Journal of Petrology | Volume 44 | Number 1 | Pages 3-38 | 2003
© Oxford University Press 2003
Alkali Picrites Formed by Melting of Old Metasomatized Lithospheric Mantle: Manîtdlat Member, Vaigat Formation, Palaeocene of West Greenland
1GEOLOGICAL SURVEY OF DENMARK AND GREENLAND, ØSTER VOLDGADE 10, DK-1350 COPENHAGEN K, DENMARK
2GEOLOGICAL MUSEUM, ØSTER VOLDGADE 5–7, DK-1350 COPENHAGEN K, DENMARK
3MINERALOGISK–GEOLOGISK MUSEUM, SARS GATE 1, N-0562 OSLO, NORWAY
4GEOLOGICAL INSTITUTE, UNIVERSITY OF COPENHAGEN, ØSTER VOLDGADE 10, DK-1350 COPENHAGEN K, DENMARK
5DANISH LITHOSPHERE CENTRE, ØSTER VOLDGADE 10, DK-1350 COPENHAGEN K, DENMARK
RECEIVED December 5, 2001; ACCEPTED July 1, 2002
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESPETROGRAPHY AND MINERALOGYWHOLE-ROCK GEOCHEMISTRYDISCUSSIONCONCLUSIONSAPPENDIX: CALCULATION OF MELTING...SUPPLEMENTARY DATAREFERENCES |
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Alkaline picrites and basalts constitute 20–200 m of lava flows and hyaloclastites in the middle part of an
2 km thick succession of tholeiitic picrites and basalts formed during continental rifting of West Greenland around 60 Ma. The alkaline rocks, found only in northern Disko, have phenocrysts of olivine + chromite ± clinopyroxene; lava flows contain abundant groundmass clinopyroxene and plagioclase, whereas pillow breccias contain abundant fresh alkali basaltic glass. Six compositional types are present; all are strongly but variably enriched in incompatible trace elements [Ba, U, Nb, Ta, light rare earth elements (LREE)], yet their major elements, with relatively high SiO2 and Al2O3 and low Na2O, do not suggest an origin by small degrees of mantle melting. The isotope compositions are unusual, with negative 
Nd and mostly negative Sr (below the mantle array), high 206Pb/204Pb (below the Northern Hemisphere Reference Line), and mostly negative 
Os. The most likely source for the alkaline magmas is old metasomatized lithospheric mantle in which melting was induced by the passing hot, asthenosphere-derived, tholeiitic magmas. Simple mass-balance calculations suggest that the melting assemblages consisted of 60% pargasitic amphibole, 26–30% clinopyroxene, 9% olivine and 1% apatite. Mica in the source is required for only the least enriched magma type. For the most enriched magmas small amounts of Ba–U–Nb–Sr–LREE-rich oxides (lindsleyite and hawthorneite) are required in the melting assemblage and dominate the Pb isotope compositions. The various magma types and the partly complementary relation between them suggest that the lithospheric mantle had an ordered structure, possibly with old metasomatic zones formed by successive trapping of elements in migrating fluids. KEY WORDS: alkali picrite; amphibole melting; Greenland; lithosphere melting; metasomatism
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESPETROGRAPHY AND MINERALOGYWHOLE-ROCK GEOCHEMISTRYDISCUSSIONCONCLUSIONSAPPENDIX: CALCULATION OF MELTING...SUPPLEMENTARY DATAREFERENCES |
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Alkaline extrusive rocks are a volumetrically insignificant component in most large igneous provinces (LIPs). LIPs are overwhelmingly tholeiitic in character and are considered to be formed by relatively high degrees of melting, mainly of the asthenosphere, and usually in the presence of a mantle plume (e.g. Saunders et al., 1997). The alkaline magmas are generally considered to be products of lower degrees of melting; however, their mantle sources, whether asthenospheric, lithospheric, or both, are contentious. Alkaline extrusives have often been emplaced late in the formation of the LIPs, e.g. the meimechites capping the Siberian basalt plateau (Arndt et al., 1995, 1998), and the nephelinites and basanites capping the East Greenland basalt plateau and nunataks (Brooks et al., 1979; Brown et al., 1996; Bernstein et al., 2000). Both the Siberian and the East Greenland alkaline extrusives have been interpreted to be of asthenospheric origin (Brooks et al., 1979; Arndt et al., 1995, 1998; Brown et al., 1996; Bernstein et al., 2000) whereas in other LIPs, such as the Deccan and Yemen, lithospheric components are thought to be involved (e.g. Mahoney et al., 1985; Baker et al., 1997). Probably, the alkaline rocks in LIPs are polygenetic.
During Palaeogene rifting and continental break-up in the North Atlantic, large volumes of flood basalts were extruded on the continental margins of both West and East Greenland. The magmas are thought to be generated mainly by melting within the impacting Iceland mantle plume (e.g. review by Saunders et al., 1997). In contrast to East Greenland, the West Greenland flood basalts do not include any nephelinites but terminate with transitional to mildly alkaline basalts (Clarke & Pedersen, 1976; Larsen, 1977). The voluminous main succession is uniformly tholeiitic except within one limited interval in the middle part, which, in addition to tholeiitic picrites, contains three close-lying levels with alkali picrites and alkali basalts with distinctive and highly unusual geochemical and isotopic characteristics. This paper explores the petrogenesis of the alkaline melts and the nature of their unusual mantle sources.
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GEOLOGICAL SETTING |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESPETROGRAPHY AND MINERALOGYWHOLE-ROCK GEOCHEMISTRYDISCUSSIONCONCLUSIONSAPPENDIX: CALCULATION OF MELTING...SUPPLEMENTARY DATAREFERENCES |
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The West Greenland basalts form part of the Nuussuaq Basin, a fault-controlled extensional basin at the continental margin where Precambrian basement is covered by Cretaceous–Palaeocene sediments overlain by volcanic rocks (Fig. 1; Chalmers et al., 1999). The major part of the onshore volcanic succession was erupted during a short time period around 60 Ma (Storey et al., 1998). The lower part of the 2–4 km thick volcanic succession is dominated by highly magnesian picritic rocks (the Vaigat Formation), whereas the upper part consists of more evolved, plagioclase-phyric basalts of the Maligât Formation (Clarke & Pedersen, 1976). The magnesian magmas of the Vaigat Formation were generated at high temperatures and very high production rates in the asthenosphere and passed swiftly through the lithosphere (Gill et al., 1992; Holm et al., 1993; Larsen & Pedersen, 2000; Pedersen et al., 2002). Most magmas escaped contamination, although a number of discrete crustal contamination episodes led to the formation of subordinate units of siliceous basalts and magnesian andesites (Pedersen, 1985a; Pedersen et al., 1996; Lightfoot et al., 1997). The volcanism of the Vaigat Formation occurred in three main cycles which formed the three main stratigraphic members; from older to younger these are the Anaanaa, Naujánguit and Ordlingassoq Members. The interval with alkaline rocks is the upper Naujánguit Member to lower Ordlingassoq Member, and the main part of the alkaline rocks forms one stratigraphic unit, which is formalized as the Manîtdlat Member (Pedersen, 1985b). The geographical distribution of these rocks is limited to northern Disko (Fig. 1).
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Alkaline volcanic rock units
Alkaline rocks occur at three close-lying stratigraphic levels (Fig. 1). They are divided into a number of types as shown in Table 1. The total estimated volume of the alkaline rocks is around 30 km3; with possible extensions to the west and north the original volume may have been up to 50 km3, about 0·05% of the original volume of onshore basalts.
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The oldest alkaline unit is a volcanic neck with a few associated lava flows of alkali basalt in the Stordal area (Fig. 1). These rocks (Stordal type) are of very limited extent and volume and are interbedded within picritic lava flows of the uppermost Naujánguit Member.
The second alkaline unit is an up to 20 m thick series of olivine-rich alkali picrite flows within the lowermost part of the Ordlingassoq Member. This unit (Type 0) can be followed over 20 km along the north coast of Disko and was most probably produced from one volcanic centre. It is overlain by 50–60 m of tholeiitic picrites, which are in turn overlain by the rocks of the Manîtdlat Member.
The third alkaline unit is the Manîtdlat Member, which forms a purplish brown marker horizon within the grey tholeiitic picrites of the Ordlingassoq Member (Fig. 2). The Manîtdlat Member is found within a 30 km by 20 km area in northern Disko (Fig. 1). It is on average about 50 m thick and represents eruptions from several volcanic centres, most probably fissure eruptions because all the lava flows are of pahoehoe type and no traces of volcanic edifices or explosive activity have been found. No eruption sites are known but they must be local. The volcanic rocks were produced from at least four alkaline centres and two tholeiitic centres, which interfinger laterally. The alkaline rocks are picrites and alkali basalts with four element enrichment patterns (Types 1a, 1b, 2 and 3; see Table 1). The youngest rocks are clinopyroxene-phyric basalts (ankaramites, Type 3), probably erupted from a volcanic centre in the southern part of the area. A dyke cutting the whole Manîtdlat Member is tholeiitic but somewhat enriched. Overlying the Manîtdlat Member are tholeiitic picrites of the Ordlingassoq Member.
Most of the volcanic rocks were erupted subaerially. However, the volcanic front was prograding laterally towards the east and SE into a volcanic-dammed lake (Pedersen et al., 1998), so that the thin lava flows pass laterally into thick hyaloclastite breccia deposits (Fig. 2). Consequently, all the magma types discussed here except the Stordal type exist in hyaloclastite facies with extremely fresh glassy rocks; in contrast, many of the subaerial lava flows are affected by zeolite-facies metamorphism. Stordal type glass is found in a chilled neck contact. Mantle xenoliths have not been found despite dedicated search.
