 1976; Green and Falloon, 2005; Niida and Green, 1999). 679
This “inherited enrichment” model has the advantage of not raising any space-680
time problem or requiring contemporaneous subduction. It is also not excluded, bearing 681
in mind the long and adventurous tectonic history of the CPR, that even earlier 682
subduction events with pre-Mesozoic age could have left their fingerprints on the mantle 683
beneath the region (Dobosi et al., 1998; Downes et al., 1992; Embey-Isztin et al., 1993a; 684
Rosenbaum et al., 1997). Given the “inherited enrichment” model, a contemporaneous 685
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(middle Miocene) subduction is not crucial to explain the enrichment in the source region 686
of the middle Miocene volcanic rocks of the “Western segment”. Subduction along 687
Eastern Carpathians could only influence the magma generation by its roll-back effect, 688
triggering significant thinning of the lithosphere. 689
Some previous studies in the CPR (Harangi et al., 2006; Seghedi et al., 2005) also 690
raised the possibility of “inherited” source enrichment for Tertiary volcanics of the 691
“Western segment”. They also proposed that this source was remelted during 692
athenospheric upwelling, however, they did not attempt to link the proposed inherited 693
enrichment to any former subduction in the CPR. As we showed earlier this “inherited 694
enrichment” model is not unknown, however, its detailed application to the middle 695
Miocene volcanics of the MHZ integrating geophysics, xenoliths and igneous rocks is a 696
new effort. 697
698
8. Conclusions 699
700
Middle Miocene volcanic rocks in the “Western segment” of the CPR show 701
subduction-related geochemistry, in common with the Paleogene-early Miocene rocks of 702
the same area. Magma generation in the middle Miocene was triggered by 703
contemporaneous extension; however the subduction responsible for the source 704
enrichment has been the subject of intense debate. We propose that the subduction-705
related geochemical character is an inherited feature, where the metasomatism happened 706
before the Miocene along the subduction of either the Budva-Pindos or Vardar Ocean. 707
The major continental blocks (i.e, Tisza and Alcapa) together with their metasomatised 708
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mantle lithosphere were transported subsequently to the Carpathian embayment during 709
the late Oligocene-early Miocene. Considerable thinning of the lithosphere happened 710
during the first extension phase in the middle Miocene, which induced extensive melting 711
in the already metasomatised mantle. The produced mafic melts were trapped in the 712
lower crust, resulting in massive underplating. The heat flux associated with this 713
underplating process led to granulite-facies metamorphism and anatexis of the lower crust 714
as being represented by partially molten granulite and pyroxenite xenoliths from the 715
“Western segment”. Wide geochemical variations in middle Miocene magmas are due to 716
source heterogeneity and various ratios of mantle and lower crustal melts involved in 717
magma genesis. The inherited enrichment in the source is supported by the diminishing 718
subduction-related geochemistry of the magmas with time. Accordingly, Paleogene arc 719
magmas display very clear subduction overprint on their geochemistry. Subduction 720
character is weaker for the middle Miocene volcanics of the “Western segment” showing 721
only little traces of subduction-induced enrichment in the source. This model also 722
suggests that subduction, which may have been able to supply sediments and volatiles 723
down to the mantle beneath the “Western segment”, may not have developed along the 724
entire Western Carpathians. 725
726
Acknowledgements 727
728
L. Csontos is especially acknowledged for the discussions which outlined the 729
geodynamic framework of this paper, without his help this paper would only have 730
remained an idea. The authors would like to thank H. Downes and N. Tailby for their 731
