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Journal of Petrology | Volume 43 | Number 11 | Pages 2049-2074 | 2002
© Oxford University Press 2002
Partial Melting and Assimilation of Dolomitic Xenoliths by Mafic Magma: the Ioko-Dovyren Intrusion (North Baikal Region, Russia)
1INSTITUT FÜR GEOWISSENSCHAFTEN, JOHANNES GUTENBERG-UNIVERSITÄT, D-55099 MAINZ, GERMANY
2MAX PLANCK-INSTITUT FÜR CHEMIE, ABTEILUNG GEOCHEMIE, D-55020 MAINZ, GERMANY
3INSTITUTE OF EXPERIMENTAL MINERALOGY, RUSSIAN ACADEMY OF SCIENCE, RU-142432 CHERNOGOLOVKA, RUSSIA
4GEOLOGICAL INSTITUTE, RUSSIAN ACADEMY OF SCIENCE, RU-670047 ULAN-UDE, RUSSIA
Received March 9, 2001; Revised typescript accepted April 24, 2002
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL...RESULTS AND INTERPRETATIONDISCUSSIONSUMMARYAPPENDIXREFERENCES |
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A petrological study was carried out on Mg-skarn-bearing dunite cumulates that are part of the Neo-Proterozoic Ioko-Dovyren intrusion (North Baikal region, Russia). Skarn xenoliths contain brucite pseudomorphs after periclase, forsterite and Cr-poor spinel. Fine-grained forsterite–spinel skarns occur with the brucite skarns or as isolated schlieren. Field relationships reveal that the Mg-skarns formed from silica-poor dolomitic xenoliths by interaction with the mafic magma of the Ioko-Dovyren intrusion. Rapid heating of dolomitic xenoliths by the mafic magma caused the decomposition of dolomite into calcite + periclase, releasing much CO2. Further heating quantitatively melted the calcite. A periclase-rich restite was left behind after extraction of the low-density, low-viscosity calcite melt. The extracted calcite melt mixed with the surrounding mafic melt. This resulted in crystallization of olivine with CaO contents up to 1·67 wt %. A local decrease in the silica concentration stabilized CaAl2SiO6-rich clinopyroxene. Brucite/periclase-free forsterite–spinel skarns probably originated by crystallization from the mafic melt close to the xenoliths at elevated fO2. The high fO2 was caused by CO2-rich fluids released during the decomposition of the xenoliths. The above case study provides the first evidence for partial melting of dolomite xenoliths during incorporation by a mafic magma.
KEY WORDS: dunite; dolomite assimilation; partial melting
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL...RESULTS AND INTERPRETATIONDISCUSSIONSUMMARYAPPENDIXREFERENCES |
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Chemical signatures of primitive mafic magmas are primarily determined by the source composition (i.e. enriched or depleted mantle domains) and the conditions of melt formation (e.g. fluid saturation and oxygen fugacity; see, e.g. Green & Ringwood, 1967; Lloyd & Bailey, 1975; Mysen & Boettcher, 1975). Classical examples are the ‘MORB signature’ of tholeiitic basalts (e.g. Sun et al., 1979) and the ‘subduction signature’ of calc-alkaline series (e.g. Miyashiro, 1974). On the basis of modelling results (e.g. Patchett, 1980; Huppert & Sparks, 1985) and geochemical evidence (e.g. Thirlwall & Jones, 1983; Thompson et al., 1984; Puchtel et al., 1997), it is well established that intracontinental mafic and ultramafic magmas can experience significant crustal contamination and, therefore, modification of their original characteristics. Mineralogical and geochemical evidence for basalt contamination by silica-rich crustal material has been detected in many zoned plutons and layered intrusions (e.g. Gray et al., 1981; Stewart & DePaolo, 1992; Francis, 1994; Reiners et al., 1996; Tegner et al., 1999; Maier et al., 2000). Layered intrusions are particularly suitable objects for such investigations because they provide detailed insight into the evolutionary processes of crustal magma chambers (e.g. Turner & Campbell, 1986).
In the early days of igneous petrology, carbonate assimilation was a popular model to explain variations in the composition of mafic magmas, in particular in the context of the origin of alkaline rocks (e.g. Daly, 1910; Shand, 1930). Modern studies of the trace element and isotope compositions of alkaline igneous rocks and mantle xenoliths indicate that this model is untenable. However, several studies dealing with limestone assimilation by mafic magma (e.g. Tilley, 1952; Joesten, 1977; Baker & Black, 1980; Joesten et al., 1994; Owens, 2000) noted major changes in the chemical composition and mineralogy of the mafic magma, such as the resorption of primary igneous minerals (e.g. olivine, clinopyroxene), and the crystallization of Ca–Al-rich clinopyroxene, wollastonite and nepheline. Calc-silicate minerals such as wollastonite, andradite, grossular, spurrite, rankinite, merwinite and melilite were described as having formed at the magma–limestone interface. Actual melting of calcite has never been observed.
In this paper, we present the results of detailed petrological investigations of Mg-skarn-bearing dunites that are part of the Neo-Proterozoic Ioko-Dovyren intrusion. On the basis of geochemical data, model calculations and published experimental results, we provide evidence for (1) the metamorphic transformation of dolomite xenoliths to Ca-free Mg-skarns by quantitative extraction of a calcite melt, (2) the mixing of the calcite melt with the mafic magma, and (3) the effect of CO2 produced by metamorphic breakdown reactions of the xenolith assemblages on the chemical equilibrium between mafic magma and crystallizing minerals.