In this paper, the alkaline units will be collectively referred to as the Manîtdlat Member.
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ANALYTICAL TECHNIQUES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESPETROGRAPHY AND MINERALOGYWHOLE-ROCK GEOCHEMISTRYDISCUSSIONCONCLUSIONSAPPENDIX: CALCULATION OF MELTING...SUPPLEMENTARY DATAREFERENCES |
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Microprobe analyses of phenocrysts, glass inclusions and matrix glasses were made on a JEOL Superprobe at the University of Copenhagen, using a combination of wavelength-dispersive (WDS) and energy-dispersive (EDS) detection systems. Normal operating conditions were 15 kV acceleration voltage, 15 nA beam current, 20 s total counting time for WDS and 60 s live time for EDS analyses. High-precision analyses of Ca, Cr, Ni and Ti in olivine and Ni, Ti and V in chromite were made using 100 nA beam current and 40 s total counting time (WDS); the lower limits of detection for these conditions are 15–30 ppm (Pedersen, 1985a). Glasses were analysed with 15 nA beam current and extended total counting times (WDS) for Ti (60 s), K (60 s) and P (120 s), yielding reproducibilities of 0·03 wt % TiO2 and 0·02 wt % K2O and P2O5 (2
on 11 repeat analyses on glass standards). Sixty whole-rock samples from the alkaline units and 15 samples of the contemporaneous tholeiitic picrites were analysed for major and trace elements by X-ray fluorescence spectrometry (XRF). Major elements were determined on fused glass discs at the Geological Survey of Denmark and Greenland, as described by Kystol & Larsen (1999). FeO was determined by titration. Trace elements were determined on pressed powder pellets at the Geological Institute, University of Copenhagen, using a Philips PW 1400 spectrometer and standard analytical methods with USGS reference materials for calibration.
A subset of 13 alkaline and four tholeiitic samples, mainly fresh pillow breccias, was analysed for trace elements by inductively coupled plasma mass spectrometry (ICP-MS) and for Sr, Nd and Pb isotopes. The ICP-MS analyses were performed on a Perkin–Elmer SCIEX Elan 6000 at the University of Durham, using methods described by Turner et al. (1999). Reproducibility, based on replicate digestion of samples, varied from 1·5% to 3% for most analyses.
Sr, Nd and Pb isotope ratios were determined on unleached samples on a VG354 instrument at the Mineralogisk–Geologisk Museum, Oslo, using methods described by Griffin et al. (1988). Average values for repeated standard analyses during the analytical period were 87Sr/86Sr = 0·71023 ± 3 (2 SE) for NBS987 and 143Nd/144Nd = 0·511112 ± 5 (2 SE) for the J&M standard batch no. S819093A. The Pb standard NBS981 gave 206Pb/204Pb = 16·897 ± 0·005, 207Pb/204Pb = 15·434 ± 0·005 and 208Pb/204Pb = 36·540 ± 0·015 (2 SE).
Seven samples were analysed for Os isotopes at Geocentre Copenhagen. Samples were spiked with an 188Os- and 187Re-enriched solution and digested in inversed (14N HNO3:10N HCl = 3:1) aqua regia in carius tubes at 230°C for 1 week (Shirey & Walker, 1995). Os was distilled from aqua regia directly into 8N HBr (Nägler & Frei, 1997) and purified following Roy-Barman & Allègre (1994). Os isotope analyses were performed on a VG Sector 54 solid-source negative thermal ionization mass spectrometer, using a multi-collector static routine and single multiplier peak jump mode for small Os beams. Re was purified using the liquid extraction method of Cohen & Waters (1996) and the concentrations were measured by multiple-collector (MC)-ICP-MS on an Axiom instrument, using Ir-doped sample solutions for controlling mass fractionation of Re through monitoring the 190Ir/194Ir ratio. Procedural blanks for Re were <30 pg and for Os <3 pg.
Mössbauer analyses were performed on handpicked glass chips at the Royal Veterinary and Agricultural University, Copenhagen. The spectra were obtained at 295 and 80 kV, using a constant acceleration spectrometer, and were fitted using three Fe2+ doublets and one Fe3+ doublet.
Analytical results are presented in Tables 2–6. The complete dataset is available for downloading from the Journal of Petrology web site at http://www.petrology.oupjournals.org (Electronic Appendix A).
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PETROGRAPHY AND MINERALOGY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESPETROGRAPHY AND MINERALOGYWHOLE-ROCK GEOCHEMISTRYDISCUSSIONCONCLUSIONSAPPENDIX: CALCULATION OF MELTING...SUPPLEMENTARY DATAREFERENCES |
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The alkaline rocks have simple mineralogies with olivine and chromite as the earliest phenocryst phases, just as in the contemporaneous tholeiites (Larsen & Pedersen, 2000). In contrast to the tholeiites, clinopyroxene was the next phase to crystallize, followed by plagioclase. The glassy rocks are 1–3 cm thick pillow rims with phenocrysts (up to 1–2 mm) and microphenocrysts embedded in 70–80 vol. % clear, pristine, pale yellow glass, which has yielded well-defined 39Ar–40Ar ages of 60·3 ± 1·0 Ma and 60·7 ± 0·5 Ma (Storey et al., 1998). In the glassy rocks the amount of clinopyroxene increases gradually with decreasing Mg content of the rocks, from none or just a few tiny microphenocrysts in the glasses of the most Mg-rich rocks (Type 0), through increasingly frequent microphenocrysts, often in clusters, in Type 1 and 2 glasses, to up to 1 mm phenocrysts in the glasses of Type 3 where clinopyroxene is the dominant phenocryst phase. Plagioclase is absent from the glassy rocks of Type 0, and it forms microlites in the Type 1 and 2 glasses, and sparse microphenocrysts in the Type 3 glasses. The Stordal rocks are aphyric to olivine–clinopyroxene microphyric.
In the crystalline lava samples the groundmass consists of clinopyroxene, plagioclase, olivine, Fe–Ti oxides and apatite in intersertal, intergranular or subophitic textures. No primary mica or amphibole has been found. The alkaline and calcic character is reflected in very high modal proportions of clinopyroxene (17–40% normative di), with crystals often showing hourglass zoning and purplish colours. Late-stage segregation veins contain purple, zoned, prismatic clinopyroxene crystals, zoned plagioclases often heavily zeolitized, semi-skeletal magnetite and ilmenite and frequent apatite crystals, all embedded in a matrix of fine-grained zeolite–smectite aggregates. In contrast to the fresh glassy pillow breccias, the crystalline samples are affected by zeolite-facies metamorphism, and interstices are filled with green and brown smectites and colourless fine-grained aggregates of zeolites and Ca-hydrosilicates. These interstices probably include the breakdown products of nepheline. Vesicles are filled with massive zeolites. Sulphides form secondary pyrite in alteration zones, and primary sulphide liquid drops infrequently preserved as tiny 1–5 µm spherules in glass inclusions in olivine phenocrysts in the Mg-rich Types 0 and 1a.
Olivine
Olivine comprises several textural types similar to those from the tholeiitic rocks described by Larsen & Pedersen (2000). Most olivines are clear euhedral to subhedral to skeletal phenocrysts; some have inclusion-filled zones and healed cracks, and others have cores speckled with numerous small inclusions of oxides and sometimes glass. Some olivines are obviously xenocrystic, with anhedral and serrated outlines.
The olivines span the compositional range mg-number 92·3–77·4 (Table 2) with a compositional gap around mg-number 90. The most magnesian olivines (mg-number >90) are found in the most magnesian rocks and there is a crude correlation between the olivine compositional range within a sample and the bulk-rock MgO contents, as also found in the tholeiitic rocks (Larsen & Pedersen, 2000). All the olivines, including those with mg-number >90, have glass inclusions and high contents of CaO and Cr2O3, indicating a magmatic origin. Possible mantle xenocrysts would have very low contents of CaO and Cr2O3 (Larsen & Pedersen, 2000) and have not been found.
The minor elements MnO, CaO, Cr2O3 and NiO, measured with high precision, show distinct differences between olivines in alkaline and tholeiitic rocks (Fig. 3). First, olivines in the alkaline rocks show a far greater scatter than those in the tholeiites. Second, the main olivine populations in the alkaline rocks have distinctly higher contents of CaO (and MnO, not shown), similar or higher Cr2O3, and lower NiO than the tholeiitic olivines. Within a single sample, olivine crystals with widely different minor-element contents and zoning patterns may exist side by side. A few olivine crystals, often the larger ones, have minor-element contents similar to those of the tholeiitic olivines. This is particularly evident for some olivines with low CaO. The zoning patterns in some individual crystals are also shown in Fig. 3, and the significance of the data is discussed below.
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Oxides
The primary magmatic oxides are brown, semi-transparent chromites, which occur as small euhedral crystals enclosed in olivine and, in Type 0 picrites only, also within the matrix glass. Some olivine cores are speckled with tiny oxide inclusions ranging from dust-size particles to greenish brown bleb-like grains and opaque vermicular grains. Rare opaque oxide xenocrysts, fringed with clinopyroxene crystals, occur free in the matrix glass.
