Evolution of Mafic Alkaline Melts Crystallized
in the Uppermost Lithospheric Mantle: a Melt
Inclusion Study of Olivine-Clinopyroxenite
Xenoliths, Northern Hungary
ZOLTA¤N ZAJACZ
1*, ISTVA¤N KOVA¤CS
2, CSABA SZABO¤
3,
WERNER HALTER
1AND THOMAS PETTKE
41
DEPARTMENT OF EARTH SCIENCES, INSTITUTE OF ISOTOPE GEOCHEMISTRY AND MINERAL RESOURCES, ETH ZU«RICH, 8092 ZU«RICH, SWITZERLAND
2
RESEARCH SCHOOL OF EARTH SCIENCES, BUILDING 61 MILLS ROAD, THE AUSTRALIAN NATIONAL UNIVERSITY, CANBERRA, ACT 0200, AUSTRALIA
3LITHOSPHERE FLUID RESEARCH LABORATORY, INSTITUTE OF GEOGRAPHY AND EARTH SCIENCES, EO«TVO«S
UNIVERSITY BUDAPEST (ELTE), PAZMANY SETANY 1/c, BUDAPEST, HUNGARY, H-1117
4UNIVERSITY OF BERN, INSTITUTE OF GEOLOGICAL SCIENCES, BALTZERSTRASSE 1^3, CH-3012 BERN, SWITZERLAND
RECEIVED APRIL 7, 2006; ACCEPTED FEBRUARY 2, 2007 ADVANCE ACCESS PUBLICATION APRIL 16, 2007
Olivine-clinopyroxenite xenoliths exhumed in alkali basalts (sensu lato) in the No¤gra¤d^Go«mo«r Volcanic Field (NGVF), northern Hungary, contain abundant silicate melt inclusions. Geothermobarometric calculations indicate that these xenoliths crystallized as cumulates in the upper mantle near the Moho. These cumulate xenoliths are considered to represent a period of Moho underplating by mafic alkaline magmas prior to the onset of Late Tertiary alkaline volcanism in the Carpathian^Pannonian region. The major and trace element compositions of silicate melt inclusions in olivine display an evolutionary trend characterized by a strong decrease in CaO/Al2O3. The parental melt of the cumulates was a
basanite formed by low-degree (2%) partial melting of a garnet peridotite source. The compositional trend of the silicate melt inclu-sions, textural features, and modelling with pMELTS show that the parental melt evolved by major clinopyroxene and minor olivine crys-tallization followed by the appearance of amphibole simultaneously with significant resorption of the earlier clinopyroxene and olivine. The resulting residual melt was highly enriched in Al2O3, alkalis
and most incompatible trace elements. This type of melt is likely to infiltrate and react with surrounding mantle peridotite as a metaso-matic agent. It might also form high-pressure pegmatite-like bodies
in the mantle that might be the source of the amphibole and sanidine megacrysts also found in the alkali basalts of the NGVF. Preferential remelting of the later-formed (i.e. lower temperature) mineral assemblage (amphibole, sanidine, residual glass) might have signif-icantly contaminated the host alkaline mafic lavas, increasing their Al2O3and total alkali contents and, therefore, reducing their MgO,
FeO and CaO content.
KEY WORDS: silicate melt inclusions; geochemistry; petrogenesis; No¤gra¤d^Go«mo«r Volcanic Field; Pannonian Basin
I N T R O D U C T I O N
Olivine-clinopyroxenite and clinopyroxenite xenoliths, and clinopyroxene and amphibole megacrysts are common within the alkali basalts in the Pannonian Basin (Embey-Isztin et al., 1990, 1993; Downes et al., 1995; Dobosi & Jenner, 1999; Dobosi et al., 2003; To«ro«k et al., 2003; Kova¤cs et al., 2004) and worldwide (Frey & Prinz, 1978; Irving & Frey, 1984; Wilson & Downes, 1991; Shaw
& Eyzaguirre, 2000; Chen et al., 2001; Bondi et al., 2002; Riley & Bailey, 2003; Liu et al., 2005). In most studies they have been interpreted as high-pressure crystallization pro-ducts of melts similar in composition to the host alkaline basalts, based on thermobarometric calculations, isotopic ratios and estimated equilibrium melt compositions (Irving, 1974; Stuckless & Irving, 1976; Frey & Prinz, 1978; Shaw & Eyzaguirre, 2000; Chen et al., 2001; Bondi et al., 2002; Dobosi et al., 2003; Riley & Bailey, 2003; Downes et al., 2004; Kova¤cs et al., 2004; Shaw et al., 2005). Alternatively, the olivine-clinopyroxenite xenoliths have been interpreted as the products of metamorphic segrega-tion during anatexis of the surrounding mantle region (Chen et al., 2001). Alkali feldspar phenocrysts have also been reported in a number of alkali basalts (Frey & Prinz, 1978; Aspen et al., 1990; Riley & Bailey, 2003), but the origin of the siliceous melt from which they crystal-lized has not yet been identified. Only Irving (1974) sug-gested that they might form as near-solidus crystallization products of primitive alkaline melts in the mantle.
In the case of crystallization of basaltic melts, separation of the mafic minerals forms cumulus ultramafic rocks com-posed of clinopyroxene, olivine and usually Al-spinel; such rocks are not representative of the parental melt, thus mineral^melt equilibria are needed to constrain the geochemical character of the parental melt.
Here we study silicate melt inclusions (SMIs) from olivine-clinopyroxenite xenoliths hosted in alkali basalts of the No¤gra¤d^Go«mo«r Volcanic Field (NGVF), northern Pannonian Basin (Hungary), to directly obtain the composition of the parental melt of the xenoliths. Most previous studies of SMIs have focused on melt inclusions in phenocrysts in volcanic rocks (e.g. Nielsen et al., 1995; Kamenetsky & Clocchiatti, 1996; Kamenetsky et al., 1997; Xu et al., 1999; Norman et al., 2002; Bertagnini et al., 2003; Cervantes & Wallace, 2003; Danyushevsky et al., 2003; Maclennan et al., 2003; Schiano, 2003; Schiano et al., 2004). Here we present data on SMIs within the constituent minerals of cumulate xenoliths derived by fragmentation of a magma body that crystallized slowly near the Moho.
The major goals of this paper are fourfold: (1) to provide additional information on the genesis of Type-II (mostly clinopyroxenite and olivine-clinopyroxenite) xenoliths; (2) to obtain a better understanding of the evolution of alkali basaltic systems; (3) to test of the representative nature of olivine-hosted melt inclusions in these xenoliths; (4) to pro-vide new information on the evolution of the Carpathian^ Pannonian region. We reproduced the original composition and the chemical evolution of the parental melt based on electron microprobe analyses (EPMA) of homogenized melt inclusions and laser ablation inductively coupled plasma mass spectrometry (LA-ICPMS) analysis of unheated multiphase inclusions in olivine. Petrogenetic
models, based on the major and trace element composition of the SMIs and the rock-forming minerals, are presented and used to constrain the formation and evolution of the parental melt.
G E O L O G I C A L S E T T I N G
The No¤gra¤d^Go«mo«r Volcanic Field (NGVF) at the north-ern margin of the Pannonian Basin (Northnorth-ern Hungary^ Southern Slovakia) (Fig. 1) is mostly composed of Plio-Pleistocene alkali basalts (sensu lato). It is surrounded by the Miocene, subduction-related, Inner Carpathian calc-alkaline volcanic belt. The Pannonian Basin was formed as a result of lithospheric stretching in two major exten-sional episodes: (1) an Early to Middle Miocene passive rifting phase driven by subduction rollback (Royden et al., 1983a, 1983b; Csontos, 1995; Fodor et al., 1999); (2) a Late Miocene syn- to post-rift phase accompanied by astheno-spheric upwelling and thinning of the mantle lithosphere (Royden et al., 1983a, 1983b; Huismans et al., 2001). The alkali basalts were erupted in the post-extensional phase and sampled the lithospheric section beneath the region, transporting a large quantity of peridotite (Type-I), ultramafic cumulate (Type-II) and crustal xeno-liths to the surface (Embey-Isztin et al., 1993; Szabo¤ & Taylor, 1994; Embey-Isztin & Dobosi, 1995; Zajacz & Szabo¤, 2003; Kova¤cs et al., 2004; Seghedi et al., 2004) (Fig. 1). The xenoliths investigated in this study were collected from the Nagy-Salgo¤ volcano. The locality is a well-preserved basanite volcanic neck containing a high abundance of olivine-clinopyroxenite and wehrlite xenoliths. The host rock is Pliocene (527 Ma) in age (Szabo¤ et al., unpublished data).
