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Journal of Petrology | Volume 44 | Number 7 | Pages 1247-1280 | 2003
© Oxford University Press 2003
Quantification of Magmatic and Hydrothermal Processes in a Peralkaline Syenite–Alkali Granite Complex Based on Textures, Phase Equilibria, and Stable and Radiogenic Isotopes
1 INSTITUT FÜR GEOWISSENSCHAFTEN, AB MINERALOGIE UND GEODYNAMIK, EBERHARD-KARLS-UNIVERSITÄT, WILHELMSTRASSE 56, D-72074 TÜBINGEN, GERMANY
2 INSTITUT DE MINÉRALOGIE ET GÉOCHIMIE, UNIVERSITÉ DE LAUSANNE, UNIL–BFSH2, CH-1015 LAUSANNE, SWITZERLAND
* Corresponding author. Telephone: +49 (0)7071 2972930. E-mail: markl{at}uni-tuebingen.de
RECEIVED AUGUST 28, 2002; ACCEPTED FEBRUARY 19, 2003
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONThe Puklen complex of the Mid-Proterozoic Gardar Province, South Greenland, consists of various silica-saturated to quartz-bearing syenites, which are intruded by a peralkaline granite. The primary mafic minerals in the syenites are augite ± olivine + Fe–Ti oxide + amphibole. Ternary feldspar thermometry and phase equilibria among mafic silicates yield T = 950–750°C, aSiO2 = 0·7–1 and an fO2 of 1–3 log units below the fayalite–magnetite–quartz (FMQ) buffer at 1 kbar. In the granites, the primary mafic minerals are ilmenite and Li-bearing arfvedsonite, which crystallized at temperatures below 750°C and at fO2 values around the FMQ buffer. In both rock types, a secondary post-magmatic assemblage overprints the primary magmatic phases. In syenites, primary Ca-bearing minerals are replaced by Na-rich minerals such as aegirine–augite and albite, resulting in the release of Ca. Accordingly, secondary minerals include ferro-actinolite, (calcite–siderite)ss, titanite and andradite in equilibrium with the Na-rich minerals. Phase equilibria indicate that formation of these minerals took place over a long temperature interval from near-magmatic temperatures down to
300°C. In the course of this cooling, oxygen fugacity rose in most samples. For example, late-stage aegirine in granites formed at the expense of arfvedsonite at temperatures below 300°C and at an oxygen fugacity above the haematite–magnetite (HM) buffer. The calculated 
18Omelt value for the syenites (+5·9 to +6·3
) implies a mantle origin, whereas the inferred 18Omelt value of <+5·1 for the granitic melts is significantly lower. Thus, the granites require an additional low-18O contaminant, which was not involved in the genesis of the syenites. Rb/Sr data for minerals of both rock types indicate open-system behaviour for Rb and Sr during post-magmatic metasomatism. Neodymium isotope compositions (
Nd1170 Ma = -3·8 to -6·4) of primary minerals in syenites are highly variable, and suggest that assimilation of crustal rocks occurred to variable extents. Homogeneous Nd values of -5·9 and -6·0 for magmatic amphibole in the granites lie within the range of the syenites. Because of the very similar neodymium isotopic compositions of magmatic and late- to post-magmatic minerals from the same syenite samples a principally closed-system behaviour during cooling is implied. In contrast, for the granites an externally derived fluid phase is required to explain the extremely low Nd values of about -10 and low 18O between +2·0 and +0·5 for late-stage aegirine, indicating an open system in the late-stage history. In this study we show that the combination of phase equilibria constraints with stable and radiogenic isotope data on mineral separates can provide much better constraints on magma evolution during emplacement and crystallization than conventional whole-rock studies. KEY WORDS: peralkaline; phase equilibria; assimilation; hydrothermal; Li-amphiboles; Greenland; Gardar
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INTRODUCTION |
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Felsic alkaline igneous rocks can be divided into two principal groups: (1) quartz- and feldspar-bearing, silica-oversaturated rocks; (2) feldspar- and feldspathoid-bearing, silica-undersaturated rocks. The genesis and origin of these two groups is believed to be different. Silica-undersaturated, alkaline, intrusive complexes commonly have isotopic compositions that reflect a magma source in the mantle (e.g. Perry et al., 1987; Kramm & Kogarko, 1994; Dunworth & Bell, 2001). In many cases, contamination or assimilation processes seem to be of minor importance and, consequently, undersaturated alkaline rocks are often interpreted as differentiated residues of benmoreitic or nephelinitic magmas formed in the upper mantle (Larsen & Sørensen, 1987; Kramm & Kogarko, 1994; Stevenson et al., 1997; Frisch & Abdel-Rahman, 1999) under relatively dry conditions and low oxygen fugacities (Harris, 1983; Caroff et al., 1993).
In contrast, based on the close spatial association of silica-undersaturated and silica-oversaturated rocks in many alkaline igneous provinces worldwide, the origin of silica-oversaturated alkaline to peralkaline rocks is often explained by crustal contamination of mantle-derived undersaturated magmas (e.g. Davies & Macdonald, 1987; Foland et al., 1993; Harris, 1995; Mingram et al., 2000; Schmitt et al., 2000; Späth et al., 2001). Studies on alkaline rocks of the midcontinental rift system in North America demonstrated that not only upper crust, but also granulite-facies rocks of the lower crust might interact with alkaline magmas (Heaman & Machado, 1992).
Various experimental studies (e.g. Piotrowski & Edgar, 1970; Sood & Edgar, 1970; Edgar & Parker, 1974; Kogarko & Romanchev, 1977, 1982; Scaillet & Macdonald, 2001) have shown that silica-undersaturated and -oversaturated peralkaline melts have a crystallization interval down to temperatures of 400°C; the exsolution of a fluid phase from a residual melt is believed to take place in the very late stages of magmatic evolution. There is general consensus that volatiles play a major role in the evolution of alkaline to peralkaline magmas, for both their chemical and physical evolution. Effects of late-stage fluids in some alkaline to peralkaline intrusions of the Gardar Province have been described by Parsons et al. (1991), Finch et al. (1995), Coulson (1997), Markl (2001), Markl et al. (2001) and Markl & Baumgartner (2002). The late-stage fluids expelled from alkaline to peralkaline intrusions are highly enriched in alkalis and incompatible elements. This may result in the formation of aureoles of fenite around alkaline intrusions by interaction with the surrounding country rocks (e.g. Rock, 1976; Kunzendorf et al., 1982; Kresten, 1988; Morogan, 1989) or the autometasomatic formation of secondary mineral assemblages at the expense of primary magmatic minerals within the solidified part of the intrusion itself (e.g. Salvi & Williams-Jones, 1990; Boily & Williams-Jones, 1995; Chakhmouradian & Mitchell, 2002). Recent studies indicate that the sources and isotopic compositions of such a fluid phase may be highly variable (e.g. Boily & Williams-Jones, 1994; Bea et al., 2001) and different isotope systems show variable behaviour with regard to late- or post-magmatic alteration. The resulting isotopic disequilibria between different minerals within a single rock sample may be used to distinguish between primary magmatic and secondary late- to post-magmatic processes.