The compositional variation of the oxides is illustrated in Fig. 4 and Table 2. The primary magmatic chromites have cr-numbers [atomic 100Cr/(Cr + Al)] between 40 and 60, generally lower than in the tholeiitic rocks (Fig. 4a). The dyke chromites have high cr-numbers. There is a good correlation between the mg-number of the chromites and the enclosing olivines (not shown); chromites with mg-number 75–80 are enclosed within olivines with mg-number 91–92·5, and the iron-rich olivine xenocrysts contain correspondingly iron-rich chromites.
The primary magmatic chromites show differences in the iron oxidation states between the magma types (Fig. 4b; Table 2). The Type 1a chromites have the highest Fe2+/
Fe whereas those in Type 1b have the lowest. With progressive crystallization (decreasing mg-number) the chromites become more reduced (higher Fe2+/Fe), in contrast to the tholeiitic chromites, which become more oxidized (lower Fe2+/Fe).
The opaque oxide xenocrysts are ferroan magnesioferrites, which have thin rims that are more reduced and more magnesian than the centres. They are extremely low in both Cr and Ti (Table 2, numbers 16 and 17).
The vermicular and bleb-like oxides in some olivine crystals range from ferroan magnesioferrite in olivine cores through greenish bleb-like Al-rich spinel sensu stricto to chromian spinel approaching normal magmatic compositions in the olivine rims (Fig. 4c). These oxides are not magmatic but are solid-state high-temperature oxidation and re-equilibration products, as discussed below.
Clinopyroxene and plagioclase
The clinopyroxenes show complicated oscillatory zoning, which is a feature very commonly found in alkaline rocks. The microphenocrysts generally correspond to the outermost 2–3 zones of the larger phenocrysts. Compositionally, however, all clinopyroxene phenocrysts span a relatively narrow range, En42–48 Fs4–10 Wo42–52, and they are thus classical diopsides (Table 2). There is little or no difference between the clinopyroxenes from the various chemical rock types, except that those of Type 2 tend to have slightly higher Ti and slightly lower Wo.
Plagioclase microphenocrysts with slight normal zoning (in Type 3) and microlites (in Types 1–3) span the compositional range An87–71 Or0·7–2·3. There are no differences between the various rock types.
Primary sulphides
Tiny globules of sulphide preserved within glass inclusions in olivine have chemical compositions close to Fe–Ni monosulphide with small amounts of Cu. Most globules are too small for ‘clean’ microprobe analyses, but energy spectra show about equal amounts of Fe and Ni in sulphide in olivine with mg-number 92, and successively decreasing Ni in sulphides in olivines with mg-number 88–85. Similar globules are also present in the tholeiitic rocks. The globules seem to be too large to be exsolved from the trapped liquid and are considered to be trapped as liquid sulphide together with the silicate melts.
Matrix glasses and glass inclusions
Matrix glasses were analysed in all compositional types. Glass inclusions in olivine phenocrysts, representing melts at earlier stages of crystallization, were analysed in Mg-rich samples of Types 0, 1 and 2. The glass inclusions were not homogenized and have lost olivine as a result of post-entrapment crystallization, but as long as there are no other daughter minerals the incompatible element ratios of the trapped melts should be unaffected.
The alkaline matrix glasses (Table 3) are homogeneous and relatively fractionated, with 6·4–7·9 wt % MgO and elevated contents of the incompatible elements Ti, Na, K and P relative to the bulk rocks. In a K2O–TiO2–P2O5 triangular diagram (Fig. 5), the matrix glasses of the various alkaline types have well-defined K2O–TiO2–P2O5 ratios (Fig. 5a), with Types 0, 1a and 1b being closely similar. The tholeiitic glasses with their very low K2O and P2O5 contents plot close to the TiO2 apex. The inclusion glasses (Fig. 5b) show considerable scatter, but the main trend goes from the matrix glasses towards the TiO2 apex. A plot of CaO/Al2O3 versus the composition of the olivine (Fig. 5c) shows that low-CaO olivines have glass inclusions with low CaO/Al2O3; these are the same as those with high TiO2/P2O5 in Fig. 5b.
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The glass inclusions in the high-CaO olivines have CaO/Al2O3 ratios of 0·9–1·2. Except for Type 0 there is a general tendency of declining CaO/Al2O3 with increasing crystallization (decreasing olivine mg-numbers) to low values in the matrix glasses, distinctly lower than in the bulk rocks. This is ascribed to the formation of the abundant clinopyroxene microphenocrysts in the matrix glasses in these rocks.
Crystallization temperatures and oxidation states of alkaline vs tholeiitic melts
During quenching the last olivine rims that formed in contact with the alkaline matrix glasses have mg-numbers varying from 83·5 in the Type 0 picrites to 79·6 in the Type 3 ankaramites. When the oxidation state of iron in the glass is derived by assuming an olivine–melt Fe–Mg distribution coefficient at 1 atm of 0·30 (Roeder & Emslie, 1970), quench temperatures can be calculated after Ford et al. (1983) and range from 1190°C in the Type 0 picrites to 1150°C for the ankaramites. In most cases the measured and calculated olivine compositions are very similar, indicating equilibrium between olivine rims and glass. In comparison, the quench temperatures for the tholeiites are in the range 1210–1180°C for rocks without plagioclase phenocrysts (Larsen & Pedersen, 2000).
The complicated oxides indicate early intratelluric events of oxidation of the olivine phenocrysts, leading to oxidation-exsolution of magnesioferrite (Khisina et al., 1995), and subsequent solid-state re-equilibration at low oxygen fugacities, leading to formation of Fe3+-poor, Al-rich spinel and then to Fe3+-poor Cr-spinel (Fig. 4c). During both primary chromite crystallization and re-equilibration, the oxides from the alkaline rocks show progressively decreasing oxidation states, in contrast to the evolution trend in the chromites from the tholeiitic rocks (Fig. 4b). Mössbauer analysis of matrix glasses (Table 3) also shows higher Fe2+/Fe in the alkaline than in the tholeiitic glasses. Thus, there is circumstantial evidence that the alkaline melts were significantly more reduced than the tholeiitic melts.
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WHOLE-ROCK GEOCHEMISTRY |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESPETROGRAPHY AND MINERALOGYWHOLE-ROCK GEOCHEMISTRYDISCUSSIONCONCLUSIONSAPPENDIX: CALCULATION OF MELTING...SUPPLEMENTARY DATAREFERENCES |
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Major elements
Representative XRF analyses are shown in Table 4, and plots of the major elements vs MgO are shown in Fig. 6. The alkaline rocks contain 7–23 wt % MgO, and all have much higher contents of P2O5 and K2O than the contemporaneous tholeiites. They form a number of compositional groups (types, presented in Table 1), which are particularly evident in the P2O5 and TiO2 diagrams. Most of the alkaline rocks have low SiO2 and high CaO compared with the tholeiites. Types 0 and 3 show the greatest relative enrichment in CaO and P2O5 but no enrichment in TiO2. In contrast, Type 2 rocks have relatively low CaO and P2O5 but the greatest enrichment in TiO2 and K2O. The alkaline rocks have only slightly lower Al2O3 than the tholeiites, whereas the levels of FeO* are the same or slightly lower. Perhaps surprisingly, the alkaline rocks are not enriched in Na2O; indeed, many have lower Na2O contents than the tholeiites. Although Na2O is somewhat scattered as a result of secondary alteration, low Na2O is a primary feature of the magmas and is also seen in the analyses of the fresh matrix glasses (Table 3 and Fig. 6). The tholeiitic glasses have the same Na2O contents as the glasses of Types 0 and 2; the highest Na2O contents are seen in the glasses of Types 1 and 3 that were quenched after significant clinopyroxene crystallization. The K2O diagram shows considerable scatter, much of which is due to secondary redistribution of K in the lava samples, which, as described above, are often altered whereas the pillow breccias are fresh. The data from the pillow breccias alone strongly suggest that the Type 2 melts were generated with higher K2O than the other types. The dyke has tholeiitic abundances of most of the major elements but has slightly elevated contents of P2O5 and K2O.
Three Stordal samples and four Type 1a samples have increased SiO2 and decreased CaO relative to other rocks of the same type; together with other elemental fingerprints this suggests these samples are crustally contaminated, as discussed below.
For each rock type, the compositional variation seen in Fig. 6 is dominantly caused by olivine fractionation and accumulation; the clinopyroxene-phyric Type 3 rocks also show evidence of clinopyroxene fractionation, or perhaps accumulation, in the changed slope of the CaO trend. CaO/Al2O3 is not changed by olivine fractionation or accumulation, and the different levels of this ratio in the various magma types (Fig. 7) may be features of the primary magmas. Types 0, 1b and 3 have the highest CaO/Al2O3 ratios, up to 1·35, caused by both high CaO and low Al2O3 (Fig. 6). The sloping CaO/Al2O3 trend in the Type 3 rocks is produced by clinopyroxene fractionation or accumulation. Type 2 and Stordal have CaO/Al2O3 ratios corresponding to the tholeiites.
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All samples except for three evolved ones contain <3 wt % total alkalis, and according to the IUGS classification (Le Bas, 2000) the rocks are simply picrites and basalts. The alkaline character is better reflected in the CIPW norms, particularly of the matrix glasses (Fig. 8). None of the Type 2 and Stordal bulk rocks are ne normative; Type 1 is mixed, and all rocks of Type 0 and Type 3 are ne normative although the maximum ne is only 5·3%. However, all matrix glasses except Stordal and the dyke are ne normative, with maximum ne = 9·8–10·2% in Type 3. The three high-Si Stordal lavas are slightly Q normative (not plotted). Figure 8 also shows the high contents of normative diopside, the Type 3 ankaramites attaining a maximum of 40% di. The Type 3 glasses have somewhat decreased di because of clinopyroxene fractionation, whereas the Type 0 glasses with 39% di were quenched just before clinopyroxene saturation was reached.