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valuable suggestion on an earlier version of the manuscript. Sz. Harangi and M. Lustrino 732
is acknowledged for sharing their manuscripts with us. We express our gratitude to I. 733
Seghedi, Sz. Harangi, A. Szakács, O. Vaselli, G. Bada, M. Kázmér, M. Grad and M. 734
Beltrando for the fruitful discussions on the evolution of the CPR and the wider Alpine 735
realm. Gy. Falus, E Bali and Z. Zajacz from the Lithosphere Fluid Research Lab are 736
particularly acknowledged for their continuous support during the past years. Two 737
anonymous reviewers are also acknowledged for their constructive criticism that helped 738
to improve our manuscript. The manuscript also benefited form the thorough editorial 739
handling of R. Stephenson. This work is partially supported by an A. E. Ringwood 740
Memorial Scholarship and an Australian International Postgraduate Research Scholarship 741
to I. Kovács. This is publication 36th in the series of the Lithosphere Fluid Research 742
Laboratory at Eötvös University, Budapest, Hungary. 743
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Figure caption 744
745
Fig. 1. Major tectonic units and spatial distribution of middle Miocene to subrecent 746
igneous rocks of the wider Carpathian-Pannonian region. The map is modified after Fig. 747
1 in Kovács et al. (2007). Volcanic provinces as “Western” and “Eastern segments” of 748
the middle Miocene igneous rocks (see text for more details) and localities of some 749
Paleogene-early Miocene rocks are also indicated. Plio-Pleistocene alkali basalts with 750
mantle and lower custal xenoliths with which the study is concerned are also highlighted 751
(BBH – Bakony-Balaton highland, LHP – Little Hungarian Plain volcanic field, NG – 752
Nógrád-Gömör volcanic field, SB – Styrian basin, ETB – East Transylvanian basin). 753
Abbreviation for the middle Miocene igneous occurrences of the “Western segment”: B = 754
Börzsöny, Bü = Bükk, C = Central Slovakian Volcanic Field, Cs = Cserhát, M = Mátra, 755
SB = Styrian Basin, SBP = South Pannonian Basin, T = Tokaj, V = Visegrád Mts. 756
Abbreviations for the major faults are: PAF-Periadriatic fault, BF-Balaton fault, MHF-757
Mid-Hungarian fault, SF-Sava fault, MBF-Main Balkan fault, VF-Vardar fault. 758
759
Fig. 2. Total alkalis vs. silica diagram for middle Miocene volcanic rocks of the 760
“Western segment” (Le Bas et al., 1986). Data are taken from references in Table 1. 761
762
Fig. 3. a) Chondrite normalized (Nakamura, 1974) rare earth element pattern of the 763
middle Miocene igneous rocks of the “Western segment”; b) Primitive mantle normalized 764
(McDonough and Sun, 1988) trace element pattern of the middle Miocene igneous rocks 765
of the”Western segment” (data are taken for a) and b) from references in Table 1); c) 766
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Chondrite normalized (Nakamura, 1974) rare earth element pattern of the Paleogene-767
early Miocene igneous rocks of the Middle Hungarian zone (see Fig. 1 for the location of 768
the Mid-Hungarian zone); d) Primitive mantle normalized (McDonough and Sun, 1988) 769
trace element pattern of the Paleogene-early Miocene igneous rocks of the Mid-770
Hungarian zone (data are taken for c) and d) from Kovács et al., 2007 and references 771
therein). 772
773
Fig. 4. a) Ba/Zr vs. Nb/Zr diagram for the middle Miocene igneous rocks of the “Western 774
segment”. Diagram is modified after Pereplov et al. (2005) and data are from references 775
in Table 1.; b) Ba/Zr vs. Nb/Zr diagram for the Paleogene-early Miocene igneous rocks 776
of the Mid-Hungarian zone. Diagram is modified after Pereplov et al. (2005) and data are 777
from Kovács et al. (2007) and references therein; c) Nb/Y vs. Th/Y diagram for the 778
middle Miocene igneous rock of the “Western segment”. Diagram is modified after 779
Seghedi et al. (2005) and data are from references in Table 1; d) Nb/Y vs. Th/Y diagram 780
for the Paleogene-early Miocene igneous rock of the Mid-Hungarian zone. Diagram is 781
modified after Seghedi et al. (2005) and data are from Kovács et al. (2007) and references 782