P E T R O G R A P H Y O F T H E
C U M U L AT E X E N O L I T H S A N D
S I L I C AT E M E LT I N C L U S I O N S
In the NGVF several eruptive centres contain clinopyrox-ene- and olivine-rich xenoliths. In some locations amphi-bole is also typical and has been interpreted as a metasomatic alteration product of the clinopyroxene (Kova¤cs et al., 2004). The xenoliths selected for a detailed melt inclusion study are mostly olivine-clinopyroxenites, and subordinately wehrlites. Regardless of variations in the clinopyroxene to olivine ratio, the xenoliths are textur-ally similar. The volume proportion of olivine ranges from 20% to 40% (Table 1). Both the clinopyroxene and the olivine are euhedral or subhedral and their grain sizes are between 05 and 5 mm (Fig. 2a). The edges of the olivine crystals are resorbed, whereas the clinopyroxenes typically display a narrow reaction rim (Fig. 2b and c). In addition to the major rock-forming clinopyroxene and olivine, two types of interstitial phase assemblages are present: (1) a
Fig. 1. Map of the No¤gra¤d^Go«mo«r Volcanic Field. The most important alkaline basalt (sensu lato) occurrences are shown; the studied xenoliths were collected from Nagy-Salgo¤.
light-coloured assemblage composed of euhedral alkali feldspar, clinopyroxene, Ti-magnetite, apatite and residual glass (Fig. 2d and e); (2) a dark-coloured assemblage of partially decomposed amphibole displaying evidence of breakdown to rho«nite, nepheline, clinopyroxene, olivine and Ti-magnetite. In some xenoliths the amphibole is completely decomposed and only the residual minerals are present (Fig. 2f). Interstitial phase assemblage (1) is typical of all the xenoliths whereas interstitial phase assemblage (2) is observed in small quantities only in a few xenoliths (Table 1). Despite the uniform texture and mineralogy there is a significant variation in the melt and fluid inclusion content of the xenoliths. Out of five large (3^5 cm in diameter) unaltered xenoliths selected for investigation, only two contained SMIs in olivine large enough for analysis (410 mm), whereas clinopyroxenes are generally rich in SMIs and CO2inclusions (Table 1). As the
main focus of our study is on olivine-hosted SMIs, these two xenoliths (NNS07 and NNS08) have been selected for detailed analytical work.
Olivines contain SMIs either as single inclusions or in clusters. They are generally absent from the cores of large olivine crystals. The inclusions are always rounded and almost isometric (Fig. 3a^c), ranging from 8 to 40 mm in diameter. They occur as partially recrystallized multiphase assemblages of clinopyroxene (20^30 vol. %), Al-spinel (1^3 vol. %), rho«nite (10^20 vol. %) and/or Ti-rich pargasi-tic amphibole (5^15 vol. %), together with residual glass (40^60 vol. %) and a CO2 bubble (4^8 vol. %). Small
sulphide blebs are also present in several inclusions (Fig. 3c).
Clinopyroxene crystals are packed with elongated and oriented SMIs. Their size ranges from submicroscopic to 30 mm in diameter (Fig. 3d). They are usually present from grain cores to rims. SMI size distribution sometimes shows periodic zonation with alternating zones of small (submicroscopic to 5 mm in diameter) and large inclusions (from 5 mm to 30 mm in diameter) following the growth direction of the clinopyroxene. They are partially recrys-tallized and composed of clinopyroxene (0^20 vol. %), Ti-rich pargasite (0^10 vol. %), Al-spinel (1^3 vol. %), Ti-magnetite (0^2 vol. %), residual glass (65^95 vol. %) and a CO2bubble. The CO2fluid to melt ratio is variable,
suggesting accidental entrapment of a previously exsolved fluid phase along with the melt (heterogeneous entrap-ment) (Fig. 3d). Individual CO2 inclusions have also
been identified in clinopyroxene. Most of the CO2
inclu-sions have leaked; however, some smaller ones (generally 58 mm) preserve two phases or liquid CO2 at room
temperature. The presence of CO2 was determined by
Raman spectroscopy; significant H2O was not detected in
the fluid phase.
The compositions of the daughter phases in the SMIs in both olivine and clinopyroxene are given in Table 2.
A N A LY T I C A L T E C H N I Q U E S
SMIs were reheated using a Linkam TS1500 stage at the Lithosphere Fluid Research Laboratory, Eo«tvo«s University, Budapest. Heating was carried out in a nitro-gen atmosphere on separated and doubly polished olivine grains. The initial high (308C/min) heating rate was
Table 1: Modal composition (vol. %) and inclusion content of the studied xenoliths
Xenolith: NNS07 wehrlite NNS08 NNS09 olivine- NNS10 olivine- NNS11
olivine-Rock type: olivine-clinopyroxenite clinopyroxenite clinopyroxenite clinopyroxenite
Clinopyroxene 53 54 63 76 68 Olivine 42 33 31 20 24 Amphibole 0 11 4 0 3 Intercumulus phase assemblage 5 2 2 4 5 Inclusions in olivine abundant primary SMIs and a few sulphide inclusions primary SMIs in low abundance small (510 mm) pseudosecondary SMIs sulphide inclusions in low abundance small (510 mm) pseudosecondary SMIs Inclusions in clinopyroxene abundant primary SMIs and CO2fluid inclusions and a few sulphide inclusions abundant primary SMIs and CO2 fluid inclusions abundant primary SMIs and CO2 fluid inclusions abundant primary SMIs and CO2 fluid inclusions
abundant, but very small (mostly 51 mm) SMIs
Fig. 2. Petrography of the xenoliths. (a) Texture of a typical olivine-clinopyroxenite xenolith (NNS08); (b) mode of occurrence of the intercu-mulus phase (NNS07); (c) BSE image of a typical clinopyroxene crystal in the cumulate with Ti-rich late-stage rim (NNS08); (d) BSE image of the intercumulus phase (NNS07); (e) image of the intercumulus phase in reflected light (NNS07); (f) breakdown assemblage after amphibole in transmitted light (NNS08). cpx, clinopyroxene; ol, olivine; ic, intercumulus phase; san, sanidine; gl, glass; Ti-mt, Ti-magnetite; rh, rho«nite; ne, nepheline.
progressively decreased to 58C/min approaching the melt-ing temperature of the daughter phases. The stage allowed a cooling rate of 3008C/min in the high-temperature range (above the glass transition temperature).
Homogenized SMIs and the daughter phases of the SMIs were analyzed at the University of Vienna on a Cameca SX-100 electron microprobe using a 15 kV accel-eration voltage, 20 nA beam current, 20 s peak and 10 s background count time for each element. On all heated SMIs, at least two analyses were carried out to test their homogeneity. Compositions of the rock-forming minerals and interstitial phase assemblages were determined at ETH Zu«rich using a JEOL SUPERPROBE JXA-8200 electron microprobe. A 15 kV acceleration voltage and 15 nA beam current were used and a 20 s counting time was applied for each element. For mineral analysis a 2 mm beam diameter was used; for glasses, the beam diameter was enlarged (to 5 or 10 mm) and the beam current was reduced to 10 nA. The following standards were used:
wollastonite for SiO2and CaO; forsterite for MgO;
aegir-ine for Na2O; sanidine for K2O; apatite for P2O5; natural
oxides for FeO, TiO2and MnO.
LA-ICPMS analysis of the SMIs and their host clinopyr-oxene and olivine was carried out at ETH Zu«rich, using an ArF excimer UV (193 nm) laser source at 70mJ output energy and a repetition rate of 10 Hz. The beam diameter varied between 20 and 60 mm for SMI measurements, whereas for bulk measurement of the interstitial phase assemblages a beam diameter of 110 mm was used. SMIs were analyzed and the data reduction was performed using the method of Halter et al. (2002) (see Appendix for further details). Only unexposed SMIs were measured. A beam dia-meter slightly larger than the actual size of the inclusion was used to avoid incomplete sampling of the heterogeneous SMIs. Melt inclusions deeper than 30 mm beneath the sec-tion surface were not measured. Inclusion signals shorter than 15 readings were not considered. In the case of SMI measurements, a 10 ms dwell time was used for all elements.
Fig. 3. Petrography of the melt inclusions. (a) Multiphase SMIs in olivine (NNS07); (b) close-up of a non-exposed multiphase melt inclusion in olivine (NNS07); (c) melt inclusion containing sulphide bleb in olivine (RL) (NNS07); (d) clinopyroxene crystal packed with elongated and oriented melt inclusions and dark CO2fluid inclusions (NNS07).
However, for the clinopyroxene host measurements we used an extended (20 ms) dwell time for Yb and Lu because of their low concentrations. Quantification of the interstitial phases was based on analysis of all the major elements and normalization to 100% without application of any internal standard (Halter et al., 2002). For external standardization the NBS SRM 610 glass standard was used.
Uncertainties are variable for each single analysis, but for SMIs are mainly in the range 5^20 relative % for concentrations of large ion lithophile elements (LILE), high field strength elements (HFSE) and light rare earth elements (LREE), and 15^50 relative % for heavy REE (HREE), which are present in very low concentrations. In the case of clinopyroxene, amphibole, and bulk
intercumulus phase analyses uncertainties are below 15 relative% for most elements.
C O M P O S I T I O N S O F T H E
R O C K - F O R M I N G M I N E R A L
P H A S E S
Major mineral constituents
The clinopyroxene is diopsidic and displays almost homo-geneous cores surrounded by reaction rims (Fig. 2c). The compositional transition between the cores and the rims is not continuous and it shows clustering of mg-numbers around 82^85 (core) and 78^80 (rim), respectively
Table 2: Average composition of the daughter minerals and the residual glass in olivine- and clinopyroxene-hosted silicate
melt inclusions in xenolith NNS07 (in wt %)
Daughter mineral: clinopyroxene amphibole rho¨nite
Host of SMI: olivine (4)
s cpx (1) cpx (2) s olivine (5) s cpx (1) SiO2 437 17 488 433 23 257 07 270 TiO2 463 101 107 455 035 1053 137 651 Al2O3 1129 160 685 1339 107 1793 007 1773 Cr2O3 021 019 012 003 001 025 116 029 FeO 645 122 691 1076 026 200 14 211 MnO 012 009 016 016 006 012 005 018 MgO 1109 112 1431 1331 088 1333 030 1393 CaO 2167 054 2107 1005 094 1080 063 1103 Na2O 103 029 056 308 031 162 031 210 K2O 010 007 001 067 014 001 001 000 Total 10030 9985 9927 10035 9979
Daughter mineral: Al-spinel Ti-magnetite residual glass
Host of SMI: olivine (1) cpx (2) s cpx (2) s olivine (5) s cpx (4) s
SiO2 008 029 012 068 075 571 28 667 139
TiO2 049 097 001 904 041 021 010 015 005
Al2O3 635 470 050 943 482 291 17 217 053
Cr2O3 079 466 175 234 091 n.a. n.a.