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GEOLOGICAL SETTING |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL...RESULTS AND INTERPRETATIONDISCUSSIONSUMMARYAPPENDIXREFERENCES |
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The Ioko-Dovyren layered intrusion is situated at the southeastern margin of the East Siberian platform within the Olokit trough (e.g. Konnikov et al., 1994; Fig. 1a). Low-grade Neo-Proterozoic sandstones, schists and carbonates of the Ondoko, Itykit and Asektamur suites were intruded. The metasediments are overlain by basic to intermediate, tholeiitic volcanic rocks belonging to the Inyaptuk and Synnyr suites. Similar ages, 700 ± 20 Ma (U–Pb zircon) for tholeiitic basalts (Neimark et al., 1991) and 673 ± 22 Ma (Sm/Nd whole rock) for gabbros of the Ioko-Dovyren intrusion (Amelin et al., 1996), have been interpreted to be the result of a cogenetic relationship of the effusive and the intrusive series (e.g. Konnikov et al., 1994; Kislov, 1998).
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The Ioko-Dovyren intrusion is about 26 km long and has a thickness of roughly 3·5 km. It is composed of four zones named after the dominant cumulate rock types. From bottom to top these are the plagiolherzolite zone, dunite zone, troctolite zone and gabbro–gabbronorite zone (Fig. 1b; Konnikov et al., 1994). Granophyric gabbronorites form a belt of sills and dykes below and above the intrusion. The subvertical dip of the magmatic layering of the intrusives is concordant with bedding in the enclosing metasedimentary country rocks and is the result of a post-intrusive, regional tectonism at 
550 Ma [see Amelin et al. (1996) for discussion].
Xenoliths of Mg-rich skarns ranging from a few centimetres to 20 m in size are widespread in the dunite zone (Fig. 1b). Two types of skarns were observed (Pertzev & Shabynin, 1979): the first is composed of brucite, olivine and spinel without Ca-bearing minerals; the second is composed of olivine and spinel, sometimes including monticellite and diopside. Pressures between 0·5 and 1·5 kbar have been proposed for monticellite-bearing skarn assemblages (Alexandrov, 1990). Efimov et al. (1986) observed high-temperature melilite–olivine–Ti-fassaite–perovskite associations in calc-silicate xenoliths from the troctolite zone, and calculated a pressure of <1·5 kbar. A minimum pressure of 1 kbar is indicated by the thickness of 3·5 km of the Ioko-Dovyren intrusion.
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SAMPLING AND ANALYTICAL PROCEDURES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL...RESULTS AND INTERPRETATIONDISCUSSIONSUMMARYAPPENDIXREFERENCES |
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A total of 105 igneous rock samples of the Ioko-Dovyren intrusion and 72 samples of typical country rocks and skarn xenoliths were collected during a field trip in 1999. Thirty-eight dunite samples and 13 skarn xenoliths from two sections through the upper 450–500 m of the dunite zone were selected for detailed mineralogical and geochemical investigations. Mg-rich skarn xenoliths are concentrated in this part of the dunite zone. The sections are located in the ‘Central Creek’ (hereafter CC) and ‘Bol’schoi Creek’ (hereafter BC; see Fig. 1b) where continuous outcrops are present. The transition between the dunite and the troctolite zone is preserved in the CC section, whereas it is cut by a shear zone in the BC section. Hence, correlation based on igneous stratigraphy is approximate.
Polished thin sections were used for the petrographical description of the rocks and for microprobe analyses. Analyses were performed with a JEOL JXA 8900 RL-Superprobe at the University of Mainz, using a 20 kV acceleration voltage and a 20 nA beam current for olivine and spinel, and 15 kV and 12 nA for clinopyroxene and plagioclase. Natural minerals (albite, haematite, olivine, orthoclase, rhodonite, wollastonite) and high-purity oxides (corundum, Cr2O3, NiO, periclase, rutile) and metals (Zn) were used for standardization. The PhiRhoZet program of JEOL was used for data reduction. Up to 35 individual analyses of each type of mineral were made for each thin section. A few samples were strongly altered. In these rare cases, only 1–2 analyses of fresh mineral relics were made.
Major and trace element compositions of whole-rock samples were determined by X-ray fluorescence (XRF) using a Philips PW 1404 X-ray spectrometer at the University of Mainz. Reproducibility is better than 0·4% relative for major elements, based on comparison with results of reference samples and on duplicate analyses. Relative deviations between measured and recommended trace element concentrations of international standards range from 1 to 10% [see Wenzel et al. (1997) for details]. Carbon and sulphur were determined using a LECO CS-225 IR apparatus, with a reproducibility better than ±0·5% for C and ±5% for S. Loss on ignition (LOI) was determined gravimetrically at 1000°C. The analytical precision is better than 5%.
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RESULTS AND INTERPRETATION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL...RESULTS AND INTERPRETATIONDISCUSSIONSUMMARYAPPENDIXREFERENCES |
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Petrography
Dunitic cumulates
The dunite zone is composed of dunites, plagiodunites and wehrlites. Samples taken away from skarn xenoliths contain olivine and spinel as cumulus minerals, and plagioclase and clinopyroxene as intercumulus minerals. Textures indicate a crystallization sequence of spinel followed by olivine, then plagioclase and finally clinopyroxene. Intercumulus plagioclase is generally absent in and just above the skarn-bearing horizons, whereas poikilitic clinopyroxene is concentrated at these levels. Anhedral Fe–Ni–Cu sulphides are frequently observed in samples from the upper part of the sections. They often have irregular rims of magnetite and haematite. The samples from the dunite zone are ortho-, meso- and adcumulates according to the criteria of Irvine (1982) (Table 1). There are obvious variations in the textural characteristics even on the scale of a thin section. Parts with cumulus minerals enclosed by intercumulus phases occur together with parts that show adcumulate textures with equilibrated grain boundaries. Mineral compositions of a typical skarn-free dunite, a skarn-bearing dunite (in the following, ‘uncontaminated’ and ‘contaminated dunite’), and a dunite close to skarn xenoliths (<1 m) are given in Tables 2–5.