therein; e) 86Sr/87Sr vs. 143Nd/144Nd diagram for middle Miocene igneous rocks of the 783
“Western segment”. Data are from references in Table 1.; f) 86Sr/87Sr vs. 143Nd/144Nd 784
diagram for Paleogene-early Miocene igneous rocks of Mid-Hungarian zone. Data are 785
from Kovács et al. (2007) and references therein. 786
Fields for different kinds of Ocean Island Basalts (OIB) are from Dupuy et al. 787
(1988), Palacz and Saunders (1986) and Weaver (1991) for HIMU, Humphris and 788
Thompson (1983), Storey et al. (1988) and Weaver (1991) for EM1 and Dostal et al. 789
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(1982) and Palacz and Saunders (1986) for EM2, respectively. Compositions of enriched 790
and normal middle ocean ridge basalt (E-MORB and N-MORB) are from McDonough 791
and Sun (1988). 792
793
Fig. 5. Shematic summary of major tectonic events in the Carpathian-Pannonian region 794
since the late Eocene from a “mantle perspective”. Main tectonic features as subduction 795
zones, sutures and faults are indicated on the map. Abbreviations are the followings for 796
tectonic units: EA – Eastern Alps, SA – Southern Alps, Ti – Tisza, Da – Dacia, Di – 797
Dinarides, Ap – Apulian indenter; for tectonic lines: PAF – Periadriatic fault, BF – 798
Balaton fault, MHF – Middle-Hungarian fault, ZZF – Zagreb–Zemplén fault, VF – 799
Vardar fault. Please note that ZZF and MHF is basically the same fault, the northern part 800
of which is called as the MHF. Rotations of the Tisza-Dacia block are only schematic; 801
see Patrascu et al. (1994) for more details. 802
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1287
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Table 1. Location, age and petrology of the Middle to Late Miocene igneous rocks in the vicinity of the Middle Hungarian zone
Reference Locality* Age Petrology Póka et al., 1988 V, B, Cs, M, C, T Karpathian-Late Sarmatian rhyolite, andesite, basalt, dacite Salters et al., 1988 T, B, M, Cs Karpathian-Late Sarmatian basaltic andesite-rhyolite Szabó et al., 1992 B, Bü, C, Cs, M, T, V Karpathian- Sarmatian rhyolite, andesite, basalt, dacite Pamic et al., 1995 SPB Karpathian-Late Badenian trachyandesites, basalts, andesites Downes et al., 1995 B, M, T, Cs Karpathian-Late Sarmatian andesite, dacite Harangi et al., 1995 LHP Late Miocene trachybasalt, trachyandesite, trachyte Harangi, 2001a SB Badenian trachyandesite, trachyte Harangi et al., 2001a B, C, V Badenian andesite, dacite, rhyodacite Seghedi et al., 2004 Bü, T Karpathian-Sarmatian rhyolite, dacite
Póka et al., 2004 Cs Karpathian-Early Sarmatian andesite, dacite, rhyodacite, dacite
Zelenka et al., 2004 buried under Miocene sediments along the MHZ Karpathian-Sarmatian andesite, dacite, rhyodacite, dacite,
basalt Gméling et al., 2005 C, B, V Badenian basaltic andesite-andesite Harangi et al., 2005 Bü Karpathian-Sarmatian rhyolite** * B = Börzsöny, Bü = Bükk, C = Central Slovakian Volcanic Field, Cs = Cserhát, LHP = Little Hungarian Plain, M = Mátra, SB = Styrian Basin, SBP = South Pannonian Basin, T = Tokaj, V = Visegrádi Mts. ** glass shards and not bulk-rock analyses
Table
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Southern Alps
Tauern
PAF
Eastern Alps
Eastern Carpathians
Southern Carpathians
Serb
o-M
acedonian
massif
Dinarides
TISZA
DACIA
ALCAPA
Mid-Hun
garia
n zon
e
100 km
Moesian Platform
Bohemian Promontory
Tran
sdan
ubian
Ran
ge
Apuseni Mts.
Mecsek Mts.
BF
DF
SF
Enga
dine
Middle Miocene-subrecent magmatic rocksexposed/subcrop
Plio-Pleistocene alkali basalts
xenoliths described in the text
Paleogene magmatite, nameeffusive/intrusive
Early Miocene magmatiteexposed/subcrop
Main faults, nameexposed, subcrop
PAFVF
MBF
MHF
Western Carpathians
LHPB
Kovács and Szabó, Fig. 1 (.ai)
Villány Mts.
European foreland
Alpine-Carpathian foredeep
Alpine-Carpathian flysch belt
Penninic ophiolites
ALCAPA units
TISZA-DACIA units, Serbo-MacedonianSouthern Alpine, Dinaric, Adriatic platform units
Bosnian flysch, Slovenian trough
Mesozoic mélange and ophiolites
BBH
C
CsNG
MBü
T
V
SB
SPBDrava Basin.