FeO 184 346 294 744 64 099 013 042 011
MnO 012 022 001 048 018 n.a. n.a.
MgO 1748 1068 298 202 079 030 024 020 038
CaO 005 038 005 077 044 028 020 067 025
Na2O n.a. n.a. n.a. 947 064 475 300
K2O n.a. n.a. n.a. 306 040 328 161
Total 10088 9884 9917 10050 9788
Number of analyses. n.a., not analyzed.
(Table 2). Both the jadeite component and TiO2 content
increase from core to rim (Table 3). The reaction rim of the clinopyroxene can be distinguished on back-scattered electron (BSE) images (Fig. 2c).
Trace element concentrations (Table 4) show no obvious trends as a function of mg-number; however, rim measure-ments display the lowest mg-number and the highest incompatible trace element concentrations. It should be
Table 3: Composition of selected xenolith-forming clinopyroxenes (in wt %) and results of geothermobarometry calculations
using the method of Putirka
et al. (2003)
Xenolith: NNS07 NNS08
Grain: 1 1 2 2 3 1 1 2 2 3 3
Location: core ovg core ovg core core ovg core ovg core ovg
SiO2 490 460 489 472 488 487 445 495 472 518 481 TiO2 125 392 141 293 118 148 481 091 177 062 201 Al2O3 606 702 597 629 626 694 920 541 796 411 558 Cr2O3 031 017 021 024 030 024 019 026 019 016 039 FeO 514 618 513 588 574 501 518 542 591 519 603 MnO 007 005 008 005 009 007 005 010 009 012 005 MgO 149 128 151 134 150 145 124 147 133 154 137 CaO 226 229 224 231 219 224 229 224 222 227 232 Na2O 059 082 051 073 055 058 084 047 069 044 069 K2O 003 003 002 001 002 001 004 001 002 001 001 P2O5 001 003 000 001 001 001 002 001 000 001 001 Total 9995 9989 9983 9984 9986 10003 10018 9917 9936 10057 9985 mg-no. 838 787 840 802 823 838 810 828 800 841 802 P (kbar) 116 87 109 99 113 116 12 103 111 98 99 T (8C) 1197 1123 1192 1130 1198 1199 1152 1186 1145 1179 1129
Cations per 6 oxygens
Si 181 172 181 176 181 180 166 184 176 189 180 Al(IV) 0189 0277 0190 0235 0192 0203 0341 0158 0237 0108 0203 Al(VI) 0075 0033 0070 0042 0081 0098 0063 0080 0113 0069 0043 Fe3þ 0077 0077 0072 0075 0075 0058 0063 0053 0069 0030 0086 Fe2þ 0082 0116 0086 0108 0102 0097 0098 0115 0115 0128 0102 Mg 0821 0715 0834 0744 0826 0797 0690 0815 0741 0840 0763 Mn 0002 0002 0002 0002 0003 0002 0002 0003 0003 0004 0002 Ti 0035 0110 0039 0082 0033 0041 0135 0025 0050 0017 0056 Ca 0895 0919 0888 0925 0869 0887 0914 0893 0888 0890 0929 K 0000 0001 0001 0001 0001 0001 0002 0001 0001 0000 0001 Na 0042 0060 0037 0053 0039 0041 0061 0034 0050 0031 0050 Cr 0009 0005 0006 0007 0009 0007 0005 0008 0005 0005 0012 End-members (mol%) DiHd 0775 0776 0770 0800 0744 0750 0737 0783 0732 0812 0816 EnFs 0064 0028 0075 0026 0092 0072 0025 0074 0062 0078 0024 CaTs 0033 0000 0033 0000 0042 0057 0002 0046 0063 0039 0000 Jd 0042 0033 0037 0042 0039 0041 0061 0034 0050 0031 0043 CaTi 0078 0138 0078 0118 0075 0073 0169 0056 0087 0034 0101 CrCaTs 0009 0005 0006 0007 0009 0007 0005 0008 0005 0005 0012 ovg, overgrowth.
End-members are calculated with the method of Putrika (1999). DiHd, diopside þ hedenbergite; EnFs, enstatite þ ferrosilite; CaTs, Ca-Tschermak; Jd, jadeite; CaTi, CaTi2AlO6; CrCaTs, CaCr2SiO6.
Table 4: Trace element composition of the clinopyroxenes (in ppm)
Xenolith: NNS07 Grain: 1 1 1 1 1 1 2 2 2 2 3 3 Sc 57 60 52 58 56 59 52 59 50 52 60 56 V 307 400 306 321 281 280 273 289 212 154 280 288 Cr 1355 2437 1487 841 2912 3733 4202 3091 1796 651 2703 2211 Ni 205 211 137 183 169 126 206 190 162 126 123 136 Cu 1390 5374 628 971 5326 452 596 5209 221 641 1760 330 Rb 5048 5043 5036 036 5028 5010 5053 5034 5023 5008 5010 5012 Sr 34 40 31 37 37 41 39 44 42 59 38 42 Y 128 127 111 133 114 115 102 118 92 113 126 115 Zr 30 18 29 36 30 33 25 45 35 50 32 44 Nb 022 5024 5026 029 5028 016 5029 021 5014 5006 012 049 Ba 161 124 5073 129 5061 5018 5167 5123 5049 5033 5022 131 La 140 078 181 251 201 185 222 462 324 482 255 285 Ce 57 36 54 62 57 61 67 111 87 119 70 75 Nd 51 79 65 56 58 66 54 80 70 81 69 67 Sm 219 344 199 291 265 202 199 294 149 256 221 377 Eu 052 132 043 092 095 089 062 127 086 101 091 081 Gd 134 244 314 177 375 248 243 378 189 248 211 315 Yb 121 158 5068 105 095 094 160 130 063 111 077 076 Lu 5010 015 018 018 010 012 5012 020 5008 014 015 007 Hf 158 111 085 134 121 106 060 172 165 164 129 125 Th 015 5013 033 016 025 017 016 5011 012 018 036 034 U 5013 5012 5011 5019 5009 009 5015 5011 5007 5005 011 018 Pr 114 097 096 137 138 110 121 144 115 156 107 121 Er 156 268 075 197 071 139 103 173 077 119 123 168 Ho 053 059 048 079 037 047 051 042 033 043 061 060 Tb 056 040 059 032 048 045 042 034 046 042 044 033 Dy 212 367 222 316 268 262 193 314 193 183 313 261 Tm 5011 5018 027 5010 028 020 5012 5009 012 015 019 016 Xenolith: NNS08 Grain: 1 1 1 1 1 2 2 2 2 3 3 3 Sc 70 78 80 69 71 79 74 83 82 70 82 77 V 354 321 315 279 244 337 344 342 352 338 315 286 Cr 667 1467 1699 1038 1186 2210 1031 3368 1719 2847 1741 1515 Ni 59 90 85 101 86 107 78 88 96 138 106 111 Cu 173 247 204 136 167 191 173 220 203 5215 300 193 Rb 5013 012 007 5017 5010 011 5015 014 5005 5019 5007 010 Sr 69 58 58 68 76 62 56 62 51 58 57 52 Y 162 112 113 106 121 117 131 111 135 123 111 94 Zr 58 34 34 50 64 34 50 35 40 41 32 26 Nb 050 035 022 024 030 035 055 029 012 046 019 020 Ba 102 077 023 5053 5038 042 5024 5024 5037 5085 023 049 La 328 265 213 420 429 238 291 190 261 290 195 161 Ce 114 79 70 103 131 66 93 73 88 88 70 56 (continued)noted that these analyses are not purely representative of the composition of narrow reaction rims, as the ablation pits are larger than the width of the reaction rim. This is also shown by the transitional major element (especially TiO2) content compared with EMPA data
for clinopyroxene cores and rims. Furthermore, if we plot an incompatible trace element concentration against compatible and incompatible elements, clear trends can be observed even in the grain cores that appear to be almost homogeneous with respect to major elements (Fig. 4 and Table 3). The data from the rims fall at the evolved end of these trends. Most incompatible elements display a posi-tive correlation with each other; however, for Y, Yb and Lu this correlation is not so obvious (Fig. 4). More precisely, the concentration of Y shows no systematic change, and the variation of Yb and Lu vs Ce shows a U-shaped pat-tern, with the lowest values at intermediate Ce contents. In terms of compatible trace elements, Ni shows a clear negative correlation with Ce, whereas Sc and V display a less defined distribution. The Cu content decreases exponentially with increasing Ce concentration (Fig. 4).
The olivine grains are homogeneous in composition in each xenolith, but their compositions differ between xeno-liths.Their mg-number ranges from 79 to 85 and the Ni con-tent is between 650 and 1200 ppm. The CaO concon-tent ranges from 004 to 019 wt % with no systematic variation.