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Skarn xenoliths
Skarn xenoliths occur in two separate horizons in the CC section, whereas they are found in the whole BC section. There are two types of skarns with different mineralogical compositions and textural relationships. Local contacts between the two skarn types are sharp. The first type is dominated by coarse-grained aggregates (>0·8 mm) of brucite with onion-skin structure (Fig. 2a). These aggregates are pseudomorphs after periclase. Anhedral to subhedral olivine and spinel fill triple junctions between the periclase pseudomorphs (Fig. 2a). The second skarn type occurs as rims on brucite-rich skarns and as isolated schlieren-like bodies within the dunitic cumulate. The schlieren often exhibit an almost vertical orientation (Gurulev, 1983). They are composed of fine-grained (<0·2 mm) olivine and spinel (Fig. 2b). The fine-grained skarns are locally intruded by veins or intercalated with layers (<2 mm thickness) containing coarse-grained olivine and spinel. Poikilitic monticellite enclosing olivine and spinel was observed in skarn schlieren from the BC section (Fig. 2c). Back-scattered electron images revealed that monticellite is locally intergrown with sub-microscopic clinopyroxene. Some of the olivine–spinel skarns contain small (<0·2 mm) grains of pentlandite surrounded by irregular rims of chalcopyrite. Calcite occurs only in veins cutting the skarns and dunites. These veins are clearly of late-stage, low-temperature origin. Mineral compositions of both skarn types are presented in Tables 2 and 3.
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Whole-rock chemistry
Dunitic cumulates
The major element composition of the dunite samples (Table 6) essentially reflects the accumulation of olivine. Most major and trace elements show only small variation along the CC and BC sections. Cu is well correlated with the sulphur content of the dunite samples (not shown) and reflects the distribution of sulphide minerals.
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The variation of CaO and Sr (Fig. 3) can be theoretically used as an indicator of carbonate contamination. However, plots of CaO vs MgO and Sr vs the modal content of intercumulus minerals show that these compositional variations are mainly related to the varying content of intercumulus plagioclase and clinopyroxene (Fig. 3).
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Skarn xenoliths
Periclase–forsterite–spinel skarns are very rich in MgO (>53·6 wt %; see Table 7). MgO varies between 66 and 81 wt % on an anhydrous basis. Periclase skarns have low concentrations of SiO2 (2·9–14·2 wt %), Al2O3 (1·2–4·3 wt %) and Fe2O3 (5·7–9·4 wt %). Na2O and K2O concentrations are below the detection limit (<0·01 wt %) for most samples. The periclase skarns are almost CaO free (<0·3 wt %). Trace element concentrations of periclase skarns are at low levels (Table 7), except for two samples that show elevated Co (>194 ppm), Ni (>1522 ppm) and Cu (>165 ppm) concentrations.
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The major element composition of fine-grained olivine–spinel skarns (Table 7) is dominated by MgO (41·5–49·5 wt %), SiO2 (35·3–32·3 wt %) and Al2O3 (9·2–9·7 wt %). Na2O, K2O and CaO are close to or below the detection limit. Monticellite-bearing skarns are enriched in CaO and contain up to 21 wt % CaO. The concentrations of Rb, Sr, Y, Zr, Nb and Ba are very low in these skarns, whereas Cr ranges between 73 and 550 ppm, and Ni between 32 and 910 ppm. One sample is anomalously high in Ni (>2602 ppm) and Co (>168 ppm). The Ni, Co and Cu enrichment of some skarns is most probably related to the presence of sulphides that have been detected by electron microprobe studies.
Mineral chemistry of dunites
Spinel
Spinels enclosed by cumulus olivine show a large variability in cr-number [100Cr/(Cr + Al) in mol %] from 24 to 71 in rocks from the CC section (Fig. 4a) and from 13 to 58 in the BC section (Fig. 4b). The mg-number of spinel inclusions [100Mg/(Mg + Fe2+) in mol %] is between 33 and 60 in CC (Fig. 4a) and between 40 and 81 in BC (Fig. 4b). High cr-number paired with low mg-number predominates within the lower 100 m of the dunite sections. Dunite samples close to skarn xenoliths contain spinel that has typically low cr-number (<30) and high mg-number (70–80) values.
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Spinel shows increasing cr-number with decreasing mg-number (Fig. 5). This trend is in contrast to the trend produced by fractional crystallization of olivine (Dick & Bullen, 1984). Fractionation of olivine and plagioclase is also excluded because plagioclase does not occur as a cumulus mineral in the dunite zone. Subsolidus hydrothermal alteration would cause a trend of increasing cr-number and/or decreasing mg-number (Kimball, 1990). However, the spinel grains analysed were enclosed by fresh olivine. The observed pattern (Fig. 5) is consistent with near-isothermal equilibration of spinel with olivine of nearly uniform composition (see Dick & Bullen, 1984). Equilibration temperatures of 18 spinel–olivine pairs from the lower, uncontaminated part of the dunite zone scatter between 754 and 794°C using the thermometer of Sack & Ghiorso (1991). These temperatures are too low for the solidus. Hence, diffusional re-equilibration at sub-solidus temperatures occurred in spinel despite the fact that both olivine and spinel look pristine.
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Olivine
The mg-number of cumulus olivine varies from 85·6 to 92·2 in the sections (Fig. 6a and b). Dunite samples collected close to skarn xenoliths contain the olivine that is most Mg rich (mg-number >89·5).