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Trace elements
Figure 9 shows a range of trace elements plotted against MgO. The incompatible elements Ba, Sr, Nb and La show different levels of enrichment relative to the tholeiites, and Type 1 clearly splits up in two groups (a and b). Types 0, 1b and 3 are the most strongly enriched in Ba, Sr and La, Type 1a is intermediate, and Type 2 and Stordal are the least enriched. Types 0, 1 and 3 have extremely high Nb/Zr ratios in the range 0·5–0·7. In contrast, Type 2 is relatively more enriched in Zr and has lower Nb/Zr ratios of 0·3–0·4, but still higher than the tholeiitic values of <0·14. The dyke is close to the tholeiitic values for many elements but is clearly enriched in all the incompatible elements (see also Fig. 11). It has a Nb/Zr ratio of 0·19.
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The compatible elements Ni and Cr show the effects of olivine and chromite control; however, the alkaline rocks have distinctly lower Ni contents than tholeiites with similar MgO. The diagrams for V and Sc show a maximum at 10–11 wt % MgO, indicating clinopyroxene fractionation in magmas with <10 wt % MgO, in accordance with the major-element variations.
A wider range of trace elements was obtained by ICP-MS analysis of two samples of each of the alkaline types and one of the dyke (Table 5). The two samples from each alkaline type gave closely similar results, except for the Stordal samples, of which one (264165) is a high-Si variety. These data are presented in Figs 10 and 11 as rare earth element (REE) and multi-element diagrams with comparisons.
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Two element enrichment patterns can be clearly distinguished from the multi-element diagrams. One pattern is common to the rocks of Types 0, 1a, 1b and 3. It shows extreme enrichment in Ba, U, Nb–Ta and light to middle REE (LREE to MREE), deep troughs for Rb, Th, K, Pb and Ti, and lesser troughs for Sr, P and Zr–Hf. In comparison, melilitites, nephelinites and meimechites show similar levels of enrichment for many elements but have much smoother spectra, with moderate K troughs being the most distinctive (Rogers et al., 1992; Arndt et al., 1995, 1998; Hegner et al., 1995; Wilson et al., 1995; Bernstein et al., 2000; Späth et al., 2001). Enriched picrites from Nuanetsi have incompatible element concentrations up to 200 times primitive mantle; however, their spectra show no K anomalies and have distinct Nb–Ta troughs (Ellam & Cox, 1989, 1991). Carbonatites and allegedly carbonatite-metasomatized mantle xenoliths show variable degrees of enrichment and have spiky patterns with similarities to those of the Manîtdlat Member, showing deep troughs for Rb, K, Zr and Ti, and strong relative enrichment in Ba, Nb and LREE (Nelson et al., 1988; O’Reilly & Griffin, 1988; Yaxley et al., 1991; Larsen & Rex, 1992; Le Roex & Lanyon, 1998; Coltorti et al., 2000). The Manîtdlat Member rocks have low to very low Th/U (down to 1·2 in Type 3), whereas most carbonatites have high Th/U although the ratios are very variable.
A different element enrichment pattern is seen in the Type 2 rocks. The incompatible-element enrichment is less extreme, Rb is enriched, Ba much less so, the troughs for Th, K, Pb and Sr are very small, and there is a large Nb–Ta peak and lesser peaks for Zr–Hf and Ti. This pattern is similar to those of some basanites and alkali basalts from eastern Australia (O’Reilly & Zhang, 1995; Zhang et al., 2001), and remarkably similar to patterns of amphibole from some mantle xenoliths (Moine et al., 2001).
The Stordal type has an enrichment pattern most similar to those of Types 0, 1 and 3, but with smaller troughs for Rb, Th and K. The high-Si Stordal sample has higher Rb and Pb, and lower U, Nb, Ta and REE than the other Stordal sample.
The dyke is the least enriched of the alkaline rocks. Its incompatible element contents all lie between those of the tholeiites and the other alkaline rocks, and the trace-element pattern, with low Rb and high Nb/La, mostly resembles that of Type 1a.
The uniqueness of the Manîtdlat Member magmas may be illustrated by comparison with incompatible element ratios in other strongly enriched rocks (Fig. 12). In melilitites, nephelinites and meimechites, K/Ba, Rb/Sr, Ba/Nb and La/Nb ratios vary over about one order of magnitude. In the Manîtdlat Member rocks these ratios vary over about two orders of magnitude, and whereas Type 1a and Stordal often plot together with other rocks, Type 2 lies at one extreme end with higher K/Ba and Rb/Sr and lower La/Nb and Ba/Nb than most other rocks, and Types 0, 1b, and 3 lie at the opposite extreme with lower K/Ba and Rb/Sr and higher La/Nb and Ba/Nb than almost all other rocks. The significance of the spread in Fig. 12 and the good correlation within the Manîtdlat Member is discussed below.
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Isotopes
The alkaline rocks of the Manîtdlat Member have very unusual isotope compositions (Table 5). The Sr–Nd–Pb results for the two samples analysed of each type are mutually consistent, and the results for the tholeiitic picrites of the Ordlingassoq Member, analysed simultaneously with the alkaline rocks, are in complete agreement with earlier results by Holm et al. (1993), Lightfoot et al. (1997) and Graham et al. (1998).
The tholeiitic rocks have positive Nd and negative Sr (Fig. 13) and plot within the field for the Iceland mantle plume (e.g. Stecher et al., 1999). They have been interpreted by earlier workers (Holm et al., 1993; Lightfoot et al., 1997; Graham et al., 1998) as produced by melting of the asthenospheric mantle in the proto-Icelandic mantle plume, and the data fields for the Ordlingassoq Member shown in Figs 13 and 14 thus conceivably represent the local contemporaneous asthenospheric mantle.
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The alkaline rocks (excepting the dyke) all have negative Nd, and all except Type 2 and the high-Si Stordal lava also have negative Sr (Fig. 13). Types 0, 1b and 3 have nearly identical Nd–Sr isotope compositions. Except for Type 2, all the alkaline rocks plot well below the oceanic mantle array in the Nd–Sr isotope diagram, in an area of the lower left quadrant occupied by very few uncontaminated mantle-derived rocks. None are known from the North Atlantic Igneous Province; the few we have noted are some potassic rocks from the Elkhead Mountains, Colorado (Leat et al., 1988; Thompson et al., 1989), some carbonatites and kimberlites from the Archangelsk region, NW Russia (Mahotkin et al., 2000), and some nephelinites from the Napak and Mt Elgon volcanoes in East Africa (Simonetti & Bell, 1994, 1995).
The Pb isotope compositions (Fig. 14) of the tholeiites of the Ordlingassoq Member cluster around the Northern Hemisphere Reference Line (NHRL) and fall within the Iceland field. In contrast, the alkaline rocks have high 206Pb/204Pb ratios and plot below the NHRL. The basement rocks in the area have low Pb isotope ratios (Fig. 14), and the high-Si Stordal sample is clearly displaced towards basement values. In terms of the vertical distance from NHRL as defined by Hart (1984), the Manîtdlat Member rocks have 
7/4 = -8 to -25, 8/4 = 0 for Type 2 and 8/4 = -100 to -257 for the other types, outside the range of all modern mid-ocean basalts (MORB) and ocean island basalts (OIB) (Thirlwall, 1997). We do not know of any other mantle-derived uncontaminated igneous rocks with Pb isotope ratios similar to those of the Manîtdlat Member. Many alkaline rocks have similarly high 206Pb/204Pb ratios, but for a given 206Pb/204Pb they all have higher 207Pb/204Pb and 208Pb/204Pb ratios, most of which lie close to or above the NHRL (e.g. Nelson et al., 1988; Simonetti & Bell, 1994, 1995; Hegner et al., 1995; Wilson et al., 1995; Kalt et al., 1997; Le Roex & Lanyon, 1998; Bell & Tilton, 2001; Späth et al., 2001, and compilation therein). Pb isotope ratios similar to those of the Manîtdlat Member are, however, reported from some highly metasomatized peridotite xenoliths with metasomatic oxides from South Africa (Hawkesworth et al., 1990) and in metasomatic oxide (mathiasite) from such xenoliths (Fig. 14; Griffin et al., 1999).
All the samples analysed for Os isotopes have low to very low 187Os/188Os ratios (0·1342–0·1067, Table 6) and all except the Type 1a sample have negative Os (Fig. 15). There is an inverse correlation between the Os isotope ratios and the amounts of Os present in the samples (Table 6). One sample of Type 0 has exceptionally high Os, 44 ppb. These data are as unusual as the other isotope data: 187Os/188Os ratios below 0·110 have previously not been reported from igneous rocks but only from peridotite xenoliths from old subcontinental lithospheric mantle (Pearson et al., 1995; Hanghøj et al., 2001).