The amphibole compositions plot between pargasite and kaersutite with 42^48 wt % TiO2 (Table 5). The
mg-number is between 74 and 77. Their trace element pat-terns normalized to primitive mantle show positive HFSE and LILE and characteristic negative Pb anomalies (Fig. 5). The breakdown assemblage of amphibole consists of rho«nite, Ti-rich diopsidic clinopyroxene, nepheline (Na/K 50), olivine (mg-number 840) and ilmenite (Table 5, Fig. 2f). The bulk composition of the breakdown assemblage is in good agreement with the amphibole composition in terms of both major and trace elements (Table 5).
Intercumulus phase assemblages
The alkali feldspar in the intercumulus phase is sanidine with a Na/K ratio of 06, the clinopyroxene is diopsidic and the residual glass is highly enriched in Al2O3 and
alkali oxides (Table 6). Relative to the parental melt of the cumulates, the bulk composition of the intercumulus phase assemblage has a slightly elevated SiO2content (between
509 and 553 wt %) and very strong enrichment in alkali elements (90^150 wt % Na2O þ K2O), Al2O3
(208^252 wt %) and all incompatible trace elements (Table 7), except for Y and HREE. The CaO/Al2O3ratio
is very low (012^037). The REE pattern is remarkably steep, with La/Yb ranging from 21 to 32 (Table 7).
Table 4: Continued
Xenolith: NNS08 Grain: 1 1 1 1 1 2 2 2 2 3 3 3 Nd 109 72 71 79 102 79 96 80 89 85 82 58 Sm 453 205 254 184 316 264 308 261 293 401 269 225 Eu 139 121 097 078 108 093 106 109 109 095 078 071 Gd 425 305 308 278 261 303 295 223 275 226 305 186 Yb 121 088 083 087 132 096 106 101 081 078 070 085 Lu 012 010 010 008 017 010 014 012 013 009 008 010 Hf 258 166 140 281 207 157 177 130 195 216 158 124 Th 018 013 009 025 013 006 013 004 5006 5011 5002 007 U 004 004 005 014 006 007 5004 5003 5005 5010 5002 5002 Pr 178 126 127 143 203 130 172 133 143 137 119 094 Er 178 124 103 140 143 129 149 130 140 132 101 092 Ho 073 045 050 042 045 051 061 042 051 054 046 035 Tb 067 047 046 048 045 041 044 040 046 048 048 033 Dy 376 265 254 158 331 251 250 312 304 319 254 216 Tm 014 012 015 013 016 013 014 016 013 026 013 008Because clinopyroxene is packed with melt inclusions, it was difficult to ablate the host mineral free of melt inclusions. Rb is highly incompatible in the clinopyroxene but is present in the melt. It was used as a control to ensure that no entrapped melt was included in the ablated volume.
The FeO and MgO contents are low, whereas the mg-numbers are relatively high (49^67).
Downes et al. (2004) have studied glass-bearing horn-blendite and clinopyroxenite xenoliths entrained in intra-plate alkali basalts from Germany and Israel. The bulk compositions of the intercumulus phases compared with the composition of their interstitial glasses have similar total alkali content, but higher concentrations of SiO2
and Al2O3, and lower concentrations of MgO, CaO, FeO
and incompatible trace elements.
H O M O G E N I Z AT I O N
E X P E R I M E N T S A N D R E S U LT S
The melting temperatures of the residual glass in the SMIs hosted in olivine range from 1090 to 11708C. Observation of the onset of glass melting was rather difficult; thus, these values are maximum estimates. Final melting of
clinopyroxene and rho«nite in SMIs occurred between 1190 and 12508C, whereas spinel persisted to 1290^13408C in several cases. Complete dissolution of the gas bubble did not occur in a significant proportion of the studied SMIs, which indicates heterogeneous gas entrapment or partial pressure relaxation in the post-entrapment history of the SMIs. However, decrepitation features were never observed. Despite the relatively slow cooling rate (3008C/min), partial recrystallization of the homoge-nized SMIs could not be observed, neither at microscopic scale nor by scanning electron microscope. Homogenization of clinopyroxene-hosted inclusions was not possible, because melting of the host clinopyroxene began before complete homogenization of the SMIs at atmospheric pressure.
Eighteen reheated SMIs were exposed and polished after homogenization and analyzed by electron micro-probe. All are at least 15 mm in diameter and were ana-lyzed with at least two or three points to check homogeneity. Twelve inclusions appeared to be completely homogeneous and only data obtained from these were used for interpretation.
Melt inclusion compositions were corrected by adding olivine to achieve a KdFe^Mgvalue of 032 that is
appropri-ate for olivines crystallized from basaltic magmas at shal-low mantle depths (Ulmer, 1989). This correction accounts for the post-entrapment crystallization of the inclusions (i.e. overgrowth on the inclusions’ wall). The normalized compositions of the SMIs are shown in Table 8. They show a trend that represents an interval of parental melt evolu-tion. Relationships between the composition and intra-granular textural position (core or rim) could not be evaluated, because in most crystals only one inclusion was measurable. CaO concentration shows a strong negative
Fig. 4. Trace element variation in the rock-forming clinopyroxene as a function of a selected incompatible trace element, Ce.
Fig. 5. Primitive normalized trace element diagram showing the composition of the xenolith-forming amphiboles. Metasomatic amphi-boles found in peridotite xenoliths of the region (Filakovo, Maskova, Ba¤rna) are also shown (Szabo¤ et al., unpublished data).
Table 5: Average composition of the mineral phases forming the breakdown assemblage of amphibole in xenolith NNS08
Bulk ampy (4) s cpx (5) s rh (8) s ne (4) s ol (3) s ilm (1) SiO2 397 408 03 445 15 269 18 442 03 394 03 009 TiO2 457 465 017 459 067 1050 150 015 008 010 002 5142 Al2O3 170 144 04 86 17 157 03 334 04 002 002 027 Cr2O3 007 011 001 010 004 022 006 001 002 002 003 038 FeO 783 757 013 535 034 178 130 066 005 156 055 349 MnO 008 004 001 004 001 007 001 000 000 015 001 027 MgO 133 146 04 127 07 156 07 010 001 455 05 110 CaO 108 118 03 232 03 111 02 182 021 034 003 012 Na2O 507 283 003 085 005 182 019 149 02 002 002 001 K2O 152 134 013 001 001 002 001 448 019 001 000 004 P2O5 005 004 002 002 001 002 000 002 002 005 000 001 Total 9812 9996 9968 9968 10131 9846 mg-no. 752 774 808 609 213 839 359 Sc 49 46 6 79 9 63 20 V 508 452 7 583 85 1073 84 Cr 402 631 197 544 684 779 396 Ni 196 211 11 115 66 424 69 Cu 343 67 11 57 05 51 28 Rb 172 53 01 02 04 Sr 248 515 9 111 12 54 29 Y 186 173 21 414 41 120 71 Zr 781 488 81 1659 695 993 513 Nb 320 268 26 108 96 325 20 Ba 141 269 166 96 42 40 La 78 52 097 137 124 28 417 Ce 206 167 32 442 40 82 91 Nd 165 161 17 366 37 79 82 Sm 45 43 08 114 07 28 28 Eu 159 171 018 330 029 088 080 Gd 467 440 058 941 159 288 194 Yb 142 128 029 337 020 119 067 Lu 018 017 006 034 007 025 012 Hf 221 160 035 453 167 245 204 Ta 165 143 009 132 094 209 034 Pb 082 034 010 072 087 Th 064 006 001 041 008 029 U 016 002 000 007 5013 Pr 307 273 047 714 075 132 174 Er 180 153 063 331 024 121 081 Ho 075 072 008 154 017 050 046 Tb 073 071 004 142 013 047 034Major elements in wt %; trace elements in ppm.
Major element data are from EMPA, except for the bulk composition of the breakdown assemblage, which was calculated by averaging sevaral bulk LA-ICPMS analyses (110mm beam diameter). Numbers in parenthesis show the number of analyses.
linear correlation with Al2O3 concentration; thus, the
CaO/Al2O3ratio is a good tool to describe melt evolution
and major element abundances are plotted as a function of CaO/Al2O3(Fig. 6).
The results of the LA-ICPMS analyses of single SMIs in olivine are shown in Table 9 and Fig. 7. The trace ele-ment content of the SMIs shows a trend similar to that of the major elements. Concentrations of most incompati-ble trace elements increase strongly with decreasing CaO/Al2O3 ratio, resulting in a 2^25 times enrichment
(Fig. 7). Concentrations of Sc and V decrease, and, interest-ingly, Y is also strongly depleted throughout the melt evolu-tion. In most analyses, HREE were below detection limits. The primitive mantle normalized average composition of the most primitive and most evolved SMIs is shown in Fig. 8. Even the most primitive melt composition shows considerable enrichment in all incompatible elements and a characteristic steep REE pattern.
D I S C U S S I O N
Crystallization conditions of the cumulates
The primary goal of this study is to provide an insight into the formation of the olivine-clinopyroxenite cumulate xenoliths and the high-pressure evolution of the parental mafic melts. Therefore, it is important to constrain the P^T conditions of these processes and their location in the lithospheric mantle.
Estimation of the pressure and temperature conditions of cumulate formation is possible using our SMI data and the chemical compositions of the xenolith-forming miner-als (Table 3). Assuming that the melting temperatures of the daughter phases in the SMIs have not been modified
significantly by volatile loss after entrapment, a first-order estimation of crystallization temperature can be achieved based on the heating experiments. The maximum final melting temperatures suggest that clinopyroxene satura-tion occurred at 12508C. However, because addisatura-tion of olivine to the reheated SMIs was necessary to obtain equi-librium melt composition (see above), olivine was probably the first phase to crystallize on the liquidus, implying higher liquidus temperatures. We interpret the presence of the glass (Fig. 3) in the SMIs as a result of incomplete crys-tallization at depth or melting upon reheating by the host basalt during transport of the xenoliths and quenching after eruption. Initial glass melting temperatures indicate a minimum of 10908C for the solidus temperature; how-ever, this might be overestimated because of potential H loss from the melt inclusions (Danyushevsky et al., 2002).