The Puklen complex of the Gardar Province, South Greenland, is an example of a silica-oversaturated alkaline to peralkaline intrusion comprising a heterogeneous suite of mostly quartz-bearing syenites to quartz-rich peralkaline granite. Petrographically, all rock types show a primary magmatic and a secondary late- to post-magmatic mineral assemblage. The present study is focused on the phase equilibrium constraints on crystallization parameters, on whole-rock geochemistry and on stable and radiogenic isotope compositions of mineral separates, which are used to constrain the magma sources and to decipher the magmatic and late- to post-magmatic processes in the Puklen complex. We show that whole-rock isotope data for these peralkaline rocks would be difficult to interpret and are inadequate for derivation of genetic models.
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REGIONAL GEOLOGY |
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The Gardar Igneous Province (Fig. 1a) in South Greenland represents a failed rift structure of Mid-Proterozoic (1·1–1·3 Ga) age (Upton & Emeleus, 1987). Early Proterozoic (1·7–1·8 Ga) basement granites and gneisses (Emeleus & Upton, 1976) are in places overlain by a sequence of early Gardar basalts and sandstones (Eriksfjord Formation; Poulsen, 1964). A large number of Gardar dyke rocks and about 12 major alkaline to peralkaline igneous complexes intrude the Ketilidian basement. Fluid inclusion data (Poulsen, 1964; Konnerup-Madsen & Rose-Hansen, 1984) and the preserved contacts between sediments and lavas of the Eriksfjord Formation and intrusions indicate that at least the Ilímaussaq intrusion but probably the others as well were intruded at a high crustal level of 3–5 km. The plutonic complexes are composed of (in order of decreasing abundance) syenites, nepheline syenites, alkali granites, gabbros, syenogabbros and carbonatites. With one exception (the granitic to agpaitic Ilímaussaq intrusion), the major plutonic complexes follow either a SiO2-undersaturated trend from just saturated syenites to foyaites and peralkaline or agpaitic nepheline syenites, or an oversaturated trend from augite syenites to peralkaline granites. The Puklen complex belongs to the second group.
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The Puklen complex (Fig. 1b) is a relatively small body (4 km x 2 km), which intrudes Ketilidian basement granite and early Gardar dolerite dykes. It is cut by NW-trending basic post-Gardar dyke rocks (Pulvertaft, 1961). Hence, the Puklen complex is considered to be of late Gardar age (Pulvertaft, 1961). Zircons from a pegmatite in the nearby Nunarssuit complex have been dated at 1171 ± 5 Ma (Finch et al., 2001). Figure 1b shows a generalized geological map of the complex. The field geology was described by Parsons (1972). The first magma pulse formed a suite of coarse-grained, fine-grained and porphyritic varieties of silica-saturated to -oversaturated augite syenites. Contacts between the various syenite types are in most cases gradual. Hence, the intrusion of the various syenites probably took place more or less contemporaneously. In the southern part of the intrusion, a fine-grained and leucocratic granophyre cuts the adjoining syenite. A second pulse of magma produced a homogeneous, coarse-grained peralkaline granite, which grades into or may be locally intruded by fine-grained and leucocratic microgranite.
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SAMPLES AND ANALYTICAL METHODS |
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Six syenites, two granophyres, three coarse alkali granites and two microgranites were analysed for their mineral chemical composition, whole-rock chemical compositions and
18O values; mineral separates of clinopyroxene and amphibole from selected samples were analysed for O, Sr and Nd isotope compositions. The samples were selected to cover the range of textural varieties of the different rock types in a representative way. The sample localities are shown in Fig. 1b. Additionally, two samples of hydrothermal quartz were analysed for their 18O values: sample Q-M is from a centimetre-thick quartz vein and sample Q-D from a cavity containing centimetre-sized euhedral quartz crystals. Both samples were collected from the same locality as sample GM1635.
Minerals were analysed using a JEOL 8900 electron microprobe at the Institut für Geowissenschaften at the Universität Tübingen. For calibration both natural and synthetic standards were used. The beam current was 15 nA and the acceleration voltage was 15 kV. The counting time on the peak was 16 s for major elements, and 30–60 s for minor elements (Mn, Ti, Zr, F, Cl). Background counting times were half of the peak counting times. The peak overlap between the Fe Lß and F K
lines was corrected for. To avoid Na migration under the electron beam, feldspar was analysed using a defocused beam of 15 µm diameter. Data reduction was performed using the internal 
Z procedures of JEOL (Armstrong, 1991).
Bulk compositions of coarsely exsolved feldspars and titanomagnetites were recalculated by combining image processing (NIH Image software) of back-scattered electron (BSE) images of the exsolved minerals with point analyses of the exsolved phases [see Marks & Markl (2001) for a more detailed description of the technique]. Bulk compositions for each mineral and sample were calculated using 3–5 grains.
Trace element contents were measured by in situ laser ablation–inductively coupled plasma-mass spectrometry (LA–ICP-MS) at the EU Large Scale Geochemical Facility (University of Bristol) using a method similar to that described by Halama et al. (2002). The precision of trace element concentrations, based on repeated analyses of standards, was approximately ±5%. The detection limit for Li was typically 150–250 ppm, for Rb and Sr 0·1–1 ppm, and for REE <0·1 ppm.
Whole-rock analyses were performed by standard X-ray fluorescence (XRF) techniques at the Institut für Mineralogie, Petrologie und Geochemie at the Universität Freiburg, using a Philips PW 2404 spectrometer. Pressed powder and Li-borate fused glass discs were prepared to determine contents of trace and major elements, respectively. The raw data were processed with the standard XR-55 software of Philips. Relative standard deviations are <1% and <4% for major and trace elements, respectively. Detection limits vary between 1 and 10 ppm, depending on the specific trace element.