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESPETROGRAPHY AND MINERALOGYWHOLE-ROCK GEOCHEMISTRYDISCUSSIONCONCLUSIONSAPPENDIX: CALCULATION OF MELTING...SUPPLEMENTARY DATAREFERENCES |
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Modification of the primary alkaline magmas
Identification of crustal contamination
The effects of crustal contamination on the tholeiitic volcanic rocks on Disko and Nuussuaq are well described (Pedersen, 1979, 1985a; Pedersen & Pedersen, 1987; Goodrich & Patchett, 1991; Pedersen et al., 1996; Lightfoot et al., 1997). Contamination with either Precambrian basement or Mesozoic sediments leads to increase in SiO2, K2O, Rb and Pb, and decrease in FeO*, CaO, TiO2 and Nb (but not appreciably Zr). Sr isotope ratios increase, and Nd and Pb isotope ratios decrease. Compositional shifts of this kind are seen in the three high-Si Stordal samples, one of which is analysed isotopically, and to a lesser extent in four Type 1a samples (Figs 6, 7, 9, 13 and 14). We therefore consider these samples to represent crustally contaminated alkaline magmas. Modelling of the crustal contamination process is outside the scope of the present paper.
On the other hand, the character of the alkaline rocks as such, with their strong enrichment in Ba, U, Nb, LREE and P, and low Sr isotope ratios and high Pb isotope ratios, cannot be explained in terms of contamination with any known crustal components in the region.
Evidence for mixing of tholeiitic magma into alkaline magma
The alkaline magmas were erupted within a regional zone of eruption centres that mainly produced tholeiitic magmas, and it is likely that the magmas utilized the same conduit systems. The conduits would be possible sites of mixing between alkaline and tholeiitic magma batches and their phenocrysts, of which the tholeiites contained only olivine + chromite. The heterogeneous olivine populations within single samples suggest that such mixing has taken place (Fig. 3). Low-CaO zoned olivines such as phenocryst 1 in Fig. 3 are explicable as crystals that originally formed in tholeiitic magmas and later were mixed into the alkaline magmas where they partly re-equilibrated and continued their growth. The ‘tholeiitic’ levels of Ca, Cr and Ni in the cores are preserved, whereas Mg and Fe are re-equilibrated to lower mg-numbers. The low CaO/Al2O3 in the glass inclusions in the low-CaO olivines (Fig. 5) confirms the tholeiitic character of their parental melts. These low-Ca olivines have mg-numbers of 82–92·3 (Fig. 3). Phenocrysts 2–4 in Fig. 3 are normal phenocrysts in the alkaline magma.
Because of the evidence for mixing it is necessary to evaluate the extent to which this process has modified the composition of the original alkaline magmas. The main constraint on this comes from the near-constant Nd isotope ratios in Types 0, 1b and 3; the total variation in 143Nd/144Ndi for these rocks is only 0·00002 (0·51236–0·51238, Table 5). Because the Nd contents are strongly dominated by the alkaline component, moderate amounts of tholeiitic magma can be mixed into the alkaline magma before the Nd isotope ratio changes significantly. A Type 0 alkaline end-member magma may be mixed with up to 25% tholeiitic magma before the Nd isotope ratio increases by more than 0·00002 (mixing curve in Fig. 13), and 25% is therefore considered a maximum amount of in-mixed tholeiitic magma. We have no means of quantitatively constraining the amounts of tholeiitic magma further, but based on the relative scarcity of the tholeiitic olivine crystals we consider that the amount of tholeiite in the alkaline magmas was normally 10% or less. It is possible that the tholeiitic component is dominated by olivine crystals picked up in the mush zones in the conduit systems, with very little accompanying tholeiitic melt.
With an upper limit of 25% tholeiitic magma in the alkaline magmas, the major-element composition of the unknown pure alkaline end-member melt will not be very different from that of the erupted magmas. SiO2 in the end-member will be lower than in the erupted rocks by 0·5 wt % or less; CaO/Al2O3 will be higher but still <1·26. TiO2 and Na2O are invariably low. The incompatible trace elements in the alkaline end-member will be higher by a factor of 1·3 or less; ratios of more incompatible elements (e.g. Nb/La, Th/U) will be virtually unaffected, whereas ratios of more incompatible to less incompatible trace elements (e.g. Nb/Y, Ba/Ti) will be lowered by the tholeiitic component. With an upper limit of 10% in-mixed tholeiite the alkaline magmas are practically unchanged. In conclusion, there is undoubtedly a small tholeiitic component in the alkaline magmas, but the compositional influence of this is negligible for most elements.
The dyke: mixing of alkaline magma into tholeiitic magma
Whereas the composition of an alkaline magma is fairly robust against addition of minor amounts of tholeiitic magma, the opposite is not the case. This is illustrated by the investigated dyke that cuts the entire Manîtdlat Member succession. The dyke can be interpreted in terms of mixing of alkaline magma into tholeiitic. Its major-element composition is tholeiitic (Fig. 6), whereas its incompatible trace elements (including K and P) are intermediate between tholeiitic and alkaline values, higher than those in the tholeiites by a factor of 2–4 (Fig. 11). The Nd–Sr isotope composition of the dyke is also intermediate between that of tholeiitic and alkaline rocks (Fig. 13). The low Rb and normalized Ta/La>1 suggest a relation to Type 1a, and the dyke is actually situated in an area where the alkaline rocks are solely of Type 1a. Simple mixing calculations between tholeiite (sample 326783) and Type 1a (sample 326787) give consistent results for most trace elements and the Nd isotopes, suggesting that the dyke is a tholeiitic magma that contains around 15% alkaline component.
Primary alkaline magmas
The composition of the most magnesian, possibly primary, alkaline magmas may be estimated from the most magnesian cognate olivines present. The most magnesian high-CaO olivine has mg-number 90 (Fig. 5c), which corresponds to a melt calculated to have around 15 wt % MgO, somewhat dependent on the oxidation state. Thus, samples with >15 wt % MgO most probably contain accumulated olivines whereas samples with lower MgO may represent the erupted and more or less fractionated magmas. At 15 wt % MgO the tholeiitic melts had temperatures close to 1400°C (Larsen & Pedersen, 2000). The parental alkaline melts would have had lower temperatures, loosely estimated around 1300°C.
Mantle sources and melting processes
The rocks of the Manîtdlat Member have incompatible trace element concentrations that are enriched by up to 100–300 times primitive mantle for elements such as Ba, U, Nb and La (Fig. 11). Similar enrichment levels in basic igneous rocks are normally found in melilitites, nephelinites and meimechites, and also some kimberlites. These rock types are all strongly silica undersaturated, sometimes larnite normative, plagioclase free or plagioclase poor, and are considered to be formed by small degrees of melting of enriched, volatile-bearing mantle (e.g. Nelson et al., 1988; Wilson, 1989; Rogers et al., 1992; Taylor et al., 1994; Arndt et al., 1995, 1998; Hegner et al., 1995; Wilson et al., 1995; Mahotkin et al., 2000). In comparison with these rocks, those of the Manîtdlat Member are not highly undersaturated and have relatively high contents of SiO2 and Al2O3, and low TiO2, alkalis and P2O5 (Fig. 16). The effusive eruption style and the anhydrous mineralogy do not suggest that the volatile contents of the alkaline Manîtdlat Member magmas were higher than in the tholeiites.
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Manîtdlat Member melts such as those of Types 0 and 1 have CaO/Al2O3 around 1·2 and, at the 15 wt % MgO level, 11·5 wt % Al2O3 and 1·0 wt % Na2O. According to the melting model of Herzberg & Zhang (1996), such melts cannot be produced near the solidus of an ordinary dry lherzolite at any pressure. Olivine fractionation from a deep near-solidus melt may produce enriched alkali picrites (Milholland & Presnall, 1998), but they will not be like the alkali-poor Manîtdlat Member melts. Dilution of an extreme small-degree alkaline melt with tholeiitic melt is another possibility, but this requires large amounts of tholeiitic melt (80%), which, as discussed above, is very unlikely. A more straightforward explanation is that the melting involved unusual mineral assemblages in enriched and metasomatized parts of the mantle.
Plume mantle or subcontinental lithospheric mantle?
In continental areas, the enriched mantle sources for basic alkaline rocks have been envisaged to be situated in the subcontinental lithospheric mantle, although the Nd–Sr isotopes often indicate that the lithosphere was metasomatized by asthenospheric melts shortly before the melting (e.g. Wilson, 1989). In oceanic islands, enriched mantle components such as EM1 and EM2 are considered to reside within the mantle plume that brings them to the surface and into the melting regime (e.g. Hofmann, 1997). The idea of melting of enriched plume components has also been applied to some continental basic alkaline rocks, particularly the meimechites of Siberia (Arndt et al., 1995, 1998), and the melilitites and nephelinites of central East Greenland (Bernstein et al., 2000, 2001). Indications of the location of the enriched mantle source for the rocks of the Manîtdlat Member come from the isotope compositions and the field relations.
The isotope compositions of the Manîtdlat Member rocks of Types 0, 1, 3 and Stordal plot outside the fields of the known oceanic mantle components for all of the isotope systems investigated (Figs 13–15). The few other rocks that also plot within the lower left quadrant of the Nd–Sr isotope diagram are interpreted to be either melts from enriched subcontinental lithosphere, or asthenospheric melts that have reacted with such lithosphere (Thompson et al., 1989; Simonetti & Bell, 1995; Mahotkin et al., 2000). As noted above, the Pb and Os isotope data are comparable only with data from peridotite xenoliths representing the subcontinental lithospheric mantle. The melilitites and nephelinites from East Greenland that are suggested to be derived from an enriched component intrinsic to the Iceland mantle plume have Nd–Sr isotopes that plot on the mantle array in extension of the Iceland field, and Pb isotopes that plot within to slightly above the Iceland fields (Bernstein et al., 2001, and unpublished data, 2002). If the enriched mantle source for the Manîtdlat Member was an intrinsic part of the proto-Icelandic plume, then it is a new and unique component of that plume.
The geological setting of the Manîtdlat Member is significant in this context. The alkaline magmas were generated only during a brief period in the middle of a much longer period of formation of high-degree melts from the plume (the tholeiitic picrites), and they were erupted within only a 20 km by 30 km area. If the enriched source was a localized, unique component within the upwelling plume, being the most fusible part it would have started melting at deeper levels than the depleted surroundings. This requires that the alkaline melts ascended through the overlying melting zone producing the tholeiites (and through the lithosphere) while retaining their identity without being mixed into the contemporaneous tholeiitic melts. This is contrary to the concept of pooling of melts in the melting column.
On the other hand, if the enriched mantle source was a limited volume situated within the subcontinental lithospheric mantle beneath northern Disko, the particular setting of the Manîtdlat Member volcanic rocks is easily explicable. The regional tholeiitic feeder systems progressed with time from NW to SE, and only when they moved to the Stordal–Maniillat area (Fig. 1) did the conduits start to traverse the enriched parts of the lithosphere. The passing hot tholeiitic picrites provided the necessary heat for the melting of the enriched lithosphere material, and when the low-melting parts had been removed the production of alkaline magma ceased whereas the tholeiitic magma production from the plume continued.
Age of the enriched components
It is clear from the radiogenic isotope compositions of the Manîtdlat Member rocks that the ultimate derivation of the enriched source components from a depleted asthenospheric mantle cannot be related to the 60 Ma melting event but must be older. As a consequence of radiogenic ingrowth in the enriched source the isotope ratios of the Manîtdlat Member rocks at 60 Ma correlate broadly with the appropriate parent/daughter ratios; in particular, the Type 2 rocks, which have higher Sm/Nd and Rb/Sr and lower U/Pb than the other types, also have higher 143Nd/144Nd and 87Sr/86Sr and lower 206Pb/204Pb (Figs 11, 13 and 14). The age of the enrichment is, however, very uncertain. If the Type 2 source at some initial stage had Nd, Sr and Pb isotope ratios similar to the sources for the other magma types this could have been at around 300–500 Ma. Assuming that the Sm/Nd ratios were not significantly changed during the melting event at 60 Ma, the Nd isotopes give Nd extraction ages from a depleted reservoir of 1·1–1·4 Ga. The oldest age indication comes from the Os isotopes, where the Os in the two samples with 187Os/188Os = 0·106–0·108 must have evolved in a practically Re-free environment for nearly 3 Gyr (Table 5); this is, however, a depletion age and could represent melting of sulphides in the ancient side wall to the enriched areas.
Metasomatized mantle lithologies
Minerals known from metasomatized mantle xenoliths are (besides olivine) clinopyroxene, amphibole, phlogopite, apatite, carbonate, ilmenite, rutile and a number of exotic oxides. Orthopyroxene and spinel or garnet are often rare or absent because they have reacted to form clinopyroxene, amphibole or mica (O’Reilly & Griffin, 1988; Yaxley et al., 1991, 1998; Van Achterbergh et al., 2001). If the metasomatizing agent was carbonate rich, much of the CO2 released by the decarbonation reactions will subsequently have been lost (Yaxley et al., 1991, 1998).
Melting of clinopyroxene-rich lithologies has been invoked to explain the high CaO and CaO/Al2O3 of nephelinites and melilitites (e.g. Francis & Ludden, 1990; Francis, 1991; Gibson et al., 1999). Clinopyroxenes in metasomatized mantle xenoliths may contain up to 100–600 ppm Sr, 10–100 ppm La and 20–200 ppm Ce (Erlank et al., 1987; O’Reilly et al., 1991; Hauri et al., 1993; Ionov et al., 1997; Ionov, 1998; Zanetti et al., 1999; Laurora et al., 2001; Van Achterberg et al., 2001) and may thus yield a significant contribution to the Sr and REE budget of the melt; however, elements such as Ba and Nb do not exceed 1–2 ppm in clinopyroxene, which even at 1% melting will not yield the large degrees of enrichment seen in the Manîtdlat Member.
Amphibole in metasomatized mantle xenoliths is often pargasitic. Mantle pargasite contains 0·4–2·0 wt % K2O, typically around 1 wt % K2O (O’Reilly & Griffin, 1988; O’Reilly et al., 1991; Ionov & Hofmann, 1995; Chazot et al., 1996; Ionov et al., 1997; Ionov, 1998; Zanetti et al., 1999; Laurora et al., 2001; Moine et al., 2001). Therefore, despite the presence of deep K troughs in the trace element patterns of most Manîtdlat Member rocks (Fig. 11) their source may have contained significant proportions of pargasitic amphibole. Amphiboles are capable of accommodating higher amounts of incompatible trace elements than clinopyroxene, particularly Nb and Ta (see the references cited above), and multi-element spectra of mantle pargasites show distinct peaks at Nb–Ta (Ionov & Hofmann, 1995; Laurora et al., 2001; Moine et al., 2001), rather similar to the patterns of the Manîtdlat Member Type 2 melts (Fig. 11). A large fraction of pargasite in the melting assemblage will lead to high Al contents in the melts, as seen in the Manîtdlat Member.
Phlogopite contains about 8 wt % K2O and has high concentrations of Rb and Ba. The deep troughs for Rb and K in the trace element patterns of all the Manîtdlat Member melts except those of Type 2 suggest that no or only little phlogopite contributed to the melts. The huge amounts of Ba must be provided by other phases.
Mantle apatite may accommodate high amounts of Th, U, Sr and REE (O’Reilly et al., 1991; Chazot et al., 1996; Ionov et al., 1996, 1997; O’Reilly & Griffin, 2000), but apatite does not fractionate Th from U to the degree seen in the Type 0, 1 and 3 melts (Fig. 11). Moreover, Types 1a and 1b have similar concentrations of P2O5 but very different levels of enrichment in Sr and REE (Figs 6 and 11), suggesting that an additional phase delivered these elements to the Type 1b melts.
Metasomatic oxides such as the Ba–K titanates lindsleyite and mathiasite (LIMA phases) have very high concentrations of Ba, Sr, LREE, Nb, Zr, Pb, U and Th, very high U/Th ratios, and Pb isotope ratios similar to those of some Manîtdlat Member melts (Jones, 1989; review by Haggerty, 1991; Griffin et al., 1999). A LIMA phase in the melting assemblage would explain some of the unusual trace element ratios such as the high U/Th, and also the unusual Pb isotopes of the Manîtdlat Member Type 0, 1 and 3 melts.
The lower contents of Ni in the Manîtdlat Member rocks compared with tholeiites with similar MgO contents (Fig. 6) suggest that olivine and orthopyroxene constituted a smaller part of the melting assemblage than in an ordinary lherzolite, and this is consistent with a metasomatized source.
In conclusion, the most probable mantle sources for the Manîtdlat Member melts are amphibole–clinopyroxene-rich lithologies with apatite and, for Types 0, 1 and 3, a LIMA phase. Mantle xenoliths with such lithologies (but without LIMA) have been described from eastern Australia by Wass (1979), O’Reilly & Griffin (1988) and O’Reilly et al. (1991), and have been invoked as source for the nephelinites and basanites in that region (O’Reilly & Zhang, 1995), which show several similarities to the Manîtdlat Member rocks, namely, high Al2O3 and low TiO2, P2O5, Rb/Sr and La/Nb. In the Australian xenoliths, amphibole and apatite occur both disseminated through lherzolite and concentrated in centimetre-sized veins.
Melting conditions and residual phases
Pargasitic amphibole is stable to 1050°C and 30 kbar, 96 km (Gilbert et al., 1982). In South Africa, LIMA phases were formed together with richteritic amphibole in the garnet stability field at estimated depths of 75–100 km (Haggerty, 1991). At 60 Ma, just before break-up, the lithosphere in West Greenland was strongly attenuated and only around 100 km thick (Herzberg & O’Hara, 1998). The metasomatized mantle volumes beneath West Greenland were probably situated at deep levels in the attenuated lithosphere but precise depth estimates cannot be made.
It is virtually impossibe to detect variations in the degree of melting between types because all the usual chemical indicators, such as CaO/Al2O3 and Ce/Y, reflect the source enrichment more than the degree of melting. If phlogopite was a refractory phase the less enriched Type 2 melts could have resulted from higher degrees of melting than the other alkaline melts, but then the phases such as LIMA that contributed to the peculiarities of the other melts should also be discernible in the Type 2 melt, which they are not. In all, we consider that differences in the melting assemblages are far more important than variations in the degree of melting for production of the compositional spread observed.
Many basic alkaline rocks have deep K troughs in their trace element spectra, and this has been interpreted as a result of the presence of a residual potassic phase during melting, either phlogopite (Foley & Wheller, 1990; Rogers et al., 1992) or amphibole (Späth et al., 2001). This interpretation is necessitated by the assumption that the enriched mantle source for the melts had a smooth trace element spectrum with no K trough. However, this assumption may in many cases be unwarranted; indeed the trace-element spectra of enriched mantle xenoliths are usually far from smooth and very often have deep troughs at K and sometimes at Zr and Ti (Fig. 11; Menzies et al., 1987; O’Reilly & Griffin, 1988; Yaxley et al., 1991; Hauri et al., 1993; Ionov et al., 1997; Gorring & Kay, 2000). The incompatible trace elements reside mainly in the metasomatic phases, which are often volatile-bearing and easily fusible. During melting of metasomatized mantle the metasomatic phases will melt preferentially, transferring the incompatible trace elements and thereby the spiky pattern to the melt, and leaving the depleted residual mantle with a smoothed trace element pattern and, in many cases, no residual phlogopite or amphibole but only the breakdown products from incongruent melting of these minerals. We consider that the spiky trace element spectra of the Manîtdlat Member magmas were inherited in this way from similar spectra of the enriched source.
Efficient preferential melting of metasomatic phases was demonstrated experimentally by Foley et al. (1999) for mantle assemblages with amphibole, apatite, clinopyroxene, mica and ilmenite. At 15 kbar, the solidus for pargasite-bearing assemblages is in the range 1050–1075°C, and amphibole and oxide melt completely within a few tens of degrees above the solidus. Apatite survives to slightly higher temperatures, whereas mica melts over a larger temperature range up to 1170°C. In West Greenland, the tholeiitic melts originated in the asthenosphere with temperatures of 1550°C; within the lithosphere their temperatures dropped to 1400–1300°C for melts with 15–12 wt % MgO (Larsen & Pedersen, 2000). We envisage that these melts provided sufficient heat rapidly enough to raise the temperatures in the adjacent lithosphere to 1100–1170°C, thereby inducing complete melting of amphibole, oxide, apatite and mica. The restite then consisted of clinopyroxene + olivine ± orthopyroxene and would have a sharply raised melting temperature. Farther from the heat source the metasomatic minerals may not have melted completely, but it is conceivable that melts from such areas would make only a small contribution to the total melt volume produced.
Melting assemblages for generation of the Manîtdlat Member melts
Quantification of the melting process is difficult because of the number of unknowns involved. The source may be veined, and both veins and wall rocks would melt but to very different extents (Foley, 1992; Foley et al., 1999), making parameters such as the degree of melting complex if not meaningless. But if we take the existing melts as starting points, viewing them as the sums of the contributions from the various melting minerals, we can calculate the bulk melting mineral assemblages. When we make the simplifying assumption, justified above, of complete melting of the enriched phases except clinopyroxene, simple mass-balance equations can be used in the calculations of the melt compositions. The details of these calculations are described in the Appendix and given in the Electronic Appendix B; the modelled melting assemblages are shown in Table 7, and Fig. 17 shows the incompatible trace element inventory of the various Manîtdlat Member magma types distributed between the contributing melting phases.
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For the major elements, combinations of analysed mantle minerals were used in mixing equations to approximate the major-element composition of the primary alkaline melts with 15 wt % MgO. This worked surprisingly well and showed that melts calculated as consisting of roughly 60 wt % amphibole, 30 wt % clinopyroxene, 10 wt % olivine and 1 wt % apatite are fairly close to the major-element compositions of the Manîtdlat Member melts. In detail, the various melt types require slightly different melting assemblages. Type 2 requires small amounts of mica (4%), less clinopyroxene (27%), and minerals with compositions slightly different from those of the other types. The Stordal type can be modelled only with low clinopyroxene (26%), and orthopyroxene instead of olivine because of the relatively high SiO2. All the Stordal magmas may be crustally contaminated to some extent.
For the incompatible trace elements, the known inventory of these in the primary alkaline melts was distributed between the contributing phases as detailed in the Appendix. It must be stressed that these calculations do not provide unique solutions. They characterize possible and reasonable solutions and highlight the differences between the various melt types involved, but other solutions are not thereby invalidated.
As shown in Fig. 17, amphibole contains a significant part of the trace-element inventory (and K), particularly Nb, Ta, Zr and Hf. The clinopyroxene contribution is small and that of olivine is, by our definition, nil.
A LIMA phase is required in the melting assemblages for Types 0, 1a, 1b and 3, and it must be one or both of the Ba end-members lindsleyite, Ba(Ti,Cr,Fe,Zr,Mg)21O38, or hawthorneite, Ba(Ti,Cr,Fe,Mg)12O19 (LIHA). These minerals have very high concentrations of Ba, U, Nb, Ta, REE and Sr (e.g. Jones, 1989; Haggerty, 1991; Griffin et al., 1999). They also have high TiO2 (20–60 wt %), but because only a small fraction of the mineral is required the contribution to the total Ti budget is small compared with those from amphibole and clinopyroxene. The required LIHA phase cannot be Zr rich, which could suggest it is hawthorneite. The LIHA phase is the main phase responsible for the extreme Th–U fractionation and the high Pb isotope ratios in the Type 0, 1 and 3 rocks.
In the Type 2 melting assemblage, mica is required by both the major and trace element constraints and contributes significant amounts of Rb, Ba and K to the melt. The amphibole is almost solely responsible for the large Nb–Ta peak (Fig. 17). Apatite is the main phase responsible for the contents of Th and U (and P), whereas amphibole and apatite contribute equally to the REE.
In the Stordal melting assemblage, Th and U come mainly from apatite whereas the REE and Sr are shared between apatite and amphibole, with a smaller contribution from clinopyroxene. The apatite has higher Th, U and REE than the apatite in Type 2.
For the Type 1a and 1b melts, the modelled melting assemblages are in the first round closely similar. However, compared with Type 1a, Type 1b is significantly more enriched in Ba, Sr, Pb and LREE–MREE, but not in U, Nb, Ta and Ti (Fig. 11). Thus, either the oxide phase has another composition or an additional phase is involved, which concentrates Ba, Sr, Pb and LREE–MREE. The most Ba rich of the metasomatic oxides is hawthorneite, with c. 13 wt % BaO, which occurs in association with lindsleyite (Erlank et al., 1987; Haggerty et al., 1989). It is possible that the oxide in Type 1a is lindsleyite and in Type 1b is hawthorneite, but this makes the similar U–Nb–Ta and Ti in the two types fortuitous and is difficult to model satisfactorily. Carbonates with high contents of Ba, Sr and LREE–MREE have been reported from carbonatites (Knudsen, 1991; Hornig-Kjarsgaard, 1998), in metasomatized mantle xenoliths (Jones et al., 2000), in melt inclusions in lherzolites (Van Achterbergh et al., 1999), and in doped experiments (Veksler et al., 1998). We have tentatively assigned 4 wt % carbonate with high carbonate–amphibole partition coefficients for Ba, Sr, Pb and REE to the Type 1b melt. A third possibility is a yet unknown phase, in which case much smaller amounts may be required.
The Type 0 and Type 3 melting assemblages are closely related to that of Type 1b but require slightly different proportions of the minor phases (Table 7).
Structure of the metasomatized mantle
As discussed above, a number of metasomatized source areas with different mineralogies must have been present in the lithospheric mantle to give rise to the various magma types of the Manîtdlat Member. The source areas were present within the same limited part of the lithosphere beneath northern Disko, and only there, and they melted simultaneously. Although separate they must have been closely associated, and most probably they were genetically related. The existence of the magma types suggests an ordered structure of the metasomatized mantle domains rather than a random medley of veins and patches.
O’Reilly & Griffin (1988) and O’Reilly et al. (1991) suggested that metasomatism in the lithospheric mantle is caused by volatile components released from crystallizing veins of alkaline basic magma (ultimately asthenosphere derived). The released volatile components will be rich in incompatible elements, which migrate through the mantle sidewall and precipitate in a sequence depending on the mineral stabilities and the mineral–fluid partition coefficients, as in a chromatographic column. In particular, mica will form only close to the veins, trapping Rb, K and Ti, and thereby depleting these elements in the migrating fluid, which becomes more carbonate rich with distance from the vein. Elements such as Ba, Th, U and LREE will move farther before they are precipitated, e.g. in apatite. The result is a number of metasomatic zones.
If the metasomatized mantle source for the Manîtdlat Member magmas had a similar zoned structure, both the occurrence of the different magma types and the complementary relationship between Type 2 and Type 1b seen in Fig. 12 would be explicable. Figure 18 shows the suggested model, which is slightly modified from that of O’Reilly et al. (1991) with regard to the successions of amphibole and apatite. We also assume precipitation of an oxide phase, and an additional stage with oxide + carbonate or an unknown phase. The Type 2 melts would then be derived from vein-near micaceous mantle, the Stordal and Type 1a melts from intermediate zones, and Type 0, 1b and 3 melts would be derived from the zones farthest away. The magma conduits must cross all the zones, and the zones must extend vertically to avoid mixing of the melts. Type 1a is the volumetrically dominant melt type; this suggests that the corresponding metasomatic zone is the widest of the successive zones.
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The volume of the Manîtdlat Member alkaline melts is at least 30 km3, which requires that the metasomatic zones are much wider than the few decimetres envisaged by O’Reilly et al. (1991), and that the original intrusion was on a larger scale than just ‘veins’. In a rough calculation, 30 km3 of primary melt produced by on average 10% melting corresponds to 300 km3 metasomatized mantle. The length of the feeder systems for the Manîtdlat Member, considered to be NE-directed fissures, is c. 30 km (Fig. 1). Thus, if the metasomatized areas were 30 km in length and had a depth extension of, for example, 10 km, then the total width of the metasomatic zones must be around 1 km to encompass 300 km3 of material. This allows room for the contemporaneous tholeiitic melts that were erupted within the area covered by the Manîtdlat Member lavas. It does, however, involve the corollary that a substantial amount of crystallized mafic melt must be situated at depth in the lithosphere; in eastern Australia xenolith data support such a situation (O’Reilly et al., 1988).