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The NiO content of olivine varies as a function of depth (Fig. 6a and b). Both sections start with high NiO values of 0·23–0·25 wt %. Similar high NiO values were also reported in the literature for the dunite zone below the present sections (Kislov, 1998). The CC and BC sections are similar in that olivine is relatively Ni poor between
200 and 350 m. Increasing NiO contents (0·18–0·22 wt %) were observed in the upper part of the CC and BC sections. The thickness of this NiO-rich zone is larger in the CC section. The repeated increase of the NiO content of olivine is characteristic for multiple replenishment of the magma chamber (see Irvine, 1980). Using the CaO concentrations of olivine, the CC section can be subdivided into four segments (Fig. 6a). Olivine from the lowermost 70 m of the profile contains low amounts of CaO (<0·3 wt %), which is in agreement with literature values from the lower 500 m of the dunite zone (Kislov, 1998). A first increase to 0·88 wt % is observed immediately above the first xenolith-bearing horizon. Within the following 100 m, cumulus olivine is again relatively poor in CaO (<0·3 wt %). A second increase is observed in the upper half of the second skarn-rich horizon. CaO is high (up to 1·05 wt %) in the 75 m following the skarns. Afterwards, the CaO content of olivine decreases to its initial values (<0·3 wt %). Characteristically, CaO-enriched olivines are observed in zones with relatively low concentrations of NiO.
Three samples in or on top of the skarn-bearing horizons show a much larger scatter in their CaO concentrations (Fig. 6a). Here, olivine grains show patches of CaO-poor and CaO-rich compositions (Fig. 7). The calcium distribution pattern observed is different from that of a regular growth zonation. It might reflect reaction of primary cumulus olivine with percolating intercumulus liquid enriched in CaO.
The CaO content of olivine of the BC section varies between 0·05 and 1·67 wt % (Fig. 6b). With the exception of the lowermost olivine grains, which are poor in CaO, the samples have generally high CaO contents. However, CaO concentrations of >0·7 wt % predominate in the upper half of the BC section. A decrease in the CaO content of olivine cannot be observed toward the dunite–troctolite transition, in contrast to the CC section. Again, some samples are characterized by irregular CaO distribution patterns (Fig. 6b).
Plagioclase
Intercumulus plagioclase was detected only in three thin sections from the lower part of the CC section. The anorthite content of different grains ranges from 93 to 42 mol %. Only a few grains show a regular growth zonation with a decreasing anorthite content from core to rim of up to 15 mol %. The observed pattern is interpreted to be the result of fractional crystallization of the interstitial liquid.
Clinopyroxene
Clinopyroxene, classified as augite and diopside (see Morimoto, 1988), is the dominant intercumulus mineral in both sections. Most of the clinopyroxenes contain >1 wt % Cr2O3; only a few samples show lower Cr2O3 concentrations. TiO2 concentrations of clinopyroxenes are between 0·15 and 0·75 wt %. There is no systematic variation of Cr and Ti along the sections, whereas Ca and Al show clear trends.
Clinopyroxene significantly enriched in CaO (>23 wt %) occurs in the first skarn-rich horizon of CC (one sample) and above 150 m in the CC section (Fig. 8a). Clinopyroxene from the BC section is generally rich in CaO (Fig. 8b). These CaO-rich samples are also enriched in Al2O3 (Fig. 8a and b). Aluminium partly occurs in octahedral coordination, as the content of AlIV (2 – Si, on the basis of four cations) does not follow a 1:1 line in an AlIV vs Altot diagram (not shown). Ca is well correlated with AlIV and with AlVI (Fig. 9) indicating that Al is mainly incorporated into the cpx as calcium Tschermak’s component (Cats). Uncontaminated samples contain <0·5 mol % of Cats. The Cats content increases to 4·2 mol % in skarn-bearing dunites.
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An increase in calcium in some clinopyroxenes (marked with a black star in Fig. 8a) from skarn-free dunites is, in contrast, not correlated with an aluminium increase. Ca-enriched clinopyroxenes are associated with plagioclase of relatively low anorthite content whereas clinopyroxenes with low CaO are associated with plagioclase of high An content. Therefore, the chemical variations of clinopyroxene from skarn-free dunites are probably the result of fractional crystallization of the interstitial liquid.
Mineral chemistry of the skarn xenoliths
Spinel
Spinel in both skarn types is extremely chromium poor (cr-number <0·6). Its mg-number is very high, often >95. Iron is mostly in the ferric state according to formula calculation (Droop, 1987). Equilibration temperatures of five spinel–olivine pairs scatter between 680 and 720°C, using the calibration of Sack & Ghiorso (1991). Spinel from coarse-grained veins or layers within the fine-grained parts also shows high mg-number, with a tendency towards elevated cr-number (up to 13).
Olivine
The olivine composition in the periclase skarns is almost end-member forsterite (mg-number 98·1–98·9). The NiO content of this olivine is very low (<0·03 wt %), and the CaO content varies between 0·22 and 1·04 wt %. Olivine in the fine-grained olivine–spinel skarns has a similar composition (mg-number 98·3–98·6; NiO <0·03 wt %; CaO 0·27–0·75 wt %). Veins and coarse-grained layers within the fine-grained skarns contain forsteritic olivine, with slightly increased NiO contents (up to 0·06 wt %). This olivine is variably enriched in CaO (up to 0·56 wt %).
Monticellite
Monticellite is characterized by FeO concentrations between 1·98 and 4·09 wt % and low NiO contents (<0·06 wt %). Its composition varies in the range Ca0·98–1·00Fe0·08–0·10Mg0·92–0·94SiO4.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL...RESULTS AND INTERPRETATIONDISCUSSIONSUMMARYAPPENDIXREFERENCES |
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Composition of the parental magma of the dunite cumulates
The nature of the parental melt of the dunitic cumulates is particularly important for modelling the interaction between the magma and the skarn xenoliths. It has been repeatedly argued that the parental magma of the Ioko-Dovyren dunites was of picritic or komatiitic composition (e.g. Yaroshewskiy et al., 1983; Konnikov et al., 1988, 1994). This conclusion is mainly based on the weighted mean composition of the Ioko-Dovyren intrusion. However, as fractionated melts are thought to have been lost from the magma chamber by volcanic eruptions (e.g. Konnikov et al., 1994), the initial concentration of elements such as Mg, which are preferentially incorporated by cumulus minerals, is strongly overestimated (e.g. Amelin et al., 1996).
Sills and dykes of the plagiolherzolite zone have been interpreted to indicate picritic or komatiitic affinities of the Ioko-Dovyren parental magma (e.g. Konnikov et al., 1994). However, these dykes also contain 30 wt % olivine phenocrysts (Papunen et al., 1992), and hence might reflect the original magmatic liquid plus a phenocryst admixture.
Additionally, Mg-rich olivine (mg-number >90) from the Ioko-Dovyren dunite zone has been used as an indicator for a picritic or komatiitic parental melt [see Kislov (1998) for discussion]. However, our new data do not support this hypothesis. There are separate trends of dunite olivines in an NiO vs mg-number diagram (Fig. 10). The first one, not well documented by the CaO-poor olivine of the CC section, is similar to the expected trend of olivine fractionation from a basaltic magma. CaO-poor olivine forms a second trend towards lower NiO contents at almost constant mg-number. This trend is probably produced by the segregation of an immiscible sulphide melt from the basaltic magma. Such a process removes Ni from the silicate magma and causes the crystallization of Ni-depleted olivine (see Naldrett, 1989). The third trend dominating in both sections is caused by olivine enriched in CaO. It is directed toward olivine with high mg-number (>90) combined with low NiO contents (<0·05 wt %). This trend is fundamentally different compared with the olivine crystallization path from a picritic or komatiitic parental melt (e.g. Cameron & Nisbet, 1982). Our interpretation of this third trend is presented below.
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The MgO/FeO ratio of the parental magma has been estimated on the basis of the least fractionated, Ca-poor olivine that has been probably least affected by carbonate assimilation. This olivine (mg-number 86–87) was in equilibrium with magma with an mg-number of 66·8–64·8 assuming a distribution coefficient of 0·30 for Mg–Fe partitioning between melt and olivine (Roeder & Emslie, 1970). This olivine (NiO 0·25 wt %) was derived from a magma containing 210 ppm Ni based on a nickel distribution coefficient of 9·5 between melt and olivine (see Hart & Davis, 1978). Thus, both mg-number and Ni contents are typical for basaltic melts. A basaltic parent similar to MORB was also predicted by Amelin et al. (1996) on the basis of trace element modelling results.
The program MELTS (Ghiorso et al., 1994) was used to calculate possible crystallization temperatures of the observed minerals at a pressure of 1 kbar. Three types of mafic melts were selected for the calculation: (1) mid-ocean ridge basalt (MORB); (2) plume-type basalt; (3) calc-alkaline basalt from a destructive plate margin [composition data from Wilson (1989)]. A sequence of crystallization similar to that observed in the Ioko-Dovyren intrusion (Spl 
Ol Pl Cpx) is typical for tholeiitic basalts such as MORB. A magma with 9·7 wt % MgO would crystallize spinel (mg-number 71, cr-number 43) at 1260°C, followed by olivine (mg-number 86) at 1230°C, plagioclase (79 mol % anorthite) at 1200°C and clinopyroxene (48 mol % diopside) at 1180°C. These mineral compositions agree well with our observations, except for the mg-number of spinel (Table 3), which has probably been affected by diffusional sub-solidus re-equilibration (see above). Although the cr-number of the model spinel is not re-equilibrated, it strongly depends on the Cr content of the original magma, which is unknown for the parental melt of the Ioko-Dovyren intrusion. On the basis of the modelling results, we conclude that the dunitic cumulates of the Ioko-Dovyren intrusion crystallized in the temperature interval between about 1260 and 1180°C, assuming a pressure of 1 kbar.
Metamorphic evolution of dolomite xenoliths
Previous hypotheses for the skarn formation
The skarn xenoliths observed in the dunite zone are rich in Mg and poor in Ca. Therefore, Pertzev & Shabynin (1979) excluded dolomitic country rocks as suitable precursors for the Mg-skarns. Pertzev & Shabynin (1979) discussed two models for the formation of such skarns, both based on the high-temperature metamorphism of originally magnesite- and brucite-rich rocks. The first model assumed that the carbonate host rocks of the Ioko-Dovyren intrusion contained lens-like magnesite or brucite bodies. This is very unlikely because significant occurrences of such rocks are not known in this area (Gurulev, 1983). The second model suggests a two-stage intrusion of the Ioko-Dovyren massif. Ultramafic cumulates that formed during a first intrusive stage were subsequently hydrothermally altered to brucite or magnesite. In a second step, these were metamorphosed to periclase skarns during the interaction with another magma pulse. The Ioko-Dovyren intrusion clearly formed by the influx of several batches of magma. This is indicated by the spatial variations of the NiO in olivine (Fig. 6). Nevertheless, no structural or other geological evidence was found that would indicate large time gaps between distinct pulses of magma. All dunites show a similar small degree of alteration throughout the entire profile. This does not support a hydrothermal alteration between different magma pulses. It is thus unlikely that the intrusive gap lasted long enough to allow for cooling and hydrothermal alteration of a first magma pulse.
Gurulev (1983) suggested that dolomite host rocks were altered by a hydrothermal system of an early Ioko-Dovyren intrusive pulse to brucite- and serpentine-rich rocks. Subsequent pulses of magma transformed them into periclase-rich skarns. Again, no field evidence for such a contact–hydrothermal alteration process was found.
Transformation of silica-poor dolomite xenoliths to Ca-free brucite/periclase skarns
The cumulate horizons containing the Mg-skarns can be traced to dolomite layers at the margin of the intrusion (Fig. 1b). Hence, it is very likely that dolomite indeed was the protolith of the Mg-skarn xenoliths. The silica-poor dolomite host rocks consist of dolomite (>87 wt %), some calcite (<5 wt %), and detrital quartz and sheet silicates (Gurulev, 1983). The major problem is how to transform these Ca-rich rocks into virtually Ca-free rocks, composed entirely of brucite after periclase, forsterite and spinel.
A T–X(CO2) diagram was calculated for a pressure of 1 kbar using the PerPleX program (Connolly, 1990) for the MgO–CaO–SiO2–H2O–CO2 system with excess calcite (Fig. 11). A quartz-poor protolith, typical for the host-rock dolomites, was assumed. In these rocks, quartz will be consumed during reaction with dolomite and water at low temperature, resulting in the formation of minor amounts of calcite + tremolite in a dolomite matrix. Tremolite will react out, producing the assemblage forsterite + calcite at a temperature below 570°C. The remaining dolomite will decompose to calcite + periclase + CO2 at temperatures between 620 and 800°C, depending on the composition of the fluid phase. Metamorphic carbon dioxide will lead to increasing X(CO2) of the fluid phase, which will escape from the xenoliths into the surrounding magma. At 1200°C, which is the temperature at which plagioclase would form, calcite will melt even with little H2O present [X(CO2) >0·6; Fig. 11]. As our modelling results indicate a magma temperature of at least 1260°C, it is indeed possible to completely melt the calcite in the xenolith.
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The time that is necessary to heat a 1 m dolomite sphere to 700°C assuming a temperature of the mafic magma of 1200°C and a thermal diffusivity of 10-6 m2/s for dolomite (Carslaw & Jaeger, 1959) is <2 days. About 5 days are required to reach 1100–1200°C. Compared with the lifetime of about 104-105 years of a mafic plutonic system (e.g. Stewart & DePaolo, 1992), there was ample time available to complete the metamorphism of the dolomite xenoliths and to melt the calcite.
Calcite melt has a low density of about 2000 kg/m3 and a low viscosity of about 10-3 Pa s under the P–T conditions considered in our case (Genge et al., 1995). Hence, calcite melt can be squeezed out of the xenolith by gravitational compaction caused by either the surrounding mafic magma or the dunitic crystal mush. Decomposition of dolomite and the quantitative removal of calcite melt result in a volume reduction of 70%, and a nearly Ca-free, periclase-rich restite.
Anhedral olivine and spinel fill the triple junctions between brucite pseudomorphs of periclase (Fig. 2a). This could indicate that these minerals crystallized from a silicate mafic melt, or that they recrystallized in contact with a melt. However, skarn olivine and spinel cannot simply be derived from the crystallization of the mafic magma. Olivine is distinctly more enriched in MgO and more depleted in NiO compared with olivine of the dunites. Skarn spinel is almost Cr free, in contrast to the Cr-rich spinel observed in the dunites.
Gurulev (1983) suggested a metasomatic addition of Si, Al and Fe from the mafic magma to the xenolith. Mass balance calculations show that the SiO2, TiO2, Al2O3, Fe2O3, MnO and MgO contents of the brucite–forsterite–spinel skarns agree well with the concentration of these oxides obtained by subtraction of CaO and LOI from the original silica-poor dolomite (Table 8). Consequently, there is no need to postulate mass transfer from the mafic magma into the brucite skarns. However, Na2O and K2O concentrations of the brucite skarns are generally too low compared with the silica-poor dolomite precursors. Obviously, the mobile alkali elements have been lost from the xenoliths together with the calcite melt.
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Reactions at the skarn–magma interface
The formation of forsterite–spinel skarns
Generally, forsterite–spinel skarns can have formed under metamorphic or magmatic conditions. A metamorphic origin for the forsterite–spinel skarns appears to be indicated by high 
18O values of skarn olivine (+16 to +19
SMOW). These values are just a little lower than those of dolomite host rocks (up to +24 SMOW, Krivoplyasov et al., 1982). Several studies have observed calcite–dolomite–forsterite–spinel assemblages in the contact aureoles of silica-poor, aluminous carbonate with igneous rocks (e.g. Rice, 1977; Bowman & Essene, 1982). They were formed at low pressures of 1 kbar and temperatures of 600°C. The removal of calcite as a melt from this type of xenolith at relatively low temperatures (see Fig. 11) could occur only if the fluid phase was water rich. As the metamorphic breakdown of dolomite causes increasing X(CO2) of the fluid phase, calcite will probably not melt at 600°C. A second argument against a carbonate precursor of the forsterite–spinel skarns is the relatively high Cr and Ni concentrations of these skarns (Table 7). Carbonate rocks are typically poor in Cr (11 ppm) and Ni (20–30 ppm; Turekian & Wedepohl, 1961). Hence, it is unlikely that the composition of such skarns is simply the result of partial melting of Si-rich dolomites.
Using the MELTS software package (Ghiorso et al., 1994), it was possible to demonstrate that Mg-rich olivine–spinel assemblages can crystallize from MORB-like mafic melts under specific conditions. Model spinel is generally too rich in Cr (cr-number 30) and olivine is too poor in Mg (mg-number 86–88) at oxygen fugacities typical for mafic magmas [around QFM (quartz–fayalite–magnetite) or NNO (nickel–nickel oxide)]. An elevated fO2 equal to the HM buffer (haematite–magnetite) is needed to crystallize a Cr-depleted to Cr-free spinel but the crystallization path of the melt is changed to clinopyroxene feldspar orthopyroxene. Thus, no olivine would crystallize.
In a second model, we assumed addition of 10 wt % MgO to a fractionating MORB magma at 1230°C and 1 kbar. This assumption is based on preliminary results of dolomite melting experiments in mafic magma (E. S. Persikov, personal communication, 2002). The experiments show that skarn periclase is partly dissolved by the magma. Model olivine with mg-number 88 and model spinel with cr-number 12 are predicted if the system is not buffered at a fixed oxygen fugacity. Only a model system adjusted to very high oxygen fugacity (HM) yields mineral compositions that are in agreement with the composition of olivine (mg-number >97) and spinel (mg-number 98, cr-number <1·5) observed in the fine-grained forsterite–spinel skarns. Hence, if forsterite and spinel crystallized from the magma, high oxygen fugacity would be needed.
Forsterite–spinel skarns often occur as rims around the periclase skarn xenoliths, or as schlieren above Mg-skarn-bearing horizons. Hence, we investigated the effect of CO2-rich fluids on the oxygen fugacity of the melt. We modelled the relationship between oxygen fugacity and composition of graphite-undersaturated C–O–H fluids at 1200°C and 1 kbar (see Ferry & Baumgartner, 1987; Holloway, 1987). Fugacity coefficients of H2O, CO2, CO, CH4 and H2 were calculated using modified Redlich–Kwong equations (Connolly & Cesare, 1993). The results clearly show that fluids characterized by X(CO2) >0·99 are highly oxidizing, with fO2 equal to or higher than the HM buffer. Hence, the fO2 of CO2 fluids released by decarbonation reactions is high enough to cause crystallization of forsteritic olivine and Cr-free spinel from a mafic magma. We favour crystallization from melt at high fO2, as Mg-enriched olivine and Cr-depleted spinel are also observed in dunitic cumulates adjacent to skarn xenoliths (see discussion below).
The local crystallization of monticellite and clinopyroxene within schlieren of fine-grained forsterite–spinel skarns (Fig. 2c) can be explained on the basis of the following reactions:
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The formation of Mg-enriched olivine and Cr-depleted spinel in dunites close to skarn xenoliths
Dunite samples close to skarn xenoliths (<1 m) contain olivine with elevated mg-number (>89·5) and spinel with lower cr-number (<30) compared with ‘normal dunites’ far from xenoliths. The chemical potential of MgO was calculated for the mineral pairs of a CaO–MgO–FeO–Al2O3–SiO2 system using the PerPleX software package (Connolly, 1990) so as to understand the relationships between the several dunites. Activity corrections were made for the specific composition of olivine (two-site mixing with regular solution, W = 4 kJ per site), spinel (ideal mixing in inverse spinels) and clinopyroxene (C2/c Ca-pyroxene) of ‘normal’ dunite and dunite containing Mg-enriched olivine. However, the mineral pairs of both types of dunite are not significantly different in their chemical potentials of MgO within the uncertainty of the calculation (-645 kJ/mol for uncontaminated dunite sample 99-DV44 vs -644·2 kJ/mol for dunite 99-DV92 with Mg-rich olivine). Therefore, the Mg enrichment of olivine and the Cr depletion of spinel in dunites close to skarns cannot be simply explained by assimilation of skarn periclase by the mafic melt.
Modelling results using the program MELTS (Ghiorso et al., 1994) again demonstrate the important role of the oxygen fugacity for the composition of minerals fractionating from a melt (Fig. 12). According to these calculations, the mg-number of model olivine would increase from 86 at the QFM buffer to 90 as fO2 increases and the melt is equilibrated at QFM + 2. Using a Cr2O3 content of 0·08 wt % for the model melt, the cr-number of the spinel would decrease from 36 to <20. Similar results are obtained if the increase in fO2 is combined with the addition of 2 wt % MgO to the model melt.
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log fO2 (QFM + 2). It should be noted that the cr-number of the first crystallizing model spinel is dependent on the Cr2O3 content of the model magma (0·08 wt % in our model).The trends of the calculated mineral compositions are in good agreement with the trends observed in the Ioko-Dovyren dunites. For example, a correlation of increasing mg-number of olivine and decreasing cr-number of spinel is well shown in the dunites of the BC section (Fig. 12), where most of the samples are affected by contamination. Fractional crystallization of olivine and spinel would produce a different trend of decreasing cr-number of spinel and decreasing mg-number of olivine. Although it appears in Fig. 4a that the cr-number of spinel increases in the xenolith-rich horizons of the CC section, samples taken close to skarn xenoliths show significantly lower cr-number than the remaining CC samples. Again, we conclude that the increase in fO2 reflected by the mineral data is caused by the release of metamorphic CO2 from the skarn xenoliths (see Fig. 11, and the interpretation of the forsterite–spinel skarns). The influence of oxidizing CO2-rich fluids must be restricted to a zone adjacent to and above the xenoliths.
Additional support for magma–xenolith interaction comes from preliminary oxygen isotope analysis (Krivoplyasov et al., 1982). The 18O values of olivine from dunite adjacent to skarn xenoliths (+12 to +14 SMOW) are high and indicate an exchange with the assimilated dolomite rocks (up to +24 SMOW). CO2 produced during the metamorphic breakdown of dolomite has a 18O value higher than that of the reacting carbonates (Valley, 1986). This 18O-enriched CO2 will, together with the calcite melt generated in the xenoliths, increase the 18O value of the coexisting mafic magma and olivine.
Mixing of calcite melt and mafic magma
The CaO enrichment of cumulus olivine
The calcite melt released from the xenoliths interacted with the surrounding mafic magma. Evidence for this process comes from the crystallization of Ca-rich olivine. There is a strong increase in the CaO content of olivine in and above the skarn horizons (Fig. 6a and b). In principle, the CaO concentration of olivine is dependent on its forsterite content and on melt composition (Libourel, 1999; Markl et al., 2001). It is particularly sensitive to the amount of calcium, aluminium, alkali and ferrous iron in the coexisting melt, but not to the temperature of formation. There is a restricted range in the mg-number of olivine (86–92), which does not correlate with the CaO content. In addition, there is no evidence for variable Al2O3, Na2O and K2O contents in the parental magma during crystallization of the dunite cumulates (Table 6). Hence, the increase in CaO of olivine is related to an increase of the CaO activity of the melt and/or a decrease of aSiO2. Qualitative estimates using MELTS (Ghiorso et al., 1994) show that addition of 4 wt % CaO to the parental magma will increase CaO in olivine to 0·59 wt % at 1230°C and 1 kbar. Addition of 4 wt % CaO to the magma at a lower pressure of 0·3 kbar causes an even higher content of 0·64 wt % CaO of olivine with mg-number 82–83. Addition of >4 wt % CaO would change the crystallization path and induce abundant precipitation of clinopyroxene.
The crystallization of CaO-rich olivine also depends on the volume of fractionating mafic magma that could be enriched in CaO by mixing with the calcite melt. If there was a large stagnating magma volume, as might be the case on top of the skarn-rich horizon II in the CC section, a thick zone of CaO-rich cumulus olivine could form (Fig. 6a). If, in contrast, injections of fresh, NiO-rich basalt led to a rapid dilution of the contaminated magma (as probably occurred within and above horizon I in CC), the CaO enrichment of olivine is much less pronounced. Although olivine in the BC section is generally rich in CaO (Fig. 6b), the highest concentrations are observed on top of the xenolith-bearing zone between 300 and 375 m. As in the CC section, the zone with CaO-rich olivine is characterized by low NiO concentrations, indicating a stagnant volume of fractionating mafic magma (Fig. 6b).
The formation of Cats-rich intercumulus clinopyroxene
A further important consequence of the mixing of mafic and calcite melt is the crystallization of intercumulus clinopyroxene enriched in the Ca-Tschermak’s component (Figs 8 and 9). In the following, we use the pressure- and temperature-sensitive equilibrium constant K of the reaction
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Overall, the observed changes in mineral composition of the dunites are consistent with contamination of the mafic magma with CaO. The source of the CaO is probably a calcite melt formed during the interaction of dolomite xenoliths with mafic silicate magma. This conclusion is based on field relationships and coincidence of occurrence, although any evidence for calcite melting in the strict sense cannot be provided.
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SUMMARY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL...RESULTS AND INTERPRETATIONDISCUSSIONSUMMARYAPPENDIXREFERENCES |
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On the basis of complex field and petrological investigations, we propose the following model for the formation of the Mg-rich skarns from dolomite xenoliths and the contamination of the basaltic magma (Fig. 13).
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Rapid heating of the xenoliths after incorporation into the mafic magma caused decomposition of dolomite to calcite + periclase + CO2, and, progressively, melting of calcite. Such partial melting has been predicted by experimental results (Wyllie, 1965) to occur in contact metamorphism. But so far, it has never been detected in nature. Calcite melt was squeezed out of the xenolith, as it has a low viscosity and a low density. Xenoliths collapsed by 70% in volume.
Mixing of mafic magma and calcite melt resulted in the crystallization of cumulus olivine with high CaO content and of intercumulus clinopyroxene enriched in the CaAl2SiO6 component. Metamorphic CO2 fluids released from the xenoliths led to increasing oxygen fugacities in the neighbourhood of the skarns. Dunitic cumulates containing very Mg-rich olivine and Cr-poor spinel, and fine-grained forsterite–spinel skarns formed at high fO2 in the neighbourhood of decomposing xenoliths. Saturation of the magma with a Ca-bearing phase (Cpx) caused the crystallization of monticellite and clinopyroxene in the forsterite–spinel skarn schlieren. Stabilization of Ca minerals reflects the increasing activity of CaO during the final stage of skarn formation.
Our results demonstrate that carbonate contamination can dramatically affect the redox state of the crystallizing magma. As a consequence, Mg-rich olivine will crystallize, which could be erroneously interpreted to indicate komatiitic affinities of the parental magma. Detailed geological and mineralogical investigations are necessary to exclude carbonate contamination and to reconstruct potential mantle source characteristics of the mafic magma.
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APPENDIX |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLING AND ANALYTICAL...RESULTS AND INTERPRETATIONDISCUSSIONSUMMARYAPPENDIXREFERENCES |
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ACKNOWLEDGEMENTS |
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Reviews by B. W. D. Yardley, G. Markl, M. Amelio Logon and J. Wolff significantly improved the paper. Technical support by B. Schulz-Dobrick and N. Groschopf with XRF and microprobe analyses is acknowledged. We thank C. Hauzenberger and R. Fernando for helpful discussions. Graphical styling of the figures by G. Feyerherd is very appreciated. Financial support was provided by Deutsche Forschungsgemeinschaft (DFG grant Ba 1876/1 to L.P.B.) and by the Russian Science Foundation (RFBR grant 99-05-04013 to E.G.K. and E.V.K.).
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FOOTNOTES |
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*Corresponding author. Present address: Institut für Mineralogie, Petrologie und Geochemie, Universität Tübingen, D-72074 Tübingen, Germany. Telephone: 0049-7071-2972602. Fax: 0049-7071-293060. E-mail: thomas.wenzel{at}uni-tuebingen.de
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