ETB
Nyirseg
“Eastern segment”“Eastern segment”
“Western segment”“Western segment” Magura subduction
Penninic subduction
Budva-Pindos subduction
Vardar suture
Sava-Vardar zone
subductions, sutures
Figure
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35 40 45 50 55 60 65 70 750
2
4
6
8
10
12
14
16
SiO2
Picro-basalt
Basalt
Basalticandesite
Andesite
Dacite
Rhyolite
Trachyte
Trachydacite
Trachy-andesite
Basaltictrachy-andesite
Trachy-basalt
Tephrite
Basanite
Phono-Tephrite
Tephri-phonolite
Phonolite
FoiditeK
O+
Na
O2
2
Kovács and Szabó, Fig. 2 (.cdr)
Figure
Page 51 of 54
Acce
pted
Man
uscr
ipt
1
10
100
1000
La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu.1
1
10
100
1000
10000
Cs Rb Ba Th U Nb K La Ce Pb Pr Sr P Nd ZrSm Eu Ti Dy Y Yb Lu
Mid-Hungarian zone
.1
1
10
100
1000
10000
1
10
100
1000
Mid-Hungarian zone
La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Cs Rb Ba Th U Nb K La Ce Pb Pr Sr P Nd ZrSm Eu Ti Dy Y Yb Lu
Middle Miocene igneous rocks (SiO <56wt%)2
Kovács and Szabó, Fig. 3 (.cdr)
Rock
/Chondri
teR
ock
/Chondri
te
Rock
/Pri
mit
ive
man
tle
Rock
/Pri
mit
ive
man
tle
Middle Miocene igneous rocks (SiO >56wt%)2 Paleogene-early Miocene igneous rocksof the Mid-Hungarian Zone
a b
dc
Figure
Page 52 of 54
Acce
pted
Man
uscr
ipt0111.0
10.0
1.0
1N
b/Z
r
rZ/aB
N-MORB
E-MORB
MORB & Within Plate Basalts
IAB
0111.010.0
1.0
1
Nb/
Zr
rZ/aB
N-MORB
E-MORB
MORB & Within Plate Basalts
IAB
0111.010.0
1.0
1
01
Th/
Y
Y/bN
E-MORB
0111.010.0
1.0
1
01
Th/
Y
Y/bN
E-MORB
0.702 0.704 0.706 0.708 0.710 0.712 0.7140.5120
0.5122
0.5124
0.5126
0.5128
0.5130
0.5132
87 86Sr/ Sr
441341
dN
/d
N
N-MORB
0.702 0.704 0.706 0.708 0.712 0.7140.5120
0.5122
0.5124
0.5126
0.5128
0.5130
0.5132
87 86Sr/ Sr
N-MORB
0.710
Kovács and Szabó, Fig. 4 (.ai)
Middle Miocene (SiO2<56 wt%)Middle Miocene (SiO2>56 wt%)Paleogene-early Miocene (SiO2<56 wt%)Paleogene-early Miocene (SiO2>56 wt%)
EMII
HIMU
EMI
a b
c d
e f
dN
/d
N441
341
Figure
Page 53 of 54
Acce
pted
Man
uscr
ipt
late Eocene - early Oligocene (~30 Ma) late Oligocene - early Miocene (~20 Ma)
middle Miocene (~17 Ma) middle to late Miocene (~12 Ma)
PAF
EA
SA
Di
Ap
Ti
Da
PAF
EA
SA
Di
Ti
Da
BF
ZZF
EA
SA
Di
Ti
Da
Ap
BF
?
PAF
ZZF
EA
SA
Di
Ti
Da
Ap
BF
PAFZZF
Subduction
Suture zone
Fault zones
Proposed fault zones
Rotations
Direction of block movements
volcanic activity
Legend
Kovács and Szabó, Fig. 5 (.cdr)
Alcapa
Alc
apa
Alc
apa
Alc
apa
a b
c d
Budva
-Pin
dos
subductio
n
Magurasubduction
Budva
-Pin
dos
sutu
re
Ap
Magura
subduction
"Vard
ar
sutu
re"
Magura
subductionMHF
VF VF
VF VF
Figure
Page 54 of 54