Oxygen isotope compositions of powdered whole-rock samples were determined by a conventional method modified after Clayton & Mayeda (1963) and Vennemann & Smith (1990), using BrF5 as reagent and converting the liberated oxygen to CO2.
The oxygen isotope composition of hand-picked mineral separates was measured using a method similar to that described by Sharp (1990) and Rumble & Hoering (1994). Between 0·5 and 2 mg of sample were loaded onto a small Pt sample holder and evacuated to a vacuum of 10-6 mbar. After prefluorination of the sample chamber overnight, the samples were heated with a CO2 laser in an atmosphere of 50 mbar of pure F2. Excess F2 was separated from O2 by exchange with KCl held at 150°C. The extracted O2 was collected on a molecular sieve (13X). Oxygen isotopic compositions were measured on O2 using a Finnigan MAT 252 mass spectrometer. The results are given in the standard -notation, expressed relative to VSMOW in permil (). Replicate oxygen isotope analyses of the standards (12 loads of NBS-28 quartz and 10 loads of UWG-2 garnet; Valley et al., 1995) had an average precision of ±0·1 for 18O (±2
error of the mean). The accuracy of 18O values was better than 0·2 compared with accepted 18O values for NBS-28 of 9·64 and UWG-2 of 5·8.
For Sr and Nd isotope analyses, about 10 mg of hand-picked mineral separate were spiked with mixed 84Sr–87Rb and 150Nd–149Sm tracers before dissolution under high pressure in HF at 180°C in polytetrafluoroethylene (PTFE) reaction bombs. Rb and Sr were separated in quartz columns containing a 5 ml resin bed of AG50W-X12, 200–400 mesh, equilibrated with 2·5N HCl. Sm and Nd separation was performed in quartz columns using 1·7 ml Teflon powder coated with HDEHP (di-ethyl hexyl phosphate) as cation exchange medium, equilibrated with 0·18N HCl. All analyses were made by thermal ionization mass spectrometry (TIMS) using a Finnigan MAT 262 system in static collection mode. Sr was loaded with a Ta–Hf activator and measured on a single W filament. Rb, Sm and Nd were measured with a double Re-filament configuration. The 87Sr/86Sr ratios were normalized to 86Sr/88Sr = 0·1194, the 143Nd/144Nd ratios to 146Nd/144Nd = 0·7219, and the Sm isotopic ratios to 147Sm/152Sm = 0·56081. Repeated analyses of Ames metal (Geological Survey of Canada, Roddick et al., 1992) gave a 143Nd/144Nd ratio of 0·512142 ± 22 (±2m, n = 10) and of the NBS 987 Sr standard yielded a 87Sr/86Sr ratio of 0·710264 ± 16 (±2m, n = 8). Total procedural blanks (chemistry and loading) were <200 pg for Sr and <100 pg for Nd. A decay constant of 1·42 x 10-11 a-1 for 87Rb (Steiger & Jäger, 1977) and of 6·54 x 10-12 a-1 for 147Sm (Lugmair & Marti, 1978) were used. Nd values were calculated using present-day CHUR values of 0·1967 for 147Sm/144Nd (Jacobson & Wasserburg, 1980) and 0·512638 for 143Nd/144Nd (Goldstein et al., 1984). Initial Sr and Nd isotope ratios were calculated for an age of 1170 Ma, on the basis of U–Pb ages on zircons from the Nunarssuit intrusion, which is close to the Puklen complex and consists of similar rock types (Finch et al., 2001). Calculated uncertainty in Nd units based on analytical errors is not more than 0·5. The error based on age uncertainty is of the order of 0·5–1·0 Nd unit for ages, which are 100 Myr younger or older, depending on the Sm/Nd ratio.
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PETROGRAPHY |
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The following description of the rock types is a brief summary. Detailed and comprehensive petrographic descriptions have been given by Parsons (1972). The mineralogy of the investigated samples is summarized in Table 1. Abbreviations used in the text and figures are given in the Appendix.
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Syenite suite
Syenites are highly variable with respect to colour, grain size, texture and modal mineralogy. Most varieties are equigranular; porphyritic types with subhedral feldspar phenocrysts are less common. The syenites are highly variable in quartz content, but nepheline-bearing syenites do not occur. In some places euhedral centimetre-sized quartz crystals in miarolitic cavities can be found. The dominant matrix feldspar is mesoperthite. In some large feldspar grains, unexsolved cores are still preserved. Some of the quartz-free samples contain interstitial albite. Primary mafic minerals are olivine, clinopyroxene (augite), Fe–Ti oxides, amphibole (amphibole I), apatite and zircon. Olivine, which is now replaced by orange to red iddingsite, is restricted to quartz-free samples. Parsons (1972) noted the rarity of fresh olivine. Subhedral augitic clinopyroxene is grey to green and some grains are strongly zoned. The primary Fe–Ti oxide in most syenites is ilmenite. Some samples show a magmatic two-oxide assemblage of ilmenite and titanomagnetite, and in two samples, titanomagnetite is the only primary Fe–Ti oxide. Primary titanomagnetite is always oxy-exsolved to ilmenite and magnetite in both trellis and less frequently sandwich-type forms [terminology after Buddington & Lindsley (1964)]. Dark brown to dark green amphibole (amphibole I) (Fig. 2a and b) occurs as interstitial crystals or as overgrowths on augite (Fig. 2b). Apatite is found as inclusions in alkali feldspar, augite and amphibole I. Zircon forms small subhedral crystals enclosed in alkali feldspar. In virtually all syenite samples, the primary magmatic phase assemblage is overprinted by a late-stage peralkaline phase assemblage. Augite may show green patches or rims of aegirine-rich pyroxene (Fig. 3). Amphibole I is overgrown by a later pale green amphibole (amphibole II; Fig. 2a) and the primary feldspars developed fine-grained albite along grain boundaries (Fig. 2c).
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Two samples without primary quartz (GM1615 and GM1616) contain aenigmatite. Within these samples, augite, ilmenite and amphibole I form clusters in the feldspar matrix. Albite, secondary quartz II and Ca–Fe carbonates are common as interstitial mineral phases. Aenigmatite forms rims around ilmenite (Fig. 2d). Cracks within aenigmatite are filled with fine-grained titanite and Fe-hydroxides. Titanite also occurs along the contact between ilmenite and aenigmatite. Augite close to aenigmatite is converted to bright green aegirine–augite (Fig. 3).