Clinopyroxene compositions and the SMI data for olivines (Table 8) allow the application of the clinopyroxene^liquid geothermobarometer of Putirka et al. (2003), assuming that the phases represent equili-brium compositions, which is likely because of the slow crystallization and is indicated by the homogeneous major element distributions. The equilibrium melt composition was determined from the homogenized olivine-hosted SMIs. A CaO/Al2O3ratio of 07 measured in the most
pri-mitive melt inclusions suggests that crystallization of signif-icant amounts of clinopyroxene did not happen before their entrapment (see below). Thus the primitive melts (average composition of the three most primitive olivine-hosted homogenized SMIs with the highest CaO/Al2O3
ratios) were assumed to be in equilibrium with clinopyrox-ene cores, whereas the most evolved compositions (average composition of the olivine-hosted homogenized SMIs with
Table 6: Average composition of mineral phases and glass (in wt %) in the intercumulus phase assemblage in xenolith
NNS07
sanidine (5)
s glass (5) s cpx (3) s Ti-magnetite (1) apatite (1)
SiO2 638 07 520 05 481 003 024 102 TiO2 044 005 015 003 259 030 162 002 Al2O3 185 03 299 05 492 017 061 010 Cr2O3 000 000 000 001 001 001 013 002 FeO 064 014 091 024 695 013 736 035 MnO 000 000 001 001 010 001 070 003 MgO 006 008 019 021 132 025 228 013 CaO 060 031 018 007 226 049 023 560 Na2O 428 089 1287 113 075 003 004 009 K2O 1079 159 216 025 010 005 011 034 P2O5 002 001 002 001 007 010 001 404 Total 9919 9838 9936 9411 9856
Table 7: Bulk composition of the intercumulus phase assemblages determined by LA-ICPMS in xenolith NNS07
ic1 ic2 ic3 ic4 ic5 ic6 Average (9)
SiO2 522 525 522 509 552 553 531 TiO2 116 103 098 195 125 130 128 Al2O3 252 217 224 208 210 222 222 FeO 281 342 364 402 286 229 317 MnO 002 005 005 004 003 002 004 MgO 152 357 417 293 181 132 255 CaO 812 806 738 655 417 257 614 Na2O 632 603 569 676 665 749 649 K2O 271 357 347 603 703 754 506 P2O5 020 048 043 068 039 032 042 mg-no. 492 650 672 565 530 507 569 CaO/Al2O3 032 037 033 032 020 012 028 Na2O þ K2O 903 961 916 1278 1368 1502 1155 Sc 61 113 83 144 75 43 86 V 117 131 92 210 133 70 125 Cr 62 25 94 46 12 61 50 Ni 109 74 110 71 34 33 72 Cu 147 113 183 89 77 72 113 Rb 47 76 66 133 200 187 118 Sr 819 635 816 442 336 275 554 Y 121 183 116 236 190 156 167 Zr 118 236 136 274 331 259 226 Nb 39 81 84 126 168 136 106 Ba 1552 985 894 709 390 233 794 Cs 035 099 088 165 311 175 145 La 24 41 34 54 47 37 39 Ce 40 67 52 87 76 63 64 Nd 16 26 18 34 26 20 23 Sm 30 52 32 64 49 36 44 Eu 12 15 11 16 11 06 12 Gd 27 42 27 56 40 30 37 Yb 11 17 11 23 20 17 16 Lu 016 023 015 030 026 023 022 Hf 25 50 27 57 60 46 44 Ta 20 39 40 61 75 64 50 Pb 22 35 44 38 54 52 41 Th 42 90 81 126 171 110 103 U 12 30 27 44 60 38 35 Pr 41 68 49 87 73 57 63 Er 12 17 11 24 20 16 17 Ho 045 071 044 092 074 060 064 Tb 042 068 041 084 060 050 057 La/Yb 227 246 315 238 229 215 245 Sr/Y 678 348 705 188 177 176 378 Zr/Hf 471 471 507 483 549 561 507 Nb/Ta 199 208 212 208 224 213 211
CaO/Al2O3045) were thought to be in equilibrium with
reaction rims of the clinopyroxenes. This assumption has some uncertainties, especially for the estimates based on the clinopyroxene core compositions, as these probably re-equilibrated to some extent with later residual melts via diffusion. The method is, however, more sensitive to varia-tion in the jadeite component of clinopyroxene than to the variability in the equilibrium melt composition. For the calculations of the clinopyroxene components, the method of Putirka (1999) was used. Calculated pressure and tem-perature values are shown in Table 3 along with the clino-pyroxene compositions. The average uncertainty of the method is 338C and 17 kbar (Putirka et al., 2003). This might be higher for the P^T estimations based on clinopyroxene cores, because of the uncertainty in the equilibrium melt composition. The estimated crystalliza-tion temperature for clinopyroxene cores ranges from 1185 to 12008C. This is in good agreement with the temperature data for clinopyroxene final melting in the olivine-hosted SMIs but lower than the clinopyroxene saturation tem-peratures obtained by pMELTS modelling (see below). The calculated crystallization pressures for clinopyroxene cores are between 10 and 12 kbar. The reaction rims indi-cate lower crystallization temperatures (1120^11508C) than the cores; however, the estimated pressure values are indistinguishable within the error of the method (87^120 kbar). These values suggest that crystal-lization of the cumulate bodies happened at a depth
of 29^38 km, in the uppermost lithospheric mantle close to the Moho (Huismans et al., 2001) without significant ascent in the recorded cooling and crystallization interval. We estimated the oxygen fugacity during crystallization using the method of Ballhaus et al. (1991) based on the com-position of small Cr-spinel inclusions in olivine. The values scatter between QFM þ10 and QFM þ123 log units (where QFM is the quartz^fayalite^magnetite buffer).
Do the melt inclusions represent the
composition of the parental melt?
The physical conditions of crystallization of the cumulate bodies near the Moho and their cooling rates are different from those of basaltic volcanic systems from which most previous studies of SMIs in olivine have been reported. Most of these earlier studies were carried out on composi-tionally zoned olivine phenocrysts formed by fractional crystallization in volcanic systems. In our case the homo-geneous composition of the host olivine crystals indicates equilibrium crystallization accompanied by slow cooling characteristic of underplated igneous bodies. Iron loss, sug-gested for melt inclusions hosted in magmatic olivine phe-nocrysts (Danyushevsky et al., 2000), may not be significant here because of the equilibrium crystallization of the host olivine. On the other hand, the FeO/MgO ratio in the SMIs is considered to be identical to the latest melt compo-sition, which was in equilibrium with the host olivine. In this scenario, it is assumed that the SMIs maintain
Table 8: Major element composition of the homogenized melt inclusions in olivine of xenolith NNS07 analyzed by EPMA
(in wt %)
Inclusion: 01 02 03 04 05 06 07 08 09 10 11 12 SiO2 487 471 460 451 458 448 465 472 458 473 438 451 Al2O3 227 199 227 209 209 192 205 193 188 184 181 180 TiO2 204 209 243 266 264 312 437 318 304 276 236 213 Cr2O3 0005 0001 0014 0002 0035 0012 0006 0037 0003 0023 0012 0004 FeO 480 611 640 669 660 859 704 580 845 855 849 1001 MgO 402 495 530 553 533 681 573 461 671 687 667 788 MnO 010 009 010 010 012 016 010 009 012 010 012 017 CaO 717 1050 904 938 947 873 764 1080 832 842 1228 1037 K2O 422 286 283 309 304 316 336 286 273 274 272 257 Na2O 530 478 463 481 405 436 466 459 409 412 350 346 P2O5 063 047 039 033 032 035 031 038 040 023 026 023 NiO 0012 0012 0018 0000 0000 0005 0002 0028 0013 0000 0000 0022 Total 9976 9884 9978 9862 9825 9930 10027 9890 9846 9953 9837 9997 Added host (wt %) 0 1 8 7 6 13 10 5 14 14 12 15 CaO/Al2O3 032 053 040 045 045 045 037 056 044 046 068 058 Na2O þ K2O 952 764 745 789 709 753 802 745 682 686 622 603equilibrium with the olivine host composition, which maintains equilibrium with the residual melt. Non-representative remelting of olivine crystallized on the inclusion walls during the reheating experiments or crys-tallization on the walls upon quenching can also induce
significant modifications in the melt compositions. However, these can be properly corrected for, using Fe/Mg exchange equilibria between the host olivine and the entrapped melt (Table 8). In contrast, all those major and trace elements that are incompatible in the host olivine
Fig. 6. Major element variation diagrams for homogenized SMIs (œ) as a function of CaO/Al2O3ratio. The bulk composition of the analyzed
Table 9: Results of LA-ICPMS analyses of individual unheated melt inclusions in olivine
Xenolith: NNS07 Inclusion: 01 s (%) 02 s (%) 03 s (%) 04 s (%) 05 s (%) 06 s (%) 07 s (%) 08 s (%) 09 s (%) Al2O3 1691 10 1718 12 1986 16 1927 15 1786 20 1639 13 1760 14 1950 14 1814 12 CaO 872 59 1190 45 1196 98 1059 116 1007 281 1098 63 1028 60 1077 45 755 62 Na2O 518 10 451 13 517 18 521 16 524 23 467 14 498 15 514 15 550 12 K2O 222 11 178 14 164 25 194 22 182 40 173 17 194 17 199 15 265 13 P 803 325 926 196 52327 1522 566 54314 1046 287 1654 163 1716 114 2297 101 V 387 44 386 44 599 73 523 77 420 159 409 56 337 59 302 49 388 46 Cr 673 159 426 183 1772 161 639 275 1838 332 776 160 435 230 308 210 556 158 Cu 789 71 283 124 711 165 969 146 53127 886 93 1256 75 389 104 1474 56 Rb 28 116 23 136 31 254 41 220 53 325 29 172 33 153 34 115 49 102 Sr 588 25 534 20 578 55 529 55 511 97 500 38 708 31 845 24 711 26 Y 14 178 18 147 25 182 16 306 5198 22 172 18 183 20 129 18 153 Zr 151 65 148 71 218 121 142 146 129 274 149 94 222 72 225 56 182 65 Nb 47 89 33 111 56 176 37 211 29 517 41 132 43 121 40 97 55 87 Ba 480 63 302 83 480 142 283 195 203 343 377 103 320 103 414 70 578 63 Cs 560 578 5179 593 5138 557 5115 593 568 La 21 110 15 138 22 249 19 267 18 432 23 155 22 148 26 105 26 112 Ce 42 82 34 93 61 145 24 228 42 294 48 104 54 92 58 69 65 69 Nd 21 267 26 248 53 355 5775 95 429 32 303 54 215 41 188 28 230 Sm 124 379 109 402 5285 5226 5507 192 373 5210 66 546 65 593 Eu 553 25 409 521 582 5107 31 565 596 38 381 44 346 Gd 5202 147 340 52600 5440 5380 141 440 1012 541 5182 5235 Yb 5519 5430 225 584 5853 5488 5395 5427 5425 5607 Lu 5166 5126 5149 5308 5634 5294 5247 5349 5281 Hf 5823 434 706 5614 51670 51276 5942 5845 5796 5983 Ta 135 399 129 503 217 687 5414 5725 257 461 203 540 254 346 306 326 Pb 5958 5605 5997 51606 51922 51149 51021 51291 5713 Th 5314 122 508 5320 5532 5764 5362 153 611 101 563 343 353 U 105 554 5219 5315 5545 5679 5334 5347 5291 5249 Pr 50 18 42 22 65 41 5186 5175 49 30 66 24 68 18 60 21 Er 5591 5440 5605 51227 52079 486 760 51145 51320 272 638 Ho 5189 5178 5231 5481 5581 5254 5330 5362 5210 Tb 5171 5151 5252 5550 5671 5273 5217 5293 5171 Xenolith: NNS07 Inclusion: 10 s (%) 11 s (%) 12 s (%) 13 s (%) 14 s (%) 15 s (%) 16 s (%) 17 s (%) 18 s (%) Al2O3 1486 11 1757 11 1822 13 1988 11 1689 12 1723 17 1869 13 2107 12 2013 26 CaO 1059 52 755 68 1030 32 1060 21 990 23 638 53 1089 28 1414 38 1379 159 Na2O 442 12 562 11 495 13 503 12 456 12 564 17 481 14 488 12 463 21 K2O 175 14 243 13 206 13 224 12 232 12 288 17 213 13 154 14 165 17 P 5512 1328 178 1152 108 1080 57 1982 43 2033 92 957 107 674 264 873 277 V 483 39 403 48 n.d. 00 n.d. 00 n.d. 00 n.d. 00 n.d. 00 553 35 275 153 Cr 690 147 814 139 661 121 487 83 685 74 869 134 526 119 516 154 1418 208 (continued)Table 9: Continued
Xenolith: NNS07 Inclusion: 10 s (%) 11 s (%) 12 s (%) 13 s (%) 14 s (%) 15 s (%) 16 s (%) 17 s (%) 18 s (%) Cu 534 86 1395 61 564 77 196 79 1079 37 1411 64 515 65 148 169 353 122 Rb 24 145 36 131 41 89 44 52 40 58 61 96 38 74 20 134 28 147 Sr 528 28 748 26 478 26 630 16 778 16 834 28 616 21 648 24 572 38 Y 25 147 14 198 19 122 26 62 23 69 15 172 21 95 25 114 23 236 Zr 146 73 175 71 167 60 159 38 229 34 318 58 140 53 116 72 51 346 Nb 37 100 56 91 38 95 48 51 64 47 77 87 25 93 25 120 29 133 Ba 271 88 539 68 415 61 381 39 571 35 588 68 350 54 247 85 330 88 Cs 586 5108 106 415 045 402 041 439 187 396 037 699 5109 568 La 17 132 33 103 24 101 25 59 42 49 47 94 22 85 15 128 19 194 Ce 50 80 59 76 53 65 56 39 79 36 81 70 48 57 32 89 33 162 Nd 20 250 36 215 41 162 28 116 38 107 33 234 34 140 37 179 24 419 Sm 5160 5137 48 567 76 239 104 225 58 615 68 344 91 414 5151 Eu 547 567 27 394 35 195 28 234 60 328 28 290 27 335 551 Gd 26 71 510 98 369 73 260 94 237 56 648 68 344 100 366 5144 Yb 5584 5412 5766 270 404 5627 5377 5427 5701 5381 Lu 5272 5140 5126 030 500 5078 5155 5047 5261 5118 Hf 166 5510 4845 382 3365 269 447 244 946 340 381 338 154 3767 51109 Ta 102 169 603 258 274 280 156 277 166 450 270 181 254 124 418 151 375 Pb 5910 5737 5153 294 210 194 283 404 395 232 372 5878 167 963 Th 146 454 408 314 277 255 394 127 466 124 623 219 270 202 5613 5219 U 5360 5245 073 478 150 193 5067 285 305 133 274 5197 166 649 Pr 46 19 74 19 63 16 55 11 76 10 127 15 71 13 24 24 48 36 Er 5724 422 583 223 640 234 377 132 566 5808 5402 5647 5763 Ho 5191 5206 n.d. n.d. n.d. n.d. n.d. 5228 5239 Tb 031 5172 111 432 125 244 098 288 121 537 103 354 5389 5250 Xenolith: NNS07 NNS08 Inclusion: 19 s (%) 20 s (%) 01 s (%) 02 s (%) 03 s (%) 04 s (%) 05 s (%) 06 s (%) Al2O3 2063 10 1573 12 1987 14 2000 10 2000 12 2167 11 2156 16 2146 11 CaO 1284 27 978 43 1596 45 1074 29 935 24 1018 28 1029 66 1246 20 Na2O 495 10 458 13 422 14 298 10 352 12 446 11 464 17 491 11 K2O 173 11 210 14 154 17 302 10 239 12 326 11 302 18 203 11 P 869 205 1281 124 222 2743 44 1994 40 2011 58 1490 149 1212 59 V 521 22 305 44 604 41 422 27 304 25 232 34 364 65 n.d. 00 Cr 460 97 737 106 676 145 1694 52 723 59 668 82 489 261 595 82 Cu 754 45 573 78 629 97 154 107 149 79 183 90 169 239 305 68 Rb 26 75 55 83 16 189 54 58 48 46 63 49 38 157 43 57 Sr 596 17 534 26 636 29 865 16 651 16 839 16 944 32 533 18 Y 29 63 21 119 37 114 18 91 19 65 19 79 12 252 25 69 Zr 135 41 181 57 117 88 236 37 168 33 187 37 180 94 156 41 (continued)and have low diffusion coefficients are likely to be repre-sentative of the original melt composition. As data from inclusions smaller than 15 mm were not considered and crystallization should be slow, boundary layer effects are unlikely to play a significant role (Kuzmin & Sobolev, 2004).
Composition of the parental melt
We estimated the composition of the parental melt by fitting trend lines to the SMI compositions on major and trace element plots and taking the values at a CaO/Al2O3
ratio of 07, which is the characteristic value for the most primitive SMIs. The calculated composition of the most primitive melt preserved in the SMIs is given in Table 10 and Fig. 8. As SMIs are absent from the cores of the large olivine grains, it is likely that the most primitive melt represented by the SMIs had already undergone some oli-vine and potentially clinopyroxene fractionation. This agrees well with the fact that this melt composition is rela-tively high in Al2O3and low in MgO compared with most
experimental melts produced by partial melting of garnet lherzolite (Table 10). Crystallization of a significant amount of clinopyroxene before entrapment of the most primitive SMIs is unlikely, as melting of clinopyroxene in
the SMIs terminates before remelting of an adequate quan-tity of olivine from the inclusion walls. Addition of 15 wt % olivine increases the mg-number of the melt to 74, a value expected of a primitive mantle melt in equilibrium with olivine having a mg-number of 91at 3 GPa using the Fe^Mg exchange coefficient (Kd ¼ 036 at 3 GPa) of Ulmer (1989). Before the addition of olivine we calculated the characteristic Fe2O3/FeO ratio of the melt using the
equation of Kress & Carmichael (1991) and used only the FeO content for the mg-number calculation. The resulting primitive parental melt is basanitic (436 wt % SiO2,
mg-number 74) with relatively low total alkali concentra-tions (491 wt % Na2O þ K2O) (Table 10). Similar results
can be achieved by adding both olivine and clinopyroxene; however, in this case the CaO concentration becomes too high, whereas the MgO content remains too low compared with the experimental melts of Walter (1998).
Measurement of the water content of the parent melt is not possible because of the slow crystallization, long resi-dence time at high temperatures and reheating during transportation to the surface, which could have allowed significant H loss from the SMIs by diffusion through the host mineral (Danyushevsky et al., 2002). However, crystal-lization of amphibole in the SMIs and as a cumulate
Table 9: Continued
Xenolith: NNS07 NNS08 Inclusion: 19 s (%) 20 s (%) 01 s (%) 02 s (%) 03 s (%) 04 s (%) 05 s (%) 06 s (%) Nb 39 56 51 79 40 112 52 57 43 47 61 48 99 94 32 67 Ba 284 49 450 62 405 79 740 35 438 34 673 33 610 86 290 48 Cs 5176 5343 5845 5369 013 698 066 355 5466 057 372 La 15 82 25 98 16 147 33 62 34 45 38 52 32 140 27 62 Ce 43 46 52 67 60 78 68 42 65 33 69 38 71 93 53 43 Nd 31 117 21 242 31 206 31 135 30 99 37 107 29 328 35 114 Sm 63 266 69 479 5153 5202 76 218 38 416 133 546 71 281 Eu 24 248 20 485 48 438 13 464 26 213 22 312 585 30 239 Gd 98 181 5239 5353 5249 63 253 48 418 5206 46 354 Yb 5279 5241 5208 5369 307 326 119 492 5353 496 328 Lu 5081 5157 5171 5393 013 694 5532 5150 036 495 Hf 09 408 535 51032 849 232 359 263 511 271 481 702 535 230 Ta 205 188 171 391 213 384 351 188 293 152 198 229 542 347 228 189 Pb 166 265 508 51096 192 397 205 261 213 368 5692 204 287 Th 058 589 224 172 442 498 160 241 174 483 148 5984 345 149 U 5225 141 475 5277 110 390 032 561 125 324 101 737 075 302 Pr 57 11 70 16 57 25 80 11 73 8 86 9 85 24 73 10 Er 5433 5764 5491 358 355 224 307 268 318 5616 263 375 Ho 050 400 5246 5386 5548 037 398 114 278 053 1104 n.d. Tb 097 554 076 594 244 371 5420 055 309 120 249 5218 062 378xenolith-forming phase indicates that a significant amount of water was present in the melt. Compatibility of H2O
during mantle melting is similar to that of P, Ce and Nd (Aubaud et al., 2004; Bolfan-Casanova, 2005), which would allow us to calculate the H2O content of the melt if the
H2O content of the mantle source was known. However,
there are limited data available on the water content of the subcontinental asthenosphere. Embey-Isztin et al. (1993) and Dobosi et al. (1995) have shown that Plio-Pleistocene alkaline basalts of the Pannonian^Carpathian
Fig. 7. Trace element variation diagrams for the olivine-hosted SMIs as a function of the CaO/Al2O3ratio. The bulk composition of the
region are derived from an ocean-island basalt (OIB)-like asthenospheric mantle source. Michael (1988) pointed out that H2O/(Ce þ Nd) ratios for mid-ocean ridge basalt
(MORB) and OIB vary in a narrow range of 95 30. For OIB from Loihi this ratio varies between 107 and 125 (Michael, 1988). Simons et al. (2002) drew similar conclu-sions, suggesting a typical H2O/Ce ratio of 210 20 for a
mantle plume and 150 10 for a depleted MORB mantle (DMM) component. Minor variations in the above-mentioned elemental ratios between contrasting mantle sources allow their use to estimate the water content of
the parental melt of the cumulates assuming an OIB-like mantle plume component as the source. The estimated water content is between 073 and 085 wt % based on the H2O/(Ce þ Nd) ratios of Michael (1988) and between 074
and 09 wt % using H2O/Ce ratios of Simons et al. (2002).
The presence of CO2fluid inclusions and heterogeneous
entrapment of silicate melt and CO2 in clinopyroxenes
(Fig. 3d) indicates CO2 saturation during crystallization.
Considering that primary CO2 inclusions are present in
crystal cores, CO2saturation must have occurred close to
the temperature of clinopyroxene saturation. The presence of a large amount of heterogeneously entrapped CO2
fluid^silicate melt inclusions in clinopyroxene and lack of CO2fluid inclusions in olivine indicates that nucleation of
the exsolving CO2happened on the surface of the
crystal-lizing clinopyroxene. According to the volatile saturation model of Dixon (1997) for alkali basalts with 08 wt % water, a CO2/H2O ratio of approximately three is required
to reach volatile saturation at the estimated crystallization pressure of the cumulates. This implies CO2 contents
between 220 and 255 wt % for the parental melt.
Source of the parental melt
Conditions of parental melt formation can be estimated based on the composition of the SMIs. We modelled the trace element composition of partial melts of a garnet lher-zolite and a spinel lherlher-zolite produced by variable degree of partial melting. We selected Nb, Y, Zr and Nd for the
Fig. 8. Primitive mantle normalized trace element composition of the primitive and the evolved end of the melt compositions shown by the SMIs. For comparison the average composition of the xenolith-hosting alkaline basalts of the region is also shown (Dobosi and Jenner, 1999).
Table 10: Composition estimations for the parental melt of the cumulate xenoliths
Primitive end Fe2O3/FeO 15 wt % 20 wt % 10 wt % ol
Experimental melts of Walter (1998)
of SMI trend calculated ol added ol added and 10 wt % 3 GPa, 4 GPa,
cpx added 14% melting 9% melting
SiO2 444 444 436 433 443 4617 4638 TiO2 236 236 202 190 204 091 145 Al2O3 172 172 146 137 143 1332 981 FeO 894 725 817 820 919 955 1065 Fe2O3 188 191 224 MnO 013 013 013 014 013 018 020 MgO 698 698 1280 1473 1167 1690 1858 CaO 125 125 107 101 123 1069 1031 Na2O 347 347 295 278 283 096 093 K2O 231 231 196 185 185 056 083 P2O5 022 022 019 019 018 Total 9846 9865 9902 9918 9883 mg-no. 582 632 736 762 743 759 757 CaO/Al2O3 073 073 073 073 086 080 105
Model compositions are calculated by addition of various amounts of olivine (ol) and clinopyroxene (cpx) to the composition of the the most primitive silicate melt inclusions (data in wt %).
model because the analytical uncertainty for these elements is relatively small in the SMI measurements (Table 9). Furthermore, these elements are thought to be insensitive to any potential fluid metasomatism in the source. The asthenospheric upper mantle is inferred to be fertile (i.e. not DMM), based on the isotopic composition of the alkali basalts of the region (Embey-Isztin et al., 1993; Dobosi et al., 1995; Szabo¤ et al., unpublished data). We used the primitive mantle composition of McDonough & Sun (1995) to represent the bulk composition of the source and an internally consistent mineral^melt trace element parti-tioning dataset from garnet lherzolite melting experiments at pressures between 28 and 32 GPa (Salters et al., 2002) to model the trace element concentrations in the partial melts. We assumed equilibrium partial melting. Other parameters and results of the model are shown in Fig. 9. Yttrium is compatible in garnet and slightly incompatible in clinopyroxene; consequently, the Ycontent of the melt is a good indicator of the presence of garnet in the source, but is less affected by the degree of partial melting. Niobium is highly incompatible in all the source minerals, thus it is a good indicator of the degree of partial melting; however, it is not sensitive to the phase proportions in the source. We plotted the calculated primitive parent melt composition on a Nb vs Y diagram (Fig. 9). This shows the degree of partial melting and the garnet/clinopyroxene ratio in the source. To eliminate the uncertainty on the absolute concentrations caused by olivine crystallization before SMI entrapment, we plotted the Nb/Y ratio as a function of the Nd/Zr ratio. As a consequence of the high incompatibility of these elements in olivine, its crystalliza-tion does not influence their ratios in the residual melt.
The model indicates partial melting of a garnet lherzo-lite source with a clinopyroxene/garnet ratio between three and seven. The estimated degree of partial melting is 2%.
We compared the major element composition of the melt with the experimental data of Walter (1998) for melting of garnet peridotite. The major element concentrations in the parental melt best resemble experimental melts produced by garnet lherzolite melting close to 3 GPa (Table 10). Melting at higher pressures would result in higher MgO and FeO, but significantly lower Al2O3and slightly lower
CaO concentrations because of the increasing proportion of olivine in the melting mode at the expense of clinopyr-oxene and garnet. The MgO content is still slightly lower and Al2O3content slightly higher in our modelled
paren-tal melt composition than in the experimenparen-tal melt produced at 3 GPa. This might indicate that formation of the melt occurred at even lower pressures, but still in the garnet stability field. The lowest melt fractions produced by Walter (1998) are around 10%, which is a much higher degree of partial melting than that proposed here. This explains the significantly lower incompatible major
element (K, Na, Ti) concentrations in the experimental melts, but does not affect significantly the concentrations of Al2O3, FeO, and MgO, which are buffered by the
equilibrium mineral assemblage.