Metasomatized mantle in West Greenland
The Precambrian basement in the whole of West Greenland is intruded by several occurrences of small-volume ultramafic–alkaline magmas such as lamprophyres, kimberlites and carbonatites, with ages ranging from Archaean to Eocene (Larsen & Rex, 1992). Thus, there is abundant evidence that the lithosphere contains metasomatized domains, but the extent of these is not known because their possible surface expressions, the alkaline magmas, are strongly controlled by the tectonic state of the lithosphere. The Manîtdlat Member itself is an example that metasomatized areas of the lithosphere may lie ‘dormant’ with no magmatic surface expression for hundreds of millions of years until a new tectonic regime allows melt generation and ascent to the surface.
The basement within the Nuussuaq Basin and in the stable areas east of it consists of Archaean crust of 2800 Ma age that was reworked during the Proterozoic around 1750 Ma (Kalsbeek, 1999). The region hosts two occurrences of alkaline rocks in addition to the Manîtdlat Member. The basement at Eqi (Fig. 1) contains a small swarm of weakly deformed, Palaeoproterozoic (
1780 Ma) carbonate-rich, ultramafic lamprophyre dykes (Larsen & Rex, 1992). On Ubekendt Ejland (Fig. 1), a small swarm of lamprophyre dykes dated at 34 Ma cut the 20 Myr older volcanic succession (Larsen, 1981, 1982; Clarke et al., 1983). The dykes are strongly enriched in incompatible trace elements, alkalis (up to 6·3 wt %, Fig. 16) and volatiles (5–9 wt % H2O + CO2), and they may be small-degree melts of metasomatized mantle with similarities to that of the source for the Manîtdlat Member. The age of the enrichment event is unknown. These three occurrences of alkaline rocks within 150 km of each other could perhaps be viewed as being derived from one extensive metasomatized zone in the Archaean lithosphere, mobilized at different times and by different tectonic events. However, as discussed above, the particular setting of the Manîtdlat Member clearly indicates that its enriched mantle source is an isolated domain. Thus, at least in the Disko–Svartenhuk region, the metasomatized mantle domains are individual separate entities 1000 km2 or less in horizontal extent, which lends support to the theory that the enrichment events were associated with discrete intrusions of asthenospheric magmas.
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESPETROGRAPHY AND MINERALOGYWHOLE-ROCK GEOCHEMISTRYDISCUSSIONCONCLUSIONSAPPENDIX: CALCULATION OF MELTING...SUPPLEMENTARY DATAREFERENCES |
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The enriched alkaline picrites and basalts of the Manîtdlat Member represent a discrete, short-lived melting event during the formation of the tholeiitic picrite melts of the Vaigat Formation. The ascending alkaline magmas utilized the tholeiitic conduit systems, as shown by the presence of occasional tholeiitic (low-Ca) olivine xenocrysts; however, the compositional effect of in-mixing of small amounts of tholeiitic magma is very small.
The alkaline rocks have anhydrous primary mineral assemblages of mainly olivine, chromite, clinopyroxene and plagioclase; this may be explicable by high temperatures, with matrix glasses quenched at 1150–1190°C, above the stability limit of amphibole.
Six alkaline magma types can be defined (Stordal and Types 0, 1a, 1b, 2 and 3), all showing variable enrichment in Ba, U, Nb–Ta and LREE. Five types show relative depletion in Rb, Th, K, Zr–Hf and Ti, whereas the least enriched Type 2 shows no depletion in these elements and particular enrichment in Nb–Ta and Zr–Hf. Unusual Sr–Nd–Pb–Os isotope compositions, with low 87Sr/86Sr, 143Nd/144Nd and 187Os/188Os, and high Pb isotope ratios, are outside the range of any modern asthenospheric mantle components and have resulted from melting of source rocks with long-term low Rb/Sr, Sm/Nd and Re/Os, and high U/Pb and U/Th.
The most likely origin for the alkaline melts is in old metasomatized lithospheric mantle domains rich in amphibole, clinopyroxene and apatite, some parts also containing mica and other parts containing small amounts of metasomatic oxides (lindsleyite or hawthorneite, LIHA), and possibly carbonate. The amounts of the mantle phases going into the melt can be assessed by simple mass-balance calculations, which show that the bulk of the melts is made up of 60% pargasitic amphibole, 26–30% clinopyroxene, 9% olivine and 1% apatite. The melt types most enriched in Ba, U, Nb–Ta and LREE in addition require 0·3–0·7 % LIHA phase, and some types (0, 1b and 3) require an additional phase (carbonate?) with these elements. The least enriched Type 2 melt requires 4% mica in the melting assemblage.
The enriched mantle domains must be highly structured to give rise to the various melt types. Good correlations between some incompatible element ratios suggest a common link between the melt types. A mantle with metasomatic zones produced during an old event of migration of incompatible elements away from a crystallizing alkaline intrusion and precipitation of these elements in successive zones, as in a chromatographic column, can explain the partly complementary relation between the melt types. The zone closest to the contact selectively trapped Rb, K, Zr–Hf and Ti and later gave rise to the melts of Type 2, whereas the most distal zones gave rise to the most Ba–U–Ta–LREE-enriched melts.
In the Paleocene, large volumes of hot, asthenosphere-derived tholeiitic magmas traversed the thinned and fractured lithosphere. When the NW–SE migrating active conduits traversed the enriched domains in the lithosphere, extensive melting of the incompatible-element-rich low-melting phases took place. The volume of enriched mantle was small and the low-melting component was rapidly extracted, making the alkaline event very short-lived.
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APPENDIX: CALCULATION OF MELTING ASSEMBLAGES IN THE GENERATION OF THE MANÎTDLAT MEMBER MAGMAS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESPETROGRAPHY AND MINERALOGYWHOLE-ROCK GEOCHEMISTRYDISCUSSIONCONCLUSIONSAPPENDIX: CALCULATION OF MELTING...SUPPLEMENTARY DATAREFERENCES |
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The primary alkaline melts are considered to have 15 wt % MgO, and the bulk-rock analyses are recalculated to this value by addition or subtraction of olivine; elements incompatible in olivine are simply diluted or concentrated by the recalculation. For the Type 3 (ankaramitic) melts we used a sample with 10 wt % MgO and minimal evidence for clinopyroxene fractionation. The Stordal sample has, however, fractionated clinopyroxene.
The modelling assumes total consumption of all the enriched mantle phases except clinopyroxene during melting; then the calculations can be made using simple mass-balance equations. The calculations are available as spreadsheets in Electronic Appendix B, which may be downloaded from the Journal of Petrology web site at http://www.petrology.oupjournals.org.
For the major elements, combinations of analysed mantle minerals were used in mixing equations to approach the major-element composition of the primary alkaline melts. FeO and MgO were added up to make allowance for Fe/Mg fractionation during melting.
For the incompatible trace elements, the known inventory of these in the primary alkaline melt was distributed between the contributing phases with a simple mass-balance equation:
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The fractions of amphibole, clinopyroxene and olivine contributing to the melt are taken from the major-element calculations.
The fraction of apatite is accurately determined by the P2O5 content in the melt and the assumption that apatite contains 40 wt % P2O5.
The fraction of mica is determined by the K2O and Rb contents in the melt and assumptions of 0·7–1 wt % K2O in amphibole, 8 wt % K2O in mica, and Rb contents that must be realistic. Only the Type 2 melt has enough K and Rb to include a mica component.
The fraction of Ba–Ti-oxide (LIHA) is determined by the concentrations of Ba and Nb in the amphibole. Maximum concentrations were put at 800 ppm Ba and 100 ppm Nb because very few measured mantle amphiboles contain more than this and most are well below that level (O’Reilly et al., 1991; Ionov & Hofmann, 1995; Chazot et al., 1996; Ionov et al., 1997; Ionov, 1998; Zanetti et al., 1999; Moine et al., 2001), although some higher values have been reported (Ionov & Hofmann, 1995; Raffone et al., 2001). When the concentrations of Ba and Nb in the amphibole exceeded the maximum, an LIHA phase was introduced into the melting assemblage.
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SUPPLEMENTARY DATA |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGANALYTICAL TECHNIQUESPETROGRAPHY AND MINERALOGYWHOLE-ROCK GEOCHEMISTRYDISCUSSIONCONCLUSIONSAPPENDIX: CALCULATION OF MELTING...SUPPLEMENTARY DATAREFERENCES |
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Supplementary data are available on Journal of Petrology online.
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ACKNOWLEDGEMENTS |
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We are grateful to J. Rønsbo for assistance with the set-up of the microprobe, and particularly for the high-precision analyses. S. Bernstein and W. L. Griffin put unpublished data at our disposal. T. Andersen, S. Bernstein, W. L. Griffin, H. Hansen, P. M. Holm and D. Peate provided constructive comments and discussions, as did M. Wilson and the reviewers N. Arndt, A. C. Kerr and S. Y. O’Reilly. The Geological Survey of Greenland and the Arctic Station in Godhavn provided extensive support during fieldwork. The Danish Natural Science Research Council provided the electron microprobe and the X-ray fluorescence spectrometer at the University of Copenhagen. This paper materialized as a result of a grant from the Carlsberg Foundation to L.M.L. and is published with the permission of the Geological Survey of Denmark and Greenland.
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FOOTNOTES |
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*Corresponding author. Telephone: +45 38142252. E-mail: lml{at}geus.dk
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