In sample GM1635, a green rim of aegirine–augite has a distinct grain boundary against the primary augite core (Fig. 3b). The rim can be divided into an inner (aegirine–augite I) and outer part (aegirine–augite II). The inner part contains rounded relics of primary augite (white dashed line in Fig. 3b) and shows patchy irregularities in colour. The boundary between inner and outer parts probably marks the former grain boundary of the primary augite (red dashed line in Fig. 3b). The aegirine–augite rim is overgrown by subhedral hydroandradite with oscillatory zoning patterns, which is associated with subhedral titanite and Ti-free magnetite. Primary ilmenite occurs as inclusions within augite or is overgrown by titanite. The matrix consists of recrystallized quartz and albite.
In sample GM1611, primary augite is entirely replaced by aegirine (Fig. 2e), (calcite–siderite)ss, quartz and haematite pseudomorphs after former augite. Coarse-grained albite and (calcite–siderite)ss are arranged around aegirine. Neither amphibole I nor amphibole II is present. The matrix feldspar is mesoperthite.
Granophyre
The rock is highly leucocratic and fine grained. It consists mainly of euhedral perthitic alkali feldspar overgrown by graphic intergrowths of quartz and alkali feldspar (Fig. 2f). Mafic minerals are rare and comprise augite, ilmenite, and amphibole I and II. Textures of the mafic minerals are similar to those in syenites. Interstitial zircon and especially fluorite are abundant accessory phases.
Alkali granite
Primary magmatic minerals in this rock are alkali feldspar, quartz, ilmenite and amphibole. Minor minerals and accessories are aegirine, astrophyllite, apatite, zircon, fluorite and titanite. Feldspar is mesoperthitic and graphic intergrowths of quartz and feldspar are common. Grain boundaries between quartz and alkali feldspar may be filled with granular masses of fine anhedral albite (Fig. 2c), in some cases associated with tiny needles of aegirine. Ilmenite occurs as inclusions in amphibole or aegirine. Amphibole is subhedral deep blue to dark grey arfvedsonite. Commonly it is overgrown and replaced by bottle-green aegirine (Fig. 2g). Aegirine also occurs interstitially, forming radially arranged aggregates that are partly associated with astrophyllite (Fig. 2h). Apatite is enclosed in feldspar and in amphibole. Zircon forms small subhedral grains that may be clustered into bigger groups. In some samples, titanite is an interstitial phase. Fluorite occurs as anhedral inclusions in aegirine.
In some samples small veinlets of quartz cut the early magmatic minerals. This feature corresponds to field observations of some quartz veins and lenses of 10 cm thickness in or close to the granite.
Microgranite
This rock is leucocratic and fine grained. Primary magmatic minerals are alkali feldspar, quartz, magnetite and arfvedsonite. Aegirine occurs as radiating aggregates or overgrows deep blue amphibole. However, in the same samples, amphibole may overgrow aegirine, implying that the two mafic minerals crystallized alternately or even may have coexisted in parts of the microgranites. In some of the microgranitic veins, zircon and astrophyllite are remarkably common. The latter forms fringes of small yellow needles around aegirine. Zircon occurs interstitially or as inclusions in quartz. In one of the microgranitic dykes (about 2 cm thick), compositional zoning is marked by the occurrence of amphibole and zircon in the inner parts of the dyke, whereas the margins are rich in aegirine.
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RESULTS |
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Mineral chemistry
Feldspar
Measured and recalculated bulk feldspar compositions of the Puklen rocks are shown in Fig. 4a–c. Some typical analyses are reported in Table 2. Feldspar phenocrysts in syenites are partly chemically zoned and range in composition between Ab74An7Or19 and Ab53An2Or45. Feldspar phenocrysts in samples GM1615 and GM1616 are lower in An component compared with the other syenite samples (Fig. 4a). In one sample (GM1580), early ternary feldspar is essentially unzoned, but shows a strong and steep enrichment of Ab component with almost unchanged An component at the rim (Fig. 4b). The most common matrix feldspar in the syenites is almost Ca-free patchily exsolved alkali feldspar with bulk compositions between Ab69An0Or31 and Ab43An0Or57. Late-stage albite in syenites has nearly end-member composition. The low Ca contents clearly distinguish these fine-grained aggregates from the Ab-rich rims around phenocrysts in sample GM1580 (see above).
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Feldspar compositions in granophyres range from Ab57An0Or43 to Ab44An0Or56 and are more Or rich compared with the syenites. In one sample (GM1605) of coarse alkali granite, the composition of the unexsolved core of an alkali feldspar was determined as Ab69An4Or27–Ab65An2Or23 (Fig. 4c). Compositions of matrix feldspars in coarse granites are similar to those in syenites and range between Ab63An1Or26 and Ab49An0Or51. As in syenites, late-stage albite shows almost end-member composition. Feldspar compositions in the microgranites (Ab72An1Or27–Ab49An0Or51) extend the range towards more albite-rich compositions. As in the granophyres, late-stage albite is lacking.
Olivine
The composition of olivine could not be determined by electron microprobe, as it is altered to red–orange iddingsite. Olivine in such rock types is expected to be fayalite rich (e.g. Stephenson, 1974; Larsen, 1976; Powell, 1978; Upton et al., 1985; Marks & Markl, 2001).
Pyroxene
The primary pyroxene in the syenites and granophyres is augite with >90 mol % quadrilateral (Di + Hed + En + Fs) components (Fig. 5, Table 3). The Na content of the primary augite varies between 0·03 and 0·15 atoms per formula unit (a.p.f.u.) but exceeds in some analyses Fe3+, which was calculated based on stoichiometry (four cations, six oxygens). This indicates the presence of small amounts of the jadeite molecule (up to 6 mol %) in addition to the aegirine component. Some augites show chemical zonation (Fig. 5b). XFe and Mn increase from core to rim, continuously in some crystals but stepwise in others. The Wo component is more or less constant, Na shows a continuous and smooth enrichment, and Ti decreases. In cracks or rims on augite, the aegirine component rises to 40 mol % (Fig. 5a). In such areas, augite is enriched in Na, whereas Ca, Mg and Ti are depleted.
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In the texture of Fig. 3b, the inner aegirine–augite I shows patchy irregularities in chemical composition. Similar to the other samples, it is enriched in Na and depleted in Ca, Mg and Ti. The outer aegirine–augite II shows oscillatory zoning with respect to Na and other elements and is enriched in Al compared with the primary augite (not shown). In syenite sample GM1611 (Fig. 2e), the chemical composition of the aegirine varies between Aeg72Di3Hed25 and Aeg95Di0Hed5 and lies within the range observed in the granites (see below).
Pyroxene in coarse granites and microgranites is aegirine with compositions between Aeg76Di4Hed20 and Aeg92Di0Hed8 (Fig. 5a, Table 3). Aegirine is essentially unzoned and there is no significant compositional difference between aegirine in the coarse granites and in microgranites. Al is the most important minor element; the jadeite component makes up as much as 6 mol %.
Fe–Ti oxides
Table 4 reports some typical analyses of ilmenite and titanomagnetite in the various rock types. The composition of primary ilmenite in syenites varies between Ilm80Hem6Pyr14 and Ilm95Hem1Pyr4. Recalculated compositions of primary Ti-magnetite grains in syenites range between Usp72Mag28 and Usp97Mag3. Such Ti-rich compositions have already been reported from the augite syenite of the Ilímaussaq intrusion (Marks & Markl, 2001). Granophyric samples contain ilmenite with compositions between Ilm84Hem5Pyr11 and Ilm95Hem2Pyr3.
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Ilmenite in coarse granites is rich in MnO and varies between Ilm91Hem0Pyr9 and Ilm64Hem6Pyr30. The only primary Fe–Ti oxide observed in microgranites is magnetite with Usp contents <5 mol %.
Secondary Fe–Ti oxides in late- to post-magmatic textures are essentially Ti-free magnetite in some syenite samples and end-member haematite in granites.
Amphibole
Amphibole I in syenites is ferro-richteritic to ferro-edenitic in composition (Table 5). XFe in the entire suite ranges from 0·72 to 0·99. Fluorine (up to 3·5 wt %) always dominates over chlorine (<0·5 wt %). Some grains show a pronounced chemical zonation. XFe, Si, Na and Cl increase from core to rim, whereas Al, Ca and F decrease (Fig. 6a). Amphibole I in fluorite-bearing granophyres is ferro-edenite and essentially fluorine free. The chlorine content is in the same range as in the syenites.
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Amphibole II is a ferro-actinolite. XFe varies in the same manner as in amphibole I (Table 5). Figure 6a shows a zonation profile starting in the core of an amphibole I crystal and extending into a rim of amphibole II. Compared with amphibole I, these late-stage amphiboles are lower in Na, K, Mn and Ti, and they are also depleted in halogens. Cl is always <0·1 wt %, and F is below microprobe detection limit.
Amphibole in coarse granites and microgranites is a member of the arfvedsonite–ferro-leakeite series (Table 5). XFe varies between 0·69 and 0·98. Lithium contents vary from <0·1 to 1·0 wt % Li2O. Li is negatively correlated with Fe2+ and Al(VI), but positively with Fe3+ (Fig. 6b). The main substitution mechanisms for Li in the Puklen arfvedsonites is Li + Fe3+ 
2(Fe2+, Mg, Mn) and Li + Fe3+ Al(VI) + Na. Cl is always <0·05 wt % whereas F may be up to 3·5 wt %. F is negatively correlated with XFe.
Aenigmatite
According to the scheme of Kunzmann (1999), the Puklen aenigmatites vary in composition between Aen78Wilk18Rhö4 and Aen93Wilk4Rhö3 (Aenigmatite–Wilkinsonite–Rhönite). The main inferred substitution mechanisms are Fe2+ + Ti4+ 2 Fe3+ and Ca2+ + Al3+ Na+ + Si4+. With increasing distance from precursor ilmenite, Ca and Al contents decrease, whereas Na, Si and Mn increase.
Carbonates
Carbonates are essentially Mg-free calcite–siderite–rhodochrosite solid solutions. Tiny carbonate inclusions in replaced augite are calcite-rich and vary in composition between Cc97Sid2Rhod1 and Cc87Sid5Rhod8. Large interstitial carbonate coexisting with albite is more siderite-rich (Cc72Sid25Rhod7–Cc64Sid30Rhod6).
Andradite
Hydroandradite shows oscillatory zoning (Fig. 3b) and varies in composition between Andr67Gr32Sp1Alm0Py0 and Andr91Gr6Sp1Alm2Py0.
Trace element data for metasomatized augite of sample GM1616
Figure 7 shows selected trace element data for an augite crystal in sample GM1616. On the basis of petrographic observations and mineral compositions it is suggested that late-stage metasomatic fluids have affected the augites of this sample. The unaffected core region (left side of Fig. 7b) has low Rb/Sr ratios between 0·1 and 0·4. In contrast, in the aegirine-rich metasomatized areas (right side of Fig. 7b), both Sr and Rb are enriched by about one and two orders of magnitude, respectively. Rb/Sr ratios increase with increasing Na (i.e. increasing metasomatism) to >1 in metasomatized parts of the crystal. Concentrations of Sm and Nd and Sm/Nd ratios are shown in Fig. 7c for the same analysed points. Both elements are enriched in the metasomatized areas of the crystal, but less significantly than Rb and Sr. Enrichment factors compared with unaffected core regions are about 1·2 for Sm and 1·7 for Nd, resulting in lower Sm/Nd ratios in the metasomatized areas than in unaffected core regions.
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Whole-rock geochemistry
Major and trace element analyses of the Puklen rocks are given in Table 6. Some analyses have fairly low totals, which may be attributed to the presence of H2O+, CO2 and halogens, which were not measured. Most samples are peralkaline with an agpaitic index [molar (Na2O + K2O)/Al2O3] between 0·97 and 1·25. The samples have low contents of MgO and CaO, and high concentrations of iron and alkalis. A significant gap in SiO2 content exists within the Puklen rock series, between syenites and the other rock types (see also Parsons, 1972).
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Contents of Al2O3, CaO, FeOt, TiO2, MnO and P2O5 decrease with increasing SiO2 content from syenites to the other rock types (Fig. 8a). This may suggest fractional crystallization of plagioclase, olivine, augite, magnetite–ilmenite and apatite. However, changes in element concentrations are not always systematic. For example, microgranites, which clearly postdate the coarse alkali granites, have higher contents of CaO and Al2O3 than coarse alkali granites.
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In general, concentrations of compatible trace elements such as Sc, V, Cr, Co and Ni are low. Of the high field strength elements, Zr is strongly enriched in alkali granites. The concentrations of Zr and Zn correlate positively with the agpaitic index, whereas Sr and Ba concentrations correlate negatively with the agpaitic index (Fig. 8b), which may be attributed to extensive feldspar fractionation.
Oxygen isotope data
Whole-rock and mineral 18O values are given in Table 7 and plotted in Fig. 9. The 18O values of the syenites span a large range between +4·8 and +6·9. The values for the two analysed granophyres are +4·6 and +5·0, for the three coarse alkali granites are between +5·0 and +5·8, and for the two microgranites are +4·9 and +5·5.
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In syenites and granophyres, the
18O values of mineral separates from individual samples decrease in the order quartz–augite–amphibole I–amphibole II. 18O values of quartz (+7·9 to +8·1), amphibole I (+3·9 to +4·6) and amphibole II (+2·5 to +2·8) in syenites are rather homogeneous and significantly higher than in the granophyres (+6·4 to 6·7, +2·5 to +2·7 and +1·0 to +1·7, respectively). An explanation for the homogeneous 18O values of quartz but the heterogeneous whole-rock 18O values will be given later. With the exception of syenite sample GM1616 (+4·7), augite in the syenites is also homogeneous, varying between +5·3 and +5·5. Augite in the granophyric sample GM1593 (+5·4) has a 18O value in the range of augite in the syenites. Two separates of perthitic feldspar from syenites have low 18O values of +3·5 and +3·9.
In coarse alkali granites, 18O values decrease from quartz to amphibole to aegirine. The range in measured 18O values for all three minerals is large (quartz: +5·1 to +7·3; amphibole: +2·2 to + 3·5; aegirine: +0·5 to + 2·0) compared with those in the syenites.
In microgranites, the 18O values of quartz (+6·7 and +7·2) and aegirine (+1·8) are in the same range as those in the coarse alkali granites, but amphibole has slightly lower values (+1·7 and +1·9). Both quartz and amphibole have significantly lower 18O compared with the same minerals in the syenites. The two samples of late-stage quartz veins have 18O values similar to quartz in granites (+7·0 and +7·2).
Nd isotope data
Thirteen mineral separates from eight samples were analysed for their Sm and Nd concentrations and their Nd isotopic compositions (Table 8). The calculated Nd values are all negative and highly variable. The minerals in the syenites cover a fairly large range in Nd between -3·8 and -6·4, whereas those in the alkali granites show lower and more variable values between -5·9 and -9·6. Intermediate values of -6·5 for a microgranite amphibole and -7·2 for a granophyre augite were determined. Syenites and the two alkali granites show contrasting Nd isotope behaviour. Although in a particular syenitic sample, augite and amphibole have approximately homogeneous Nd, the variability in the Nd isotope composition of the minerals between samples is large. In the two alkali granite samples, the opposite is observed: amphibole and aegirine show almost identical values in the two samples.
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Sr isotope data
Six mineral separates from two syenites and two coarse alkali granites were analysed for their Rb and Sr contents and Sr isotopic compositions (Table 8). Rb/Sr ratios of late-stage aegirine (samples GM1605 and GM1606) and metasomatic augite in sample GM1616 are higher compared with those of the primary and unaffected minerals. Augite from sample GM1590 has a highly radiogenic initial 87Sr/86Sr of 0·730, whereas augite from sample GM1616 has the lowest value of 0·590. Despite the higher Rb/Sr and 87Rb/86Sr ratios of GM1616, the present-day 87Sr/86Sr value is low. Amphiboles from alkali granites have low initial Sr isotope ratios between 0·699 and 0·696. Late-stage aegirine has even lower values of 0·642–0·631. If ages younger than 1170 Ma are assumed, the calculated initial Sr ratios change only slightly. Only ages younger than
900 Ma, which are not known from the Gardar Province (Upton & Emeleus, 1987), lead to geologically realistic initial Sr ratios of >0·700–0·703 for depleted sources or >0·703 for isotopically enriched sources (Bulk Earth at 1170 Ma = 0·703). This indicates that the calculated low initial Sr ratios are not an effect of a wrong age assumption, but reflect disturbance of the Rb–Sr system. |
CONDITIONS DURING THE MAGMATIC STAGE |
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Based on fluid inclusion studies (Konnerup-Madsen & Rose-Hansen, 1984; Markl et al., 2001) and on the reconstruction of the sedimentary and extrusive igneous overburden (Upton, 1962; Harry & Pulvertaft, 1963; Poulsen, 1964; Emeleus & Upton, 1976) the pressure of emplacement of the Ilímaussaq complex was estimated to be
1 kbar. In the lack of better estimates, we use this value as an approximation for the Puklen complex as well.
Syenites and granophyres
Temperature and silica activity
The composition of early ternary feldspar phenocrysts in some syenite samples can be used to constrain near-liquidus conditions in the parental melt. The minimum crystallization temperature of the microperthites was determined by Parsons (1972) to be in excess of 715°C. Figure 4 shows feldspar compositions plotted on the temperature-dependent feldspar solvus after Elkins & Grove (1990). Estimated minimum temperatures range from 750° to 950°C.
Equilibria between olivine, pyroxene and melt constrain silica activity of the olivine-bearing samples, which were originally quartz-free. Unfortunately, no primary olivine was preserved in the investigated samples. The QUILF program (Andersen et al., 1993) calculates equilibria involving olivine, augite, quartz, magnetite and ilmenite, and was used to estimate the composition of olivine in equilibrium with the measured augite at temperatures between 750° and 950°C. Detailed information on the theory and application of QUILF has been given by Frost & Lindsley (1992), Lindsley & Frost (1992) and Marks & Markl (2001). Calculated olivine compositions range from Fa67Fo31La2 to Fa96Fo3La1. Because this range is similar to that for olivines from other augite syenites of the Gardar Province (Stephenson, 1974; Larsen, 1976; Powell, 1978; Upton et al., 1985; Marks & Markl, 2001) we believe that these calculations are reliable and can be used to estimate silica activity for these samples.
Calculated silica activities are based on the reference state of pure quartz at P and T. Silica activity evolved systematically from 0·70 in the syenite samples with the lowest XFe in augite to 0·98 in those that contain nearly end-member hedenbergite (Fig. 10). This systematic relationship between XFe in augite and calculated silica activity in the olivine-bearing syenites indicates that an originally quartz-undersaturated melt evolved by fractionation of olivine, augite and Fe–Ti oxide along a displaced FMQ buffer towards quartz saturation. All other syenites and granites are quartz bearing but lack olivine. The above-mentioned trend of the quartz-free samples (Fig. 10) is not applicable to the quartz-bearing ones. In contrast to what one would expect in a closed system, the quartz-bearing syenites do not contain the Fe–Mg silicates with the highest XFe, but quartz-bearing syenites show the same range of XFe in augite as quartz-free samples.
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Oxygen fugacity
Oxygen fugacity (fO2) was calculated from equilibria among Fe–Ti oxides and Fe–Mg silicate minerals (olivine, augite) using the QUILF program (Andersen et al., 1993). For each sample, calculations were performed for the whole compositional range observed and consequently they yielded a range of fO2 values. Estimated fO2 is always below the synthetic FMQ buffer and varies between 0·8 and 2·3 log units below the FMQ buffer. These and even more reduced conditions seem to be typical for the Gardar syenites (Powell, 1978; Marks & Markl, 2001). Oxygen fugacity in granophyres was estimated to be around or slightly above FMQ at temperatures between 650° and 750°C.
Granites
Temperature
In coarse granites, the rarely preserved homogeneous cores of alkali feldspar indicate minimum temperatures of 750°C (Fig. 4c). The upper stability limit for F-free arfvedsonite is 700°C (Bailey, 1969). F-rich arfvedsonite is an early liquidus phase in the Puklen granites and F is expected to increase the thermal stability of arfvedsonite significantly because experimental data for richteritic compositions (Gilbert & Briggs, 1974) show a difference in the F and OH stabilities of 300°C in the low-pressure range. Thus, the crystallization of the magmatic assemblage of the granites occurred at temperatures 
750°C.
Oxygen fugacity
The only constraint on oxygen fugacity for the magmatic stage of the coarse granites is the occurrence of arfvedsonite. End-member arfvedsonite is stable only at conditions below the synthetic FMQ buffer (Bailey, 1969).
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LATE-STAGE PROCESSES IN SYENITES |
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Formation of aegirine–augite and aenigmatite in syenites
Crystallization of anhydrous minerals such as alkali feldspar, olivine, augite and Fe–Ti oxides under relatively reduced conditions (see above) led to the enrichment of Na, Fe, Si, halogens and H2O in the residual melt. As a result of increasing H2O activity, amphibole I began to crystallize as a late magmatic phase. Enrichment of Na and depletion of Al during fractionation is also indicated by the rise of (Na + K)/Al ratio from core to the rim in augite and amphibole (Figs 6a and 11). Upon reaching fluid saturation, a fluid phase (presumably H2O–CO2) must have been exsolved, which is suggested by the formation of carbonate minerals. Based on petrographical observations that late- to post-magmatic reactions mainly took place along cracks or grain boundaries and carbonates precipitated at the same time, we conclude that this fluid phase was responsible for these reactions. Because one of the most important replacement textures in our samples is the growth of aegirine at the expense of augite (Fig. 3), we assume this fluid phase was very probably Na dominated. The replacement reaction can be described as follows:
 | (1) |
 | (2) |
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The formation of aenigmatite at the expense of ilmenite was observed in the two samples (GM1615 and GM1616) showing most extensive growth of Na-pyroxene (Figs 2d and 5a). The following schematic reaction describes this process:
 | (3) |
space where aenigmatite and aegirine coexist in the absence of Fe–Ti oxides. This non-oxide field is shown in Fig. 12 (grey field) and is bounded by the labelled reactions (a)–(c). It should be noted that reaction curves in Fig. 12 are for constant unit activities of aegirine, magnetite, ilmenite, aenigmatite and sodium disilicate. Following Nicholls & Carmichael (1969), point A in Fig. 12 represents the intersection of reactions (a)–(c) for an aegirine activity of 0·5, and the dashed line indicates the intersection of reactions (a) and (b) depending on the activity of sodium disilicate. Incorporation of Fe3+ into aenigmatite shifts reaction (a) to more oxidized conditions and thus expands the non-oxide field to the vicinity of the haematite–magnetite (HM) buffer curve. In summary, the formation of aenigmatite and stabilization of aegirine–augite in some of the Puklen syenites indicates an increase of oxygen fugacity as a consequence of cooling (curved arrow in Fig. 12).
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Formation of late-stage Ca-minerals in syenites
Replacement of early magmatic Ca-bearing minerals in syenites by late-stage Na-rich minerals led to the release of Ca into the fluid phase and subsequently to the formation of the secondary Ca-bearing minerals ferro-actinolite, (calcite–siderite)ss, andradite and titanite. In some samples Ti-free magnetite and quartz are associated with these minerals. Three associations can be distinguished:
- (calcite–siderite)ss + titanite + magnetite + quartz;
- titanite + hydroandradite + magnetite + quartz;
- ferro-actinolite.
Temperature constraints
Carbonate-bearing samples GM1615 and GM1616. Large interstitial Fe-rich carbonates associated with aenigmatite in samples GM1615 and GM1616 were used for solvus thermometry after Goldsmith et al. (1962) and the re-evaluation of Anovitz & Essene (1987). The inferred minimum temperatures for these compositions are not well constrained and range between 650° and 750°C depending on the solvus used. However, these temperatures probably mark the beginning of fluid activity at relatively high temperatures.
Sample GM1611. Carbonates that were found as tiny inclusions in completely metasomatized augite in sample GM1611 indicate minimum temperatures as low as 360–320°C and are better constrained because the solvus curve for such Fe-poor compositions is reasonably well determined (Anovitz & Essene, 1987).
Carbonate-free samples. The upper stability limit of ferro-actinolite (at 1 kbar) has been experimentally determined by Hellner & Schürman (1966) to be about 550–600°C. This is the maximum temperature for the beginning of amphibole II formation during cooling.
The above-mentioned temperature constraints indicate that the formation of secondary Ca minerals occurred within a wide temperature range starting at temperatures close to the magmatic stage of 750°C and continuing down to temperatures of 300°C.
Ca-mineral constraints on fO2 and aCO2
Although the principal process of replacement of an early Ca-bearing assemblage by a late Na-bearing one appears to be the same in all samples, the secondary Ca-mineral assemblages are not. We assume that local variations in fO2, aH2O or aCO2 were responsible for this. To investigate this, we calculated an fO2–aCO2 diagram in the Ca–Fe–Si–O–H–C system involving hedenbergite, magnetite, andradite (as an approximation for hydro-andradite), ferro-actinolite, calcite, quartz, CO2 and H2O using the GEOCALC software of Berman et al. (1987) and Lieberman & Petrakakis (1990) and the database of Berman (1988). New thermodynamic data for ferro-actinolite (Ghiorso & Evans, 2002) were added to this database. A set of isothermal log fO2–log aCO2 diagrams was calculated from 300° to 600°C. Magnetite was regarded as a pure end-member in accordance with microprobe analyses. For the calculation of end-member component activities we used the solution models of Holland (1990) for hedenbergite and Cosca et al. (1986) for andradite. For calcite and ferro-actinolite, a ‘mixing on site’ model was used. As an example, the topologic relations for reactions among these minerals are shown for 500°C in Fig. 13. Between 300° and 600°C the topologic relations between reactions are similar but the position of the invariant point [Fe-Act] shifts with falling temperature from log fO2 (600°C) = -22 to log fO2 (300°C) = -38 and from log aCO2 = -0·3 to -2·4 (black dots and dashed lines in Fig. 13). However, it is important to note that the position of the invariant point relative to the FMQ buffer is generally independent of temperature and lies at about 
FMQ = +1 log unit at all temperatures. The fO2–aCO2 diagram can be used to interpret the various Ca-mineral assemblages, as follows:
- the occurrence of andradite in sample GM1635 formed by oxidation of hedenbergite points to a fluid phase with aCO2 <0·25 (log aCO2 <-0·6) (grey area in Fig. 13) if a temperature of 500°C is assumed. Oxygen fugacity in this rock increased from magmatic values (FMQ = -1 to -2) to above FMQ = +1.
- The two aenigmatite- and carbonate-bearing samples GM1615 and GM1616 indicate even higher fO2 and aCO2 values (stippled pattern in Fig. 13).
- Ferro-actinolite-bearing (amphibole II) samples were less oxidized and had lower aCO2 (dotted pattern in Fig. 13) or—if oxygen fugacity was higher than in sample GM1635—they coexisted with a fluid phase unusually rich in CO2, which is considered unlikely.
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Occurrence of titanite
In the two samples GM1615 and GM1616 titanite occurs along grain boundaries between ilmenite and aenigmatite [see reaction (3)] and as fine-grained fillings of small cracks within aenigmatite. These textures indicate that titanite formed at the expense of aenigmatite:
 | (4) |
In the andradite-bearing sample GM1635, titanite coexists with Ti-free magnetite and quartz (Fig. 3b) as an overgrowth on ilmenite. This can be expressed by the classical reaction (Wones, 1989)
 | (5) |
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SIGNIFICANCE OF Li-AMPHIBOLES IN PERALKALINE GRANITES |
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Similar to Mn, Zn or Zr, Li can be an important component in igneous alkali amphiboles, especially in peralkaline granites (Hawthorne et al., 1993, 1994). Compared with other published data (e.g. Bailey et al., 1993; Hawthorne et al., 2001) Li contents in amphiboles of the Puklen granite (0·1–1·0 wt % Li2O) are relatively high, but variable. Comparison with the detailed studies of Hawthorne et al. (1993, 1994, 2001) shows that the dominant substitution mechanism for Li in the alkali amphiboles of the Puklen peralkaline granite is M3Fe2+ + Fe2+
M3Li + Fe3+, giving rise to the ideal end-member ferro-leakeite. The Puklen amphiboles contain up to 60 mol % of this component.
Strong & Taylor (1984) distinguished two compositional trends in amphibole from silica-saturated peralkaline igneous rocks. A magmatic to subsolidus trend is characterized by a change in composition from barroisite to richterite to arfvedsonite. The most important substitution here is [4]AlCa SiNa. This substitution involves amphiboles with full A sites and this trend is proposed to occur under reducing conditions. Second, the so-called oxidation trend reaches riebeckite composition and takes place under the influence of oxidizing fluids. This also produces amphiboles with vacancies on the A site. The Fe3+/(Fe3+ + Fe2+) ratio of the two end-members of both trends are 0·20 and 0·40, respectively. As discussed by Hawthorne et al. (1993), the incorporation of Li by the mechanism mentioned above increases the Fe3+/(Fe3+ + Fe2+) ratio from 0·20 in Li-free arfvedsonite to 0·50 in ferro-leakeite, which is even more oxidized than the oxidation trend of Strong & Taylor (1984). The relatively high Fe3+/(Fe3+ + Fe2+) ratios and high Li contents found in the Puklen arfvedsonites imply an extension of the stability field for these amphiboles to higher oxygen fugacities—even above the FMQ buffer—compared with Li-free arfvedsonite, suggesting that these amphiboles presumably crystallized above the FMQ buffer.
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LATE-STAGE FORMATION OF AEGIRINE IN PERALKALINE GRANITES |
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The formation of aegirine in the Puklen peralkaline granites is a late-stage process and is a well-known phenomenon in peralkaline granites. Most commonly, primary arfvedsonite is replaced by granular aegirine (Fig. 2g), which shows many small inclusions of quartz, fluorite and haematite. This can be described by the following reaction:
 | (6) |
 | (7) |
 | (8) |
T–fO2 constraints on aegirine formation
The stability of aegirine in the presence of water is restricted to conditions between the FMQ and HM buffer curves (Bailey, 1969). The occurrence of Ti-free haematite, and relatively Ti-poor aegirine indicates oxidized conditions at or above the HM buffer. The decomposition of arfvedsonite into aegirine, quartz and haematite was reported in the Puklen rocks by Parsons (1972), and in other peralkaline granites by, for example, Boily & Williams-Jones (1994) and Schmitt et al. (2000). Reaction (6) represents an fO2 buffer in peralkaline rocks and may be used to estimate the conditions of aegirine formation. Figure 14 shows the T–fO2 dependence of this reaction at 1 kbar using thermodynamic data for aegirine, haematite, quartz and water from Robie & Hemingway (1995). fH0 and S0 for arfvedsonite were estimated using the methods of Robinson & Haas (1983) and Chermak & Rimstidt (1989). The molar volume from Hawthorne (1976) of arfvedsonite (28·17 J/bar) was used. Temperature-dependent fugacity coefficients for water are from Burnham et al. (1969). Activities for aegirine, haematite, quartz and water were assumed to be unity. The activity of arfvedsonite varied between 0·1 and 0·15 using a ‘mixing on site’ model. The uncertainty of these calculations is mainly based on the estimation of fH0. It is believed to be in the range of 0·2% (Chermak & Rimstidt, 1989). The grey field around the calculated curves in Fig. 14 marks the uncertainty if an error of ±0·2% for fH0(arf) is assