Journal of Petrology Advance Access originally published online on September 24, 2004
Journal of Petrology 2005 46(1):191-217; doi:10.1093/petrology/egh069
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Journal of Petrology vol. 46 issue 1 © Oxford University Press 2004; all rights reserved
The Grønnedal-Ika Carbonatite–Syenite Complex, South Greenland: Carbonatite Formation by Liquid Immiscibility
1 INSTITUT FÜR GEOWISSENSCHAFTEN, EBERHARD-KARLS-UNIVERSITÄT TÜBINGEN, WILHELMSTRASSE 56, D-72074,TÜBINGEN, GERMANY
2 INSTITUT DE MINÉRALOGIE ET GÉOCHIMIE, UNIVERSITÉ DE LAUSANNE, UNIL-BFSH2, CH-1015 LAUSANNE, SWITZERLAND
RECEIVED NOVEMBER 6, 2003; ACCEPTED AUGUST 2, 2004
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONPREVIOUS WORK AND GEOLOGICAL...ANALYTICAL METHODSPETROGRAPHY AND MINERAL...GEOCHEMISTRYCALCULATION OF PHYSICO-CHEMICAL...DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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The Grønnedal-Ika complex is dominated by layered nepheline syenites which were intruded by a xenolithic syenite and a central plug of calcite to calcite–siderite carbonatite. Aegirine–augite, alkali feldspar and nepheline are the major mineral phases in the syenites, along with rare calcite. Temperatures of 680–910°C and silica activities of 0·28–0·43 were determined for the crystallization of the syenites on the basis of mineral equilibria. Oxygen fugacities, estimated using titanomagnetite compositions, were between 2 and 5 log units above the fayalite–magnetite–quartz buffer during the magmatic stage. Chondrite-normalized REE patterns of magmatic calcite in both carbonatites and syenites are characterized by REE enrichment (LaCN–YbCN = 10–70). Calcite from the carbonatites has higher Ba (
5490 ppm) and lower HREE concentrations than calcite from the syenites (54–106 ppm Ba). This is consistent with the behavior of these elements during separation of immiscible silicate–carbonate liquid pairs. 
Nd(T = 1·30 Ga) values of clinopyroxenes from the syenites vary between +1·8 and +2·8, and Nd(T) values of whole-rock carbonatites range from +2·4 to +2·8. Calcite from the carbonatites has 
18O values of 7·8 to 8·6
and 13C values of –3·9 to –4·6. 18O values of clinopyroxene separates from the nepheline syenites range between 4·2 and 4·9. The average oxygen isotopic composition of the nepheline syenitic melt was calculated based on known rock–water and mineral–water isotope fractionation to be 5·7 ± 0·4. Nd and C–O isotope compositions are typical for mantle-derived rocks and do not indicate significant crustal assimilation for either syenite or carbonatite magmas. The difference in 18O between calculated syenitic melts and carbonatites, and the overlap in Nd values between carbonatites and syenites, are consistent with derivation of the carbonatites from the syenites via liquid immiscibility. KEY WORDS: alkaline magmatism; carbonatite; Gardar Province; liquid immiscibility; nepheline syenite
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONPREVIOUS WORK AND GEOLOGICAL...ANALYTICAL METHODSPETROGRAPHY AND MINERAL...GEOCHEMISTRYCALCULATION OF PHYSICO-CHEMICAL...DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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Carbonatites typically occur in close association with alkaline silicate igneous rocks, either in individual complexes or in a regional association within particular magmatic provinces (e.g. Harmer, 1999
). However, the genetic relationship between carbonatites and associated silicate rocks is complex and not completely understood (e.g. Bell, 1998; Bell et al., 1998). Proposed models for the generation of carbonatites include direct melting of a carbonate-bearing mantle source (e.g. Wyllie & Huang, 1976; Dalton & Presnall, 1998; Harmer & Gittins, 1998; Moore & Wood, 1998), derivation by immiscible separation from a carbonated silicate melt (e.g. Koster van Groos & Wyllie, 1973; Freestone & Hamilton, 1980; Kjarsgaard & Hamilton 1988; Brooker, 1998; Kjarsgaard, 1998) and crystal fractionation of a carbonated alkali silicate melt (e.g. Lee & Wyllie, 1994; Korobeinikov et al., 1998; Veksler et al., 1998a). There is, however, still a lack of robust criteria that can be used to distinguish carbonatitic melts directly derived from the mantle from those produced by differentiation of a parent silicate melt (Bell, 1998).
This study explores geochemical and petrological data from the mid-Proterozoic Grønnedal-Ika nepheline syenite–carbonatite complex in the Gardar Province, South Greenland. The complex represents the most significant occurrence of carbonatite in the province. It is well suited to study the relationship between carbonatite and associated silicate rocks because the carbonatite occurs in close spatial association with nepheline syenitic rocks. Previous studies of Grønnedal-Ika investigated the petrography, whole-rock geochemistry and mineral chemistry of the complex (Emeleus, 1964; Bedford, 1989; Goodenough, 1997; Pearce et al., 1997). In this study, we focus on new O and Nd isotope data from minerals of the syenites combined with Nd isotope data from the carbonatite to address still open questions about liquid immiscibility. In situ trace-element data from minerals are used to evaluate the constraints given by experimental and theoretical carbonate liquid–silicate liquid partitioning data. Additionally, the petrological and geochemical characteristics of the syenites are compared with other syenitic complexes in the province in order to explore differences in terms of magma evolution that can be related to carbonatite formation.
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PREVIOUS WORK AND GEOLOGICAL SETTING |
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TOPABSTRACTINTRODUCTIONPREVIOUS WORK AND GEOLOGICAL...ANALYTICAL METHODSPETROGRAPHY AND MINERAL...GEOCHEMISTRYCALCULATION OF PHYSICO-CHEMICAL...DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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The first detailed account of the Grønnedal-Ika complex was given by Callisen (1943)
, who established the intrusive character of the alkaline rocks. Emeleus (1964) reconstructed the history of the complex based on detailed mapping and petrographic observations. Later, Gill (1972a, 1972b) and Bedford (1989) obtained numerous geochemical data on minerals and whole rocks from the complex and associated phonolite dykes. C, O and Sr and a few Nd whole-rock isotopic data were obtained by Goodenough (1997), Pearce et al. (1997) and Coulson et al. (2003).
The Grønnedal-Ika complex (Fig. 1a) is 8 km x 3 km in exposed dimension and consists predominantly of layered nepheline syenites, which were intruded by a xenolithic porphyritic syenite and a plug of carbonatite (Emeleus, 1964). It was emplaced into the border zone between the Julianehåb batholith (1850–1800 Ma; Garde et al., 2002), which is part of the Ketilidian mobile belt to the south and the Archean block to the north (Fig. 1b). The country rocks comprise quartzo-feldspathic gneisses, occasional amphibolites and rare calc-silicates (Bedford, 1989). The complex was dated at 1299 ± 17 Ma by Rb–Sr (Blaxland et al., 1978). This age is consistent with field observations indicating that it pre-dates emplacement of a suite of olivine dolerite dykes (‘Brown Dikes’) for which an U–Pb age of 1280 ± 5 Ma has been determined (Upton et al., 2003). The layered and laminated syenites are considered to have formed mainly by consolidated by bottom accumulation and adhesion to steep-sided cooling surfaces (Emeleus, 1964). The original structure of the complex comprises two series of layered syenites—the Lower Series and the Upper Series, which are separated by a raft of gneiss (Emeleus, 1964). Several distinct intrusive phases can be distinguished, but gradational relationships between different rock types within intrusive units are common (Callisen, 1943; Emeleus, 1964). For a detailed description of the syenitic rocks from Grønnedal-Ika, the reader is referred to the previous studies of Emeleus (1964) and Bedford (1989). For the purpose of this study, the Upper and Lower Series syenites are combined as the Layered Syenites (LS). At one locality in the Lower Series syenites, some 20 gabbroic xenoliths, roughly 5–30 cm in diameter, were found. Porphyritic microsyenites (PMS) occur as dykes and sheets, and cut across earlier units of the complex. A xenolithic porphyritic syenite (XPS) is the youngest of the intrusive syenites. It contains abundant xenoliths of earlier syenites, trachytes, gneisses and amphibolites. Finally, a central plug of calciocarbonatite was emplaced into the syenites (Emeleus, 1964; Pearce et al., 1997). There are three major and several smaller outcrops of carbonatite that are thought to have belonged to one central plug (Emeleus, 1964). The carbonatite is xenolithic with fragments of syenites and other rocks. It consists essentially of varying amounts of calcite, siderite and magnetite; calcite is dominant. Towards the centre of the carbonatite plug, the amount of siderite increases. Large amounts of magnetite occur where mafic dykes cut the siderite-rich part of the carbonatite (Emeleus, 1964). The syenites in close contact with the carbonatite are altered and impregnated by calcite. Sodalite and sodalite ± calcite veins occur in various units of the complex. There is no field evidence for a significant time-gap between the emplacement of the syenites and the carbonatite. The last stage of magmatic activity, at about 1280 Ma, is represented by the intrusion of a variety of dykes, including lamprophyres, porphyritic dolerites and several olivine dolerites up to a few tens of meters wide. Trachytic, phonolitic and other alkaline dykes were emplaced during renewed magmatic activity during later Gardar (1250–1150 Ma) times. Intense faulting, responsible for the elongated outline of the complex, took place largely following, but in part contemporaneous with, the intrusion of these dikes (Emeleus, 1964).
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Mainly on the basis of whole-rock REE and other trace-element data, Bedford (1989)
proposed an origin for the Grønnedal-Ika carbonatites via liquid immiscibility from a nepheline syenitic liquid. He envisaged that the XPS represents the conjugate silicate liquid to the carbonatite. However, Pearce et al. (1997) concluded that the syenites and carbonatites are unrelated because of differences in their C and O isotopic compositions. They presented data suggesting that the carbonatites are, instead, linked to lamprophyre dikes. The spatial association of carbonatites and lamprophyres is a common feature in the Gardar Province, and a general genetic link between them has been proposed previously (Coulson et al., 2003). The data presented in this study are used to re-evaluate the controversy, and a particular focus is given to the XPS as a critical unit of the complex. |
ANALYTICAL METHODS |
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Electron microprobe analysis
Mineral compositions were determined using a JEOL 8900 electron microprobe (EMP) at the Institut für Geowissenschaften, Universität Tübingen, Germany. An internal
z correction of the raw data was applied (Armstrong, 1991). Both natural and synthetic standards were used for major and minor elements. Measuring times were 16 and 30 s on the peak positions for major and minor elements, respectively. The emission current was 15 nA and the acceleration voltage 15 kV. For nepheline analyses, a beam diameter of 5 µm was used to avoid errors resulting from diffusion of Na. Feldspars with microperthitic exsolution were measured with a 20 µm wide beam and an average of 2–4 analyses was taken as representative for the respective grain. The bulk compositions of oxy-exsolved titanomagnetite grains were reconstructed by combining image-processed (NIH Image software) back-scattered electron images of the exsolved grains and point analyses of exsolved ilmenite and magnetite according to the method described byMarks & Markl (2001).
Laser ICP–MS in situ trace-element measurements
In situ laser ablation inductively coupled plasma–mass spectrometry (LA-ICP–MS) analyses on 150 µm thick polished sections were performed at the EU Large-Scale Geochemical Facility (University of Bristol) using a VG Elemental PlasmaQuad 3 + S-Option ICP–MS system equipped with a 266 nm Nd–YAG laser (VG MicroProbe II). The laser beam diameter at the sample surface was approximately 20 µm. Analyses of calcite were carried out in a beam rastering mode, as ablation of calcite was rapid. Details of the data acquisition have been described elsewhere (Halama et al., 2002). The precision of trace-element concentrations, based on repeated analyses of NIST 610 and 612 glass standards, is approximately ±5% for element concentrations >10 ppm and ±10% for concentrations <10 ppm. Typical detection limits for most trace elements in this study were 0·1–1 ppm, except for Sc, Co and Zn, with detection limits in the range of 1–5 ppm. For the REE, detection limits were generally <0·5 ppm.
Oxygen and carbon isotope analysis
The oxygen isotope composition of hand-picked clinopyroxene separates was measured using a method similar to that described by Sharp (1990) and Rumble & Hoering (1994). Hand-picked clinopyroxene (1–2 mg) was loaded onto a Pt-sample holder. The sample chamber was pumped to a vacuum of about 10–6 mbar and pre-fluorinated overnight. Samples were then heated with a CO2 laser in an atmosphere of 50 mbar of pure F2. Excess F2 was separated from O2 by conversion to Cl2 using KCl held at 150°C and the extracted O2 was collected on a molecular sieve (13X). Oxygen isotopic compositions were measured on a Finnigan MAT 252 mass spectrometer at the Institut für Geowissenschaften, Universität Tübingen. The results are reported as the deviation from Vienna Standard Mean Ocean Water (V-SMOW) in the standard -notation. Replicate analyses of the NBS-28 quartz standard yielded an average precision of ±0·1 (1
) for 18O values. In each run, standards were analyzed at the beginning and the end of the sample set. A correction, generally <0·3, was then applied to the data, equal to the average difference between the mean measured value and the accepted value for the standard of 9·64.
The carbon and oxygen isotope compositions of the carbonates were measured in automated mode using a GasBench from ThermoFinnigan connected directly to a Finnigan MAT 252 mass spectrometer. The method has been described in detail by Spötl & Vennemann (2003). About 200–400 µg of sample material was recovered as powder by micro-drilling into the calcite. CO2 was obtained by reaction of the sample with several drops of 100% orthophosphoric acid and was carried via a He stream over traps to remove water vapor and a gas chromatograph-column to separate CO2 from possible interfering gases before being passed into the mass spectrometer for isotopic analysis. The C and O isotope compositions are reported in the standard -notation, relative to PDB and V-SMOW, respectively; standard analytical errors are ±0·1 (1) for both 13C and 18O.
Nd isotope analysis
For Nd isotope analyses, 15–20 mg of hand-picked clinopyroxene from the same separates used for the O analysis, and 5–8 and 40 mg of whole-rock powder of carbonatite and basement samples, respectively, were used. Details of the analytical procedures have been described elsewhere (Halama et al., 2003). 143Nd/144Nd ratios were normalized for mass fractionation to 146Nd/144Nd = 0·7219. The average 143Nd/144Nd ratio obtained for the Ames Nd-standard (Roddick et al., 1992) was 0·512119 ± 10 (1, n = 42) during the course of this study, and the La Jolla Nd standard gave 0·511831 ± 30 (1, n = 12). The total procedural blank (chemistry and loading) was <100 pg for Nd.
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PETROGRAPHY AND MINERAL CHEMISTRY |
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TOPABSTRACTINTRODUCTIONPREVIOUS WORK AND GEOLOGICAL...ANALYTICAL METHODSPETROGRAPHY AND MINERAL...GEOCHEMISTRYCALCULATION OF PHYSICO-CHEMICAL...DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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Syenites
The major mafic silicate in the Layered Syenites is euhedral to subhedral or poikilitic clinopyroxene (Fig. 2a–c). Clinopyroxenes are typically cumulus phases and commonly rimmed by intercumulus biotite. Most clinopyroxenes can be classified as aegirine–augites (Morimoto et al., 1988
), but diopsidic pyroxene and aegirine occur occasionally. Typical clinopyroxene compositions are summarized in Table 1 and compositional trends are shown in Fig. 3a–c. In some samples, clinopyroxene shows a weak, patchy zonation with limited compositional variability. Other samples contain clinopyroxenes with a marked growth zonation, with a compositional range in a single crystal as large as that in the entire sample suite. Compositions vary between Di56Hed34Aeg10 and Di3Hed2 Aeg95. In comparison with compositional trends from other alkaline complexes of the Gardar Province (Fig. 3d), the extreme enrichment in Fe2+ at low values of Na, as seen in Ilímaussaq (Larsen, 1976) and Igdlerfigssalik (Powell, 1978), is lacking in Grønnedal. Clinopyroxene compositions in a gabbroic xenolith show a small range of Di65–59Hed30–36Aeg5.
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In the XPS, clinopyroxene phenocrysts reach up to 2 mm in length and show a marked discontinuous zoning, whereas matrix clinopyroxene grains are considerably smaller (<200 µm) and optically unzoned (Figs 2d and e, and 4a and b). The clinopyroxene phenocrysts in the XPS are characterized by a relatively Mg-rich core, overgrown by a distinctly more Na- and Fe3+-rich rim with a compositional gap between (Fig. 3c). The most primitive core compositions in the XPS reach values of Di57Hed29Aeg14, whereas the most evolved rims are Di7Hed22Aeg71. Matrix clinopyroxene compositions in the XPS overlap with the more evolved phenocryst rim compositions.
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Olivine is completely absent from the samples studied, but scarce relics of Fe-rich olivine have been observed (Emeleus, personal communication). Amphiboles are also scarce (Bedford, 1989
), but they occasionally occur as poikilitic or subhedral intercumulus grains. Only one sample (GM 1531) was investigated in which poikilitic amphibole with a katophorite composition is the dominant mafic silicate phase (Fig. 2b). Biotite, usually as intercumulus grains (Fig. 2a), occurs frequently and is often associated with hematite. All syenites from the Grønnedal-Ika complex are nepheline-bearing (Emeleus, 1964), and nepheline usually occurs as euhedral prisms (Fig. 2a–c). Nepheline is chemically unzoned and its composition is fairly homogeneous within individual samples and also within the whole sample suite, including the PMS and the XPS (Table 2). There are also no significant compositional differences between phenocrysts and matrix nepheline. The compositional variability spans the range Ne73–68Ks20–15Qz7–16. The FeO and CaO contents of nepheline in all rock types are invariably <1 and <0·1 wt %, respectively. Nepheline is often replaced by the alteration products ‘gieseckite’ (fibrous, micaceous aggregates) and cancrinite (Fig. 2c). Cancrinite also occurs as individual grains, partly with poikilitic character suggestive of an intercumulus origin.
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The abundant platy crystals of alkali feldspar frequently show perthitic exsolution into almost pure albite and orthoclase. Alkali feldspar phenocrysts of up to 1 cm in length occur in both the PMS and the XPS (Fig. 2d). Some samples contain feldspar with exsolution on a µm scale, so that the original bulk composition could be obtained using a defocused beam (Table 2). These bulk compositions are fairly homogeneous throughout the different rock types and vary in the range An0–1Ab39–54 Or46–60.
Primary magmatic titanomagnetite is relatively rare and characterized by a trellis-type oxy-exsolution texture [terminology after Buddington & Lindsley (1964)] of magnetite with lamellae of the hematite–ilmenite–pyrophanite solid solution series (Fig. 4c and d). The compositional variation in the lamellae (Ilm5Pyr93Hem2–Ilm56Pyr23Hem21) is in contrast to the constant composition of the magnetite (Mag98–99Usp1–2). Reintegrated oxide compositions vary from Mag61Usp39 to Mag72Usp28 (Table 3). This range reflects the chemical variability in the lamellae and is considered to cover fairly well the overall variation in the samples, as the lamellae are volumetrically of minor importance (10–30 vol. %). Except for MnO contents of up to 11·3 wt %, concentrations of minor elements in the reintegrated oxides are low (Al2O3 
0·12 wt %, MgO 0·03 wt %, ZnO 0·24 wt %). Late- to post-magmatic hematite associated with biotite has almost a pure hematite end-member composition (Table 3, Fig. 4e).
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Minor phases present include short prismatic apatite, zircon, blue sodalite, analcite, calcite and fluorite. Calcite appears to be a primary magmatic or late magmatic interstitial phase and not an alteration product (Fig. 2c). Sodalite can occur as small, interstitial grains, partly associated with cancrinite (Bedford, 1989
), or as larger aggregates in almost pure sodalite veins. The chemical characteristics of the minor minerals relevant to this study are briefly summarized based on the work of Bedford (1989). Sodalite contains around 8 wt % Cl but only minor amounts (<0·3 wt %) of S. Cancrinite (Fig. 4f), with the ideal structural formula (Sirbescu & Jenkins, 1999) Na6Ca2[Al6Si6O24](CO3)2 + nH2O, also has low (<0·15 wt %) concentrations of S and shows some exchange of Ca for Na. Calcite in the syenites is almost pure CaCO3.
Carbonatites
The Grønnedal-Ika carbonatites can be classified as calcite carbonatite and calcite–siderite carbonatite (Woolley & Kempe, 1989) or, on the basis of whole-rock geochemical data with a maximum of 23·1 wt % Fe2O3 (Bedford, 1989), as calciocarbonatite and ferruginous calciocarbonatite (Gittins & Harmer, 1997). The carbonatites consist essentially of varying amounts of calcite, siderite and magnetite. They are poor in U–Th–Nb–REE minerals compared with other carbonatites of the world. Magnetite is exclusively secondary after original siderite as a result of decarbonation and oxidation in the vicinity of the dolerite dykes (Bedford, 1989). In places, anhedral to subhedral calcite crystals form almost 100% of the rock (Fig. 2f). Individual calcite crystals are 50–500 µm in diameter and commonly show prominent twinning. Siderite occurs as euhedral rhombs or as minute grains and is commonly rusty-brown because of surface weathering. In the siderite-bearing samples, calcite grains tend to have a cloudy appearance. Minor quantities of apatite, sphalerite, pyrite, strontianite, monazite, pyrochlore, feldspar and alkali amphibole are also present in some samples (Emeleus, 1964; Bedford, 1989; Bondam, 1992).
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GEOCHEMISTRY |
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Trace elements
Calcite
Chondrite-normalized REE patterns of calcites are shown in Fig. 5a. Three different patterns can be distinguished in the calcites from the carbonatites. One sample—a pure calcite carbonatite (GM 1502)—contains exclusively calcite characterized by a steep REE pattern with high LREE abundances (type A). There is a slight positive Eu anomaly present (Eu/Eu* = 1·25) and the degree of LREE enrichment relative to the HREE expressed as LaCN/YbCN is
69. The overall REE enrichment of this sample is within the lower range of whole-rock analyses from the Grønnedal carbonatite and typical of carbonatites in general (Woolley & Kempe, 1989). Two distinctly different types of REE patterns occur in calcites from sample GM 1498, where type A patterns are absent. The sample is a calcite–siderite carbonatite in which siderite is characterized by brown oxidation rims. Both REE patterns show a strong depletion in LREE relative to the type A pattern (Fig. 5a). The main characteristic of the type B REE pattern is a convex-upward shape. In contrast, type C calcite has about one order of magnitude lower normalized LREE and MREE abundances and exhibits an almost continuous increase from very low LREE abundances towards the HREE. There are no textural differences between types B and C. Interstitial calcite from two syenitic samples (Layered Syenites, samples GR 44 and GR 63) shows LREE-enriched patterns similar to type A, but with slight negative Eu anomalies (Eu/Eu* = 0·66) and a flatter slope (LaCN/YbCN = 10·5–26·3). In a similar fashion, the REE pattern of the XPS whole-rock matrix is flatter than that of the whole-rock carbonatites (Fig. 5a). In addition to the REE, Sr, Ba, Pb, Y and Ga also occur in significant amounts in the calcites, whereas all calcites have very low concentrations of the HFSE (e.g. Nb, Ta, Zr, Hf), mostly below the detection limits (Table 4). Type A calcite (GM 1502) is characterized by a relative enrichment of Ba, Sr, Eu and Ga compared with the calcites from the syenites (Fig. 6a). Type B and C calcites show marked enrichments in Pb and Sr compared with the neighboring elements. These features and the low REE contents of type B and C calcite are not entirely clear. They are consistent with postmagmatic loss of these elements by fluid-controlled recrystallization and remobilization of carbonate phases (Bau, 1991; Bau & Möller, 1992), for which brown rims around siderite might be an indication, or they could be related to the contemporaneous siderite crystallization itself.
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Clinopyroxene
A detailed study of the REE and trace-element abundances in clinopyroxene and amphibole in the Grønnedal-Ika syenites has been made by Marks et al. (2004a)
. In this study, we concentrate on data from clinopyroxene phenocrysts in the XPS to evaluate whether features of the trace-element chemistry indicative of silicate–carbonate liquid immiscibility can be detected. Typical trace-element concentrations in clinopyroxene cores and rims are summarized in Table 4. Chondrite-normalized REE patterns (Fig. 5b) indicate distinct differences between core and rim compositions. The core patterns are characterized by a smooth increase in normalized abundances from La to Nd and a smooth decrease from Nd to Tm or Yb, with a slight increase towards Lu. LaCN/YbCN ratios in the cores range from 4·7 to 6·9 and Eu anomalies are slightly positive, with values between 1·02 and 1·37. In the Na-rich rims, REE patterns are sinusoidal with generally lower LREE and higher HREE abundances than in the cores (Fig. 5b). LaCN/YbCN ratios range between 0·32 and 0·75. Eu anomalies in the rims are not significantly different from those in the cores and vary from 0·99 to 1·24.
Trace-element concentrations of XPS clinopyroxene cores and rims normalized to primitive mantle values (McDonough & Sun, 1995) are shown in Fig. 6b. A principal feature of all the patterns are strong negative Ba, Pb and Ti spikes and depletions in Sc, V and Co. Negative Sr spikes are present in all patterns except for two phenocryst cores. These two cores are also characterized by the lowest Zr and Hf values. All rim compositions have strong positive Zr, Hf and Sn, but negative Li anomalies. In comparison with the core compositions, rims are also enriched in Zn, but depleted in Rb, Co and Sc. All rim patterns are very similar to those of aegirine–augites from the LS, both in enrichment level and shape (Fig. 6b). Aegirines from another silica-undersaturated syenitic complex of the Gardar Province, Ilímaussaq (Marks et al., 2004a) have much lower absolute trace-element concentrations, but the shape of their patterns has a striking similarity to that of the phenocryst rims, except for Li and Pb.
Ratios of isovalent trace elements with a similar geochemical behavior (Bau, 1996) were investigated in order to identify characteristic differences between the growth of clinopyroxene cores and rims (Fig. 7). Zr/Hf ratios in the cores vary between 31·5 and 42·1, and scatter around the primitive mantle value of 37·1 (McDonough & Sun, 1995). In the rims, however, Zr/Hf ratios are higher than in the cores, with values ranging from 45·7 to 58·8. In contrast, Y/Ho ratios in individual phenocrysts are slightly higher in the cores (Y/Ho = 19·6–22·7) compared with the rims (Y/Ho = 16·8–20·5). Nb/Ta ratios scatter widely (Nb/Ta = 4·2–16·3) and there is no consistent relation between core and rim compositions.
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Carbon and oxygen isotope analyses
Goodenough (1997)
, Pearce et al. (1997), Coulson et al. (2003) and Taubald et al. (2004) have reported C and O isotopic data from Grønnedal-Ika, but data on silicates from the syenites were lacking so far. Recent oxygen isotope studies have shown that individual coarse-grained minerals are less sensitive to low-T alteration than whole-rock samples (e.g. Vroon et al., 2001). Extensive studies of mineral separates from ocean island basalts (Eiler et al., 1997) and oceanic-arc lavas (Eiler et al., 2000a) demonstrated that analyses of mineral separates consistently span a narrower range than comparable whole-rock data. At Grønnedal-Ika, whole-rock oxygen and carbon isotope data were used as an argument against liquid immiscibility between syenites and carbonatites because of their distinct differences in 18OV-SMOW and 13CPDB (Pearce et al., 1997). To evaluate this argument in the view of mineral data, the oxygen isotope compositions of fresh magmatic clinopyroxene and amphibole in the syenites were measured, augmented by C and O isotopic data of microdrilled calcite (Table 5).
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The oxygen isotopic composition of clinopyroxene from the syenites varies from 4·2 to 4·9
and that of amphibole (18O = 4·7) is within this range. The average 18O value of clinopyroxene from the layered syenites is 4·6 ± 0·2 (1). There is no significant isotopic compositional difference between the layered syenites and the late syenites. It was not possible to analyze core and rim compositions in the clinopyroxenes from the XPS separately, but, from petrographic observations, the proportion of rim composition in the clinopyroxene population should be relatively small compared with the core composition. 18O values of clinopyroxene from the XPS are in the lower range of the entire dataset. Clinopyroxene from a gabbroic xenolith (sample GR 08) is slightly, but significantly, higher in 18O than from the syenites, with 18O = 5·1.
In the studied suite of samples, the oxygen isotope composition of calcite ranges from 7·8 to 8·6 18O and carbon isotope composition varies between –3·9 and –4·7 13C. Our data overlap with the data from the literature. They are shown in Fig. 8, together with the field of primary, mantle-derived carbonatites as proposed by Taylor et al. (1967). The C and O isotopic compositions of the Grønnedal-Ika carbonates are within the upper half of the primary carbonatite field of Taylor et al. (1967).
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Nd isotope analyses
Nd isotopic analyses of mineral separates and whole-rock samples are presented in Table 6.
Nd(T) values were calculated for T = 1300 Ma based on the Rb–Sr isochron age of the complex (Blaxland et al., 1978). They range from +1·8 to +2·8 for the syenites and from +2·4 to +2·8 for the carbonatites. The average Nd(T) of the carbonatites (2·59 ± 0·20) is slightly higher, but not significantly different from the average of the syenites (2·36 ± 0·40). Thus, the Nd isotopic composition of the Grønnedal-Ika complex appears to be relatively homogeneous. However, Nd(T) is slightly lower compared with whole-rock syenite data from Grønnedal-Ika (Nd = 3·2–3·9; Goodenough, 1997) and a value for the Gardar mantle source (Nd > 4) defined by Andersen (1997). Plotted in a Sm–Nd isochron diagram (Fig. 9), all analyses of the mafic minerals and the carbonatites scatter closely around the 1300 Ma reference line. Clinopyroxene from a gabbroic xenolith (sample GR 08) yielded a distinctly lower Nd(T) value than those typical for the syenites [Nd(T) = –1·7]. A whole-rock sample of relatively unaltered basement gneiss has Nd(T) = –13·3.
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CALCULATION OF PHYSICO-CHEMICAL PARAMETERS |
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TOPABSTRACTINTRODUCTIONPREVIOUS WORK AND GEOLOGICAL...ANALYTICAL METHODSPETROGRAPHY AND MINERAL...GEOCHEMISTRYCALCULATION OF PHYSICO-CHEMICAL...DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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Temperature, silica activity and oxygen fugacity in the nepheline syenites
The quantification of intensive crystallization parameters can be used to distinguish separate phases in the evolution of the complex or individual units. Furthermore, the determination of equilibration temperatures is important for obtaining accurate estimates of the oxygen isotope composition of the melt from the mineral data, because mineral–melt isotopic exchange is temperature dependent. The quantification of silica activities and oxygen fugacities is relevant for comparison with other Gardar complexes to evaluate the possible influence of these parameters on the origin of the carbonatite.
Equilibria among Ne, Ab and Jd components in nepheline, alkali feldspar and clinopyroxene solid solutions, respectively, provide means to estimate temperatures and silica activities in nepheline syenites (Markl et al., 2001). These equilibria are believed to reflect the conditions during formation of the last mineral in a specific assemblage (Markl et al., 2001). In the absence of a reliable pressure estimate, 1 kbar was used for all calculations. This value is an estimate from the Ilímaussaq intrusion further south in the Gardar Province, based on fluid inclusion studies (Konnerup-Madsen & Rose-Hansen, 1984). For the nepheline syenites of Grønnedal-Ika, the following three equilibria are relevant:
 | (1) |
 | (2) |
 | (3) |
Following the approach of Markl et al. (2001), we assume equilibration of the magmatic mineral assemblage is achieved as long as melt remained in contact with the minerals and in the absence of growth zonation (feldspars and nephelines). Zoned clinopyroxene crystals record various stages of crystallization, and the increase in Na+ and Fe3+ in clinopyroxene rims and along cracks may reflect oxidation by residual melt or fluid phases (Marks et al., 2003). Therefore, estimates of magmatic T and aSiO2 conditions were made using clinopyroxene compositions that presumably reflect early crystallization, i.e. the most diopside-rich compositions of the matrix clinopyroxenes. Activities of jadeite in clinopyroxene (aJd) were calculated after Holland (1990) at temperatures of 800, 900 and 1000°C. For feldspars, the complete range of compositions measured was used to determine the activity of albite (aAb) after Fuhrman & Lindsley (1988), again at 800, 900 and 1000°C. Nepheline activities (aNeph) were calculated using a mixing-on-sites model and do not show a large variation, either between or within samples. Therefore, an average aNeph for each sample was used. Input parameters and results are summarized in Table 7. Calculations using minimum and maximum values for aJd and aAb resulted in four positions of the same invariant point that define parallelogram-shaped fields in T–aSiO2 space (Fig. 10). These fields are considered to represent the equilibration conditions of the assemblage.
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For the LS and the PMS, equilibrium conditions vary from 680 to 850°C and aSiO2 values of 0·28–0·41 (Table 7). Temperatures agree well with those obtained by nepheline geothermometry after Hamilton (1961)
, which yields temperatures between 700 and 850°C. Only a few nepheline analyses poor in SiO2 indicate temperatures as low as 500°C, which suggests prolonged re-equilibration (Powell, 1978). Calculations with a higher crystallization pressure of 2 kbar yield slightly higher temperatures and silica activities, but this does not affect the general conclusions of this study. For the XPS, calculations were performed with clinopyroxene compositions from phenocryst cores, phenocryst rims and matrix grains, as it is not clear from petrographic observations which clinopyroxene equilibrated with the nepheline and feldspar. Calculations with the most Mg-rich matrix cpx compositions indicate temperatures of 710–775°C and silica activities of 0·28–0·36. Equilibria with Mg-rich rim compositions yield similar but slightly elevated values of aSiO2 (0·32–0·43) and temperatures of 785–910°C. For both matrix and phenocryst rim compositions, calculated values of T and aSiO2 overlap with the intensive parameters determined for the Layered Syenites. However, equilibria based on typical phenocryst core compositions (aJd = 0·08–0·12) yield temperatures of <550°C at silica activities <0·25. It is very unlikely that the cores grew at such low temperatures, and most probably the clinopyroxene core compositions were not in equilibrium with nepheline and alkali feldspar during growth. Thus, we believe that the clinopyroxene phenocryst cores of the XPS crystallized during a distinct earlier event in the magmatic history, possibly as the only crystallizing phase or at higher P.
The abundant occurrence of cancrinite (Fig. 4f) suggests a considerable build-up in CO2 in the late stages of the solidification of the magma, as cancrinite crystallization temperatures, derived from fluid inclusion data, range from 450 to 650°C at 1·1–1·5 kbar (Sobolev et al., 1974). This enrichment of CO2 might be linked to the forcible emplacement of the carbonatite into the nepheline–syenites.
Oxygen fugacities in the nepheline syenites were estimated from the compositions of reintegrated titanomagnetites, and independent temperature determinations from Ne–Ab–Jd equilibria. In Fig. 11, theoretical oxygen buffer curves (FMQ, HM) are plotted together with their displaced positions appropriate to the activities determined from mineral analyses in this study (FMQ*, HM*, FsMQ*). FMQ* marks the minimum oxygen fugacity in the nepheline syenites, as this is the position of the buffer curve recalculated with mineral compositions typical for the syenites (aSiO2 = 0·37, aMag = 0·43) if olivine with aFa = 1 were present. At high temperatures, Fe–Ti oxides equilibrate readily with a fluid phase because of their fast reaction kinetics in hydrothermal experiments (Buddington & Lindsley, 1964), and it can be inferred that this is true for melts as well. Then, for a given temperature and magnetite–ulvöspinelSS composition, the fO2 of the fluid–melt can be read from the fO2–T graph (Fig. 11). Because this fluid–melt phase, in turn, determines the fO2 of the rock (Buddington & Lindsley, 1964), we can estimate the oxygen fugacities of the nepheline syenitic melts at their equilibration temperatures of 710–910°C with the typical oxide composition Mag60–70Usp30–20. The latter are represented as projections of tie-lines connecting coexisting magnetite–ulvöspinelSS and hematite–ilmeniteSS pairs after Buddington & Lindsley (1964) and they define a dark grey shaded field in Fig. 11. At temperatures appropriate for the syenites (700–910°C) fO2 falls between +2 and +5 log units above the FMQ buffer (cross-hatched pattern in Fig. 11). This result agrees well with the position of the FsMQ* (ferrosilite–magnetite–quartz) buffer at very low aFs values (0·01) that corresponds to the virtually absent ferrosilite component in the aegirine–augites. The occurrence of presumably late- to post-magmatic hematite indicates that oxygen fugacities increased to values >1 log unit above the HM buffer during the later stages of magmatic evolution (light grey arrow).
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Stable isotope composition of the Grønnedal-Ika magmas
Oxygen isotopes in nepheline syenites
Oxygen isotope data for mineral separates from Grønnedal-Ika and other Gardar igneous rocks (Halama et al., 2003
; Marks et al., 2003) are plotted in Fig. 12a versus the respective Nd(T) values. Compared with the Eriksfjord Formation basalts, samples from Grønnedal-Ika syenites have slightly lower 18O values at slightly higher values of Nd(T). Clinopyroxene from the gabbroic xenolith overlaps with the EF basalt samples. In contrast, most minerals from the silica-oversaturated rocks from the Puklen intrusion tend towards lower 18O values and strongly negative Nd(T) values, which were interpreted to reflect the influence of open-system late-stage magmatic processes and assimilation of crustal material (Marks et al., 2003). These factors do not seem to have affected the Grønnedal-Ika rocks. However, not only the oxygen isotopic composition of the melt, but also temperature and chemical composition of the minerals influence 18Omineral values (Zheng, 1993a, 1993b). Therefore, it is useful to calculate 18Omelt from the mineral data. The calculation of the oxygen isotopic composition of the original melts from mineral data is relatively straightforward for basaltic rocks (Kalamarides, 1986), but less so for the Grønnedal samples because of lower equilibration temperatures and the formation of diopside–aegirine solid solutions. We therefore used a different approach, which is briefly described below.
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Based on the nepheline–albite–jadeite equilibria, we assume that the equilibration temperature (Teq) for all nepheline syenitic samples is
800°C. In a first step, the mineral–water oxygen isotope fractionation was calculated after Zheng (1993a, 1993b) at Teq using a typical aegirine–augite composition of Di30Hd40Aeg30. Secondly, the rock–water oxygen isotope fractionation of a typical nepheline syenite at Teq was determined after Zhao & Zheng (2003). The two fractionation factors are then combined to derive a value for the mineral–rock oxygen isotope fractionation. For the Grønnedal-Ika nepheline syenites at 800°C, 
cpx–rock is –1·1. 18Omelt for the gabbroic xenolith was calculated using diopside–rock = –0·7, based on the gabbro–water fractionation of Zhao & Zheng (2003) at 1000°C. The overall error introduced by uncertainties in mineral composition and Teq is estimated to be <0·4 and does not affect the conclusions.
Results for the syenitic melts yield an average 18O value of 5·7 ± 0·3 with a range from 5·3 to 6·3 and 5·8 for the gabbroic xenolith (Fig. 12b). These values are indistinguishable from 18O melt compositions calculated for the EF basalts (5·5–6·3; Halama et al., 2003), typical MORB glass compositions (5·37–5·81; Eiler et al., 2000b) and unaltered whole-rock MORBs (5·7 ± 0·2; Harmon & Hoefs, 1995) and indicate a mantle derivation for the nepheline syenites with insignificant crustal contamination (Table 5, Fig. 12b). In comparison with 18Owhole-rock data from the Grønnedal-Ika syenites, which range from 8·3 to 15·1 (Pearce et al., 1997), the calculated 18Omelt values are significantly lower and more homogeneous. In our view, the small scatter of the mineral oxygen isotope data compared with the whole-rock data suggests that the isotopic signatures in the minerals are primary, whereas the whole-rock data are suspected of having undergone late- to post-magmatic alteration. Products of late- to post-magmatic alteration are presumably concentrated along crystal boundaries, which are the parts of the sample removed during sample preparation for mineral analysis.
Carbon and oxygen isotope compositions of carbonatites
Based on the very small differences between measured whole-rock C and O isotope values and carbonate phenocrysts from the Oldoinyo Lengai volcano, Keller & Hoefs (1995) concluded that the isotopic fractionation between carbonate phenocrysts and residual melt is very small. Therefore, we can assume that 13C and 18O values in calcite are roughly equal to those in the carbonatite melt and thus comparable with the whole-rock data. This is confirmed by the overlap of published calcite and whole-rock compositions (Pearce et al., 1997; Taubald et al., 2004). The C and O isotope data are compatible with a primary magmatic, mantle-derived carbonatite (Fig. 8), i.e. largely unaffected by superficial secondary processes, although some minor late- to post-magmatic alteration is possible (Taubald et al., 2004).
Carbonatitic dykes from the Igaliko dyke swarm have a larger range in 18Owhole-rock values (2·5–9·2; Pearce & Leng, 1996) than the Grønnedal carbonatites. However, the average of 4·7 ± 1·5 is significantly lower, possibly because of a high-T meteoric water influence (Fig. 8).
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DISCUSSION |
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Carbonatite petrogenesis
In many cases, silica-undersaturated mantle melts contain carbonates (e.g. Riley et al., 1996
), and such magmas may exsolve an immiscible carbonate melt at crustal or deeper levels (e.g. Freestone & Hamilton, 1980; Kjarsgaard, 1998) or potentially evolve towards carbonatitic compositions via extensive fractionation (e.g. Lee & Wyllie, 1994). Several alkaline complexes have been interpreted to contain such so-called primary carbonatites (Gittins, 1989). The term ‘primary carbonatite’ implies a near-solidus partial melt in equilibrium with CO2-rich peridotitic mantle (Eggler, 1989). Primary carbonatites are characterized by elevated Mg-numbers (Sweeney, 1994), high (Mg + Fe)/Ca ratios, moderate amounts of alkalis (Eggler, 1989) and must have compositions dominated by calcic dolomite (Lee & Wyllie, 1998). This essentially excludes calciocarbonatites, as candidates for primary carbonatitic magmas and the whole-rock carbonatite data of Pearce et al. (1997) indicate that none of the Grønnedal-Ika samples satisfy these criteria because they are characterized by low amounts of MgO (<0·8 wt %) and alkalis (<0·5 wt %). Thus, the geochemical data do not favor a primary, unfractionated origin for the Grønnedal-Ika carbonatites. The syenites also do not have primary mantle–melt compositions, but compositions that indicate extensive crystal–liquid differentiation processes at crustal levels (Bühn & Trumbull, 2003).
Some experimental data (e.g. Lee & Wyllie, 1994) support a fractionation model for the genesis of carbonatites. However, the experimental residual liquids of carbonated alkali silicate melts were fairly silica-rich and not carbonatitic in composition, indicating that fractionation alone is not a feasible mechanism for generating carbonatites from alkali silicate melts at low pressures of 0·2 and 0·5 GPa (Kjarsgaard, 1998). Furthermore, fractionation alone cannot generate the Nb and REE concentrations characteristic of carbonatites (Srivastava, 1997). At Grønnedal-Ika, clinopyroxene is the principal mafic mineral in the syenites and its crystallization will reduce the CaO content in the remaining melt, thus inhibiting the formation of large amounts of calcite. There is also a suggestive lack of mineralogical and chemical continuity between the syenites and carbonatites at Grønnedal-Ika, which is inconsistent with a fractional crystallization relationship (Le Bas, 1989). Therefore, most arguments speak against an origin of the Grønnedal-Ika carbonatites solely by fractional crystallization.
Liquid immiscibility
Immiscible carbonate-rich magmas separated from silicate magmas tend to have calciocarbonatite compositions (Lee & Wyllie, 1998). The strongly SiO2-undersaturated, peralkaline character of the Grønnedal-Ika nepheline syenites resembles the composition of the silicate liquids in early liquid immiscibility experiments (Koster van Groos & Wyllie, 1973). Based on the higher melt–solid interfacial energy of an immiscible carbonate melt compared with the coexisting silicate melt and the resulting restricted migration capability, Minarik (1998) argued that carbonatite should be the last phase in the temporal development of a syenite complex, which is clearly the case in Grønnedal-Ika.
Further support for liquid immiscibility comes from the overlap in Nd(T) values between the syenites and carbonatites (Table 6, Fig. 12). Published Sr isotope data also show similar initial 87Sr/86Sr ratios for the two rock types (Pearce et al., 1997). Significant contamination with crustal material seems unlikely, as the local Archean gneiss country rock has strongly negative Nd(T) values that would cause a shift towards lower Nd. The Nd isotope data show that the Grønnedal-Ika rocks are derived either from an isotopically slightly depleted (relative to Bulk Silicate Earth), homogeneous mantle source or a mixture of at least two mantle end-members, one isotopically depleted and one less depleted or enriched. The identical Nd(T) values of carbonatites and syenites would require melting of similar mantle end-members in the same proportions by different melting events, which is rather unlikely. Thus, the simplest explanation of the Nd isotope data is one involving magmatic differentiation. A liquid immiscibility origin between carbonatites and lamprophyres, as suggested by Pearce et al. (1997), has recently been re-evaluated by Coulson et al. (2003). Based on new geochemical data from the whole Gardar Province, they concluded, in agreement with previous work (Pearce & Leng, 1996; Pearce et al., 1997), that carbonatitic and lamprophyric magmas were derived from the same mantle source, but explained differences between carbonatites and lamprophyres by differing degrees of partial melting of this source. The Nd isotope data from Grønnedal-Ika are in agreement with this conclusion and further suggest that the syenitic rocks are also derived from a similar source.
The oxygen isotopic fractionation between immiscible silicate and carbonate melts is unknown from direct experiments, but thought to be close to the fractionation between, for example, pyroxene and calcite (Santos & Clayton, 1995). At temperatures between 800 and 900°C, calcite–cpx varies between 1·0 for aegirine and 2·1 for diopside (Chiba et al., 1989; Zheng, 1993a). Using these values and following the arguments above, we can roughly estimate the oxygen isotopic composition of the immiscible carbonatite melt from the calculated nepheline–syenitic melt composition of 5·7. It should be between 6·7 and 7·8, within the middle of the available 18Ocarbonatite melt data (Fig. 12b). Thus, the oxygen isotope data are consistent with a petrogenetic model for the origin of the carbonatites involving liquid immiscibility. This is in contrast to the conclusion of Pearce et al. (1997), who argued that the difference in 18O values between the carbonatites and the whole-rock syenitic samples contradicts a liquid immiscibility origin.
Trace-element abundances in calcite
Experimental data on the partitioning of trace elements between immiscible carbonate and silicate liquids at 0·8 and 0·9 kbar (Veksler et al., 1998b) can be used to evaluate whether certain trace-element characteristics of carbonatitic and syenitic calcite can be explained by this process. Types B and C calcite are not considered here because of their unusual chemistry and possible alteration. The similar patterns and relative enrichment levels of calcite from sample GM 1502 (type A) and the whole-rock carbonatites indicate that the relative enrichment of REE in the Grønnedal-Ika carbonatites is controlled by calcite (Fig. 5a). Type A calcite (sample GM 1502) is especially rich in Sr, Ba and P. These compositional features were also found in experimentally produced immiscible carbonatite liquids at low P (0·5 kbar) by Kjarsgaard (1998), suggesting a similar origin for the Grønnedal carbonatites. These liquids were considered as good analogues for calciocarbonatites (Kjarsgaard, 1998), as they were also high in CaO and low in alkalis. The relatively high Ba contents of type A calcite compared with calcite from the syenites is in agreement with experimental partitioning data between silicate and carbonate melts that show a strong preference of Ba for the carbonate phase (Veksler et al., 1998b). Negative Eu anomalies and lower Sr contents in calcite from the syenites are likely to be because of feldspar precipitation. The experimental data also indicate a relative preference of the LREE over the HREE for the carbonate liquid, which is qualitatively consistent with the higher LaCN/YbCN ratios in calcite from the carbonatite compared with calcite from the syenites. However, the high absolute REE concentrations in the carbonatites are difficult to reconcile with liquid immiscibility experiments, which show a preference for most REE in the silicate liquid (Veksler et al., 1998b). Nevertheless, it was argued that the REE are preferentially partitioned into the carbonate liquids because the REE are preferentially transported by (CO3)2–- and F–-complexes that are more easily formed in carbonatitic liquids (Cullers & Medaris, 1977; Möller et al., 1980). Additionally, whole-rock trace-element compositions of the syenites and carbonatites are also consistent with a petrogenetic origin involving liquid immiscibility (Bedford, 1989; Pearce et al., 1997).
Crystallization of the nepheline syenites
Differences in the chemical and petrological evolution of alkali silicate complexes, with and without associated carbonatite, have previously been related to different CO2 contents because of melt extraction from different mantle portions on a small scale or different degrees of melting (Bühn & Trumbull, 2003). Resulting compositional differences could be related to a process of silicate–carbonate liquid immiscibility (Bühn & Trumbull, 2003). Therefore, we will compare the petrology of the Grønnedal-Ika syenites with similar rocks from complexes of the Gardar Province that comprise no carbonatites (Fig. 1b), including silica-undersaturated rocks from Igdlerfigssalik (Powell, 1978), Ilímaussaq (Larsen, 1976; Marks & Markl, 2001; Markl et al., 2001), South Qôroq (Stephenson, 1973), Motzfeldt (Jones, 1984) and silica-oversaturated rocks from the Puklen intrusion (Marks et al., 2003).
T–aSiO2–fO2 conditions
In general, the Layered Syenites, the PMS and the matrix of the XPS equilibrated under very similar T–aSiO2 conditions. The calculated silica activities of the Grønnedal-Ika syenites are distinctly lower than in similar rocks from other intrusive complexes of the Gardar Province not associated with carbonatites, such as the early augite syenite unit of Ilímaussaq (aSiO2 = 0·4–0·9; Marks & Markl, 2001) and the syenitic rocks of Puklen (aSiO2 = 0·7–1; Marks et al., 2003). Oxygen fugacities during the early magmatic stage appear to be distinctly higher in Grønnedal-Ika than in syenitic rocks from Ilímaussaq (1–5 log units below FMQ; Marks & Markl, 2001), Igdlerfigssalik (2–3·5 log units below FMQ; Powell, 1978) and Puklen (1–3 log units below FMQ; Marks et al., 2003). The almost complete absence of olivine and aenigmatite, which occurs in other Gardar complexes, and the differences in the onset of Na enrichment in clinopyroxenes (Fig. 3) is probably also related to differences in oxygen fugacity (Bedford, 1989). In the Ilímaussaq augite syenite, there is also no indication of more oxidizing conditions during late-magmatic stages (Marks & Markl, 2001) (Fig. 11). However, fO2 increases during cooling relative to the FMQ buffer curve towards values above the HM buffer in the agpaitic magmas of Ilímaussaq (Markl et al., 2001) as well as in syenitic and granitic magmas of Puklen (Marks et al., 2003). These trends are similar to the one observed in the Grønnedal-Ika syenites.
Based on the argument that the silica-undersaturated character of mafic alkaline lavas can be attributed to low degrees of partial melting under CO2-rich conditions (Mysen & Boettcher, 1975; Wyllie & Huang, 1976; Gerlach et al., 1988), the very low silica activities of the Grønnedal-Ika syenites may indicate a high CO2 content in the parental magmas which could increase during magma differentiation. Build-up of CO2 is also indicated by the occurrence of calcite and cancrinite in the nepheline syenites. Further constraints on fCO2 can be derived from the equilibrium CH4 + 2O2 = CO2 + 2H2O at variable fCH4/fCO2 ratios in T–fO2 space (Fig. 11). The reaction curve approaches conditions estimated for the Grønnedal syenites from the titanomagnetite compositions only for fCH4/fCO2 < 0·1. This is consistent with the commonly observed dominance of CO2–H2O fluid inclusions in carbonatites (Andersen, 1986; Samson et al., 1995; Bühn & Rankin, 1999). On the other hand, high fCH4/fCO2 ratios of >0·9 are indicated for the Ilímaussaq syenites, which is consistent with the presence of methane-bearing fluid inclusions with low CO2 contents (Konnerup-Madsen & Rose-Hansen, 1982). A high fugacity of CO2 would be an important prerequisite for carbonatite formation at Grønnedal-Ika and a different fCO2 in the Grønnedal-Ika parental melt could be the most important difference from the Ilímaussaq parental melt.
Trace elements in clinopyroxene
A qualitative assessment of the trace-element patterns in the Grønnedal clinopyroxenes by comparison with clinopyroxenes from other Gardar complexes not associated with carbonatite reveals no significant differences in trace-element contents or patterns (Fig. 6b; Marks et al., 2004a). It therefore appears that crystal-chemical effects dominate the trace-element partitioning behavior (Wood & Blundy, 1997; Blundy & Wood, 2003), and other effects, such as temperature and melt chemistry, play a subordinate role.
Ratios of geochemical twin elements, such as Zr/Hf and Y/Ho, can be used as petrogenetic indicators for comparing intrusions with and without associated carbonatites. Bühn & Trumbull (2003) found, for instance, significantly higher Zr/Hf ratios in syenitic rocks associated with carbonatite compared with ones without and related these differences to the ability of CO2-rich fluids to fractionate Zr from Hf (Dupuy et al., 1992; Irber, 1999). The Zr/Hf ratios in the clinopyroxene cores from the XPS overlap with those of other Grønnedal-Ika nepheline syenites (Zr/Hf = 31–43) and with those from augite syenitic units at Ilímaussaq (Zr/Hf = 29–34). The higher Zr/Hf ratios in the phenocryst rims are not likely to be because of fractionation of clinopyroxene, as both elements are mildly incompatible and the effect on Zr/Hf fractionation is negligible when only one ferromagnesian mineral is present (Linnen & Keppler, 2002). Experimental evidence also shows that there is little or no Zr/Hf fractionation in alkaline depolymerized melts (Linnen & Keppler, 2002). The lower Y/Ho ratios in phenocryst rims (Fig. 7) are consistent with exsolution of a CO2-dominated fluid, which would leave the melts with successively decreasing Y/Ho ratios (Bühn & Trumbull, 2003). These differences in Zr/Hf and Y/Ho ratios suggest that clinopyroxene cores might have grown before, and rims during or after the separation of a CO2-rich fluid. A pure crystal-chemical control on these trends seems unlikely because the major element compositions of clinopyroxene rims in the XPS and matrix clinopyroxene in the Layered Syenites are very similar.
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SUMMARY AND CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONPREVIOUS WORK AND GEOLOGICAL...ANALYTICAL METHODSPETROGRAPHY AND MINERAL...GEOCHEMISTRYCALCULATION OF PHYSICO-CHEMICAL...DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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The aim of the present study of the Grønnedal-Ika complex was to explore the relationship between the syenitic and carbonatitic rocks, to constrain the mode of origin of the carbonatite and to investigate what makes the syenitic rocks distinct from other Gardar complexes. We conclude that most of the available data indicate an origin of the carbonatite via liquid immiscibility from a parental silicate liquid. Criteria that support this model are listed below.
- Field and petrographic evidence suggests that there was no significant time gap between the emplacement of the silicate rocks and the carbonatites.
- Whole-rock and mineral geochemical fractionation indices such as Mg-number and (Mg + Fe2+)/Ca confirm that the carbonatites do not represent unfractionated mantle melts. Clearly, the syenitic silicate rocks also crystallized from fractionated melts.
- Radiogenic isotope compositions (Sr: Pearce et al., 1997; Nd: this study) are similar for silicate rocks and carbonatites.
- Oxygen isotope ratios for syenitic melts, determined from a set of mafic mineral data using appropriate fractionation factors between minerals and whole rocks (Zhao & Zheng, 2003), are in the range of typical mantle values (18O = 5·7) and roughly 2 below the carbonatite melt values. This difference in 18O is within the range of theoretical expectations.
- In agreement with liquid immiscibility experiments (Veksler et al., 1998b), the normalized REE patterns of calcite from carbonatites have steeper slopes than calcite from the associated silicate rocks. The same is true for carbonatite whole rocks compared with the syenites.
- The precipitation of aegirine–augitic clinopyroxene in the XPS with elevated Zr/Hf and lower Y/Ho ratios are consistent with effects that one would expect from the exsolution of a CO2-rich fluid–melt. This may suggest that the XPS represents the conjugate silicate liquid in an immiscibility process.
- The silicate rocks associated with the carbonatites are strongly silica-undersaturated.
- The discreteness of occurrence favors liquid immiscibility rather than a fractional crystallization relationship in which continuous series of silicate to carbonate rocks would be expected (Le Bas, 1989).
- The carbonatites represent the latest phase in the complex (Minarik, 1998).
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ACKNOWLEDGEMENTS |
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We would like to acknowledge Henry Emeleus, Kathryn Goodenough and Robert Trumbull, as well as Marjorie Wilson for their constructive reviews, which greatly improved the paper. We are also grateful to Bruce Paterson, who provided invaluable help during laser ICP–MS measurements at the Large-Scale Geochemical Facility supported by the European Community–Access to Research Infrastructure action of the Improving Human Potential Programme, contract number HPRI-CT-1999–00008 awarded to Professor B. J. Wood (University of Bristol). Thanks to the courtesy of Kathryn Goodenough, Chris Bedford and Heinrich Taubald, we were able to include some of their unpublished data into this work. Additional thanks go to Chris Bedford for making available a copy of his voluminous Ph.D. thesis via Henry Emeleus. Gabi Stoschek, Bernd Steinhilber and Elmar Reitter expertly assisted with stable and radiogenic isotope measurements. Thanks go to Thomas Wenzel for his help with microprobe measurements and his critical comments that helped to improve the manuscript. Michael Marks is thanked for his pleasant company during fieldwork. The naval station at Grønnedal provided valuable logistical support. Financial funding of this work by the Deutsche Forschungsgemeinschaft (grant Ma-2135/1–2) is gratefully acknowledged.
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FOOTNOTES |
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* Present address: Institut de Physique du Globe de Paris, Case courrier 109, 4 Pl. Jussieu, 75005 Paris, France.
Corresponding author. Telephone (+49) 7071/2972930. Fax: (+49) 7071/293060. E-mail: markl{at}uni-tuebingen.de
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REFERENCES |
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TOPABSTRACTINTRODUCTIONPREVIOUS WORK AND GEOLOGICAL...ANALYTICAL METHODSPETROGRAPHY AND MINERAL...GEOCHEMISTRYCALCULATION OF PHYSICO-CHEMICAL...DISCUSSIONSUMMARY AND CONCLUSIONSREFERENCES |
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