Evolution of the melt
Based on the textural observations (Fig. 2a) and the heat-ing experiments on the SMIs, the first phase to appear on the liquidus is olivine, shortly followed by cotectic crys-tallization of clinopyroxene. The first melt inclusion entrapment may have occurred when 15% of the paren-tal magma had crysparen-tallized, as 15% olivine fractionation is required to produce the composition of the most primitive melt inclusions from a primitive mantle melt (Table 10). At this stage, olivine was still the dominant phase in the crystallizing mineral assemblage. Subsequently, cotectic equilibrium crystallization of clinopyroxene and olivine
Fig. 9. Petrogenetic model calculations. Modelled compositions of partial melts of hypothetical spinel and garnet lherzolites character-ized by primitive mantle compositions are shown as a function of the degree of partial melting and the clinopyroxene/garnet ratio in the source. Values in per cent show the degree of partial melting; italic numbers indicate the clinopyroxene/garnet ratio in the hypothetical source. The compositions of the parental melt of the cumulates and the host alkali basalts are also plotted to allow comparison with the modelled partial melt compositions. Arrows in the Nb vs Y diagram (a) show the primitive composition of the host basalt.
dominated the melt evolution until a late stage when clin-opyroxene started to react with the residual liquid to form the characteristic TiO2-rich rims and olivine started to be
resorbed (Fig. 2b and c). The clinopyroxene/olivine ratio during crystallization along the olivine^clinopyroxene cotectic must have been high, as the modal proportion of olivine ranges from 20 to 40 vol. % and 15 vol. % olivine must have been crystallized before the onset of clinopyrox-ene crystallization. Because the most evolved SMIs in oli-vine have a CaO/Al2O3ratio of 03 and are approximately
two times enriched in the most incompatible trace elements (Nb, Ta, Rb, Ba) (Fig. 7), we suggest that olivine began to react with the residual melt only after 60 vol. % of the melt had crystallized. Amphibole is an interstitial phase in the xenoliths (Fig. 2f); therefore, its saturation must have happened at a late stage of the melt evolution. In the final stages of crystallization, alkali feldspar, Ti-rich clinopyroxene, apatite and titanomagnetite crystallized from the residual melt (Fig. 2e), leading to an extremely alkaline and Al2O3-rich residual melt. The presence of
compositionally homogeneous glass in the intercumulus phase (Fig. 2d and Table 6) suggests that either there was some melt still present at the time of sampling by the host alkali basalt or the xenoliths were partially melted by the host basalts.
We modelled the evolution of the estimated primitive parental melt using the pMELTS algorithm of Ghiorso et al. (2002), assuming equilibrium crystallization at 1 GPa and a value of 08 wt % for the initial water content. According to the model olivine appears on the liquidus at 13358C, and clinopyroxene saturation occurs at 12808C. Spinel appears at 12208C. The interval between 12808C and 10408C is dominated by clinopyroxene crystallization, simultaneous with precipitation of a minor amount of oli-vine. At 10408C Ti-rich clinopyroxene starts to crystallize, followed by amphibole at 10208C. During amphibole and Ti-rich clinopyroxene crystallization, the amount of olivine decreases slightly and the amounts of normal Ti-poor clinopyroxene and spinel decrease significantly via reaction with the residual melt (Fig. 10). The pMELTS model predicts biotite saturation before the appearance of amphibole; however, biotite was not found in the xenoliths. Thus, we excluded it from the modelled stable phase assemblage. Phase compositions predicted by pMELTS for clinopyroxene have a jadeite component higher than the clinopyroxene cores in the cumulates (Table 11). The low-temperature clinopyroxene has a much higher Ti content than the Ti-rich reaction rims of cumulate clinopyroxene. Assuming equilibrium crystallization, both modelled olivine and clinopyroxene have a lower mg-number close to the solidus than those in the cumulate xenoliths. Furthermore, the clinopyroxene/olivine ratio seems to be slightly overestimated in the crystallizing phase assemblage along the clinopyroxene^olivine cotectic.
The high mg-number of olivine and clinopyroxene suggests that they were not in equilibrium with the residual melt throughout the crystallization sequence, but might have been spatially separated from the main body of the resi-dual liquid at some stage. This and potential errors in the model may explain the other discrepancies. The modelled residual melt at 80 vol. % crystallization (reached at 9808C) is characterized by extremely high Al2O3 (214
wt %) and total alkali oxide (163 wt %) and high SiO2
(547 wt %) similar to the bulk composition of the intercu-mulus phase assemblage, but is depleted in CaO, MgO and FeO to an even greater extent (Table 11). Therefore, it represents the separated residual melt, except for compati-ble elements, which could have been buffered by the sur-rounding peridotite in the main residual melt body. Re-equilibration of the residual melt with peridotite could also be a possible reason for the higher mg-number of the ferromagnesian phases than the values predicted by closed-system crystallization modelling. Below 9808C the pro-gram fails to calculate equilibrium. We conclude that the major trends of melt evolution provided by pMELTS, assuming equilibrium crystallization, are generally in good agreement with the textural^mineralogical obser-vations; however, clinopyroxene seems to be slightly over-stabilized and there are significant differences in the predicted mineral chemistry. Although the model repro-duces the major characteristics of the analytically deter-mined evolutionary trends in melt chemistry, our model assumptions of a closed system and complete equilibrium, and imperfections of the thermodynamic model, may result in significant uncertainties in the numerical values.
Fig. 10. Results of modelling with pMELTS showing the quantity of the minerals crystallized from 100 g of initial melt as a function of temperature (clinopyroxene1, Ti-poor cores; clinopyroxene2, Ti-rich reaction rims).
Garnet precipitation from
the parental melt?
Depletion in Yand HREE displayed by both the evolution-ary trends in the melt inclusions (Fig. 7) and the trace element zonation in the clinopyroxene (Fig. 4) suggests crystallization of a mineral phase in which these elements are compatible. The most obvious explanation would be the primary magmatic crystallization of garnet. Experimental studies in alkaline basaltic melts, however, have reported stability of garnet only at pressures higher than the estimated pressure conditions at which these cumulates were formed (Green & Ringwood, 1964; Green, 1969, 1973). Amphibole could be a potential host for middle REE (MREE) and HREE, yet most reported partition coefficients in the literature are below one even for these elements (Downes et al., 2004). Furthermore, amphibole appears only at a late stage of crystallization and does not display elevated HREE and Y contents (Fig. 2f and Table 5). Liu et al. (2005) reported the presence of garnet-pyroxenite veins from the lithospheric mantle of the North China craton. They interpreted the garnet-pyroxenites to have formed through reaction of an incom-patible element-rich silicic melt and mantle lherzolite. A similar scenario appears plausible in the present case, although relicts of garnet-bearing reaction products were not sampled by the host alkaline basalts or have not been found yet. Decreasing HREE and Y content could poten-tially also be explained by the interaction with garnet-bearing lower crustal granulites, the former (probably pre-Middle-Miocene) presence of which has been reported by Kova¤cs & Szabo¤ (2005) from the same volcanic field. However, Hermann et al. (2001) did not observe depletion
of HREE and Y in basaltic melts intruded into lower crus-tal granulites. Furthermore, geobarometric calculations and the absence of plagioclase in the mineral assemblage of the clinopyroxenites indicate higher pressures than those characteristic for the lowermost crust of the study region during and after basin extension in Middle Miocene^Pliocene (08^09 GPa) (Horvath, 1993).
The fate of the residual melt
A major question concerns the fate of the residual melt left behind after cumulate formation. The most likely possibili-ties are consumption by reaction with the surrounding rocks, in situ crystallization, or migration to higher levels of the lithosphere. These will now be considered in turn.
Interaction between host rock and evolved alkaline melts
as metasomatic agent
SiO2-undersaturated alkaline mafic magmas intruded into
shallow levels of the upper mantle close to the Moho are not in equilibrium with their spinel peridotite host rock. Experimental studies have shown that melts rich in alkalis tend to have high SiO2contents as a result of the
decreas-ing activity coefficient of silica in the melt at buffered (opx^ol assemblage) silica activity (Baker et al., 1995; Hirschmann et al., 1998). Thus, the first reaction to take place is the consumption of some orthopyroxene in the mantle peridotite. A similar metasomatic interaction between peridotite and basanitic melt, involving dissolu-tion of orthopyroxene simultaneous with clinopyroxene and olivine crystallization, has been described by Shaw (1999) and Shaw et al. (2005). This reaction is probably most intense during the early stages of interaction, result-ing in an increasresult-ing SiO2 content in the residual melt of
the cumulates, regardless of the crystallizing mineral assemblage (i.e. clinopyroxene/olivine ratio). Rapid Fe^Mg exchange between mantle olivine and the melt (Gaetani & Watson, 2002; Shaw, 2004) might partially buffer the MgO/FeO ratio in the residual melt, concur-rently generating less magnesian olivine in the mantle peri-dotite. The extent of melt^mantle peridotite interaction depends on the unknown spatial dimensions of the intruded melt bodies (surface/volume ratio); however, enrichment of SiO2 in the early stage of melt evolution
appears somewhat faster than expected from closed-system equilibrium crystallization modelling, and the mg-number of the ferromagnesian phases is slightly higher than expected (Fig. 6, Table 11). As shown above, with an increasing degree of crystallization, an extremely alkaline and Al2O3-rich, but CaO-, MgO- and FeO-poor
melt is generated as a result of the absence of plagioclase in the mineral assemblage. As a consequence of early volatile saturation, the H2O/CO2 ratio and the total amount
of volatiles dissolved in the melt also increased. These late-stage residual melts probably migrate through the host peridotite and may be effective metasomatic agents,
Table 11: Phase compositions at 10208C obtained from
pMelts modelling
olivine clino-pyroxene1
clino-pyroxene2
spinel amphibole apatite residual melt SiO2 381 472 352 420 519 TiO2 284 118 108 — 105 Al2O3 94 153 568 178 214 FeO 244 542 444 161 605 339 Fe2O3 225 208 917 002 042 MgO 374 118 88 168 154 140 CaO 013 193 211 131 558 170 Na2O 168 122 361 779 K2O 737 P2O5 424 055 H2O 179 298 mg-no. 732 795 779 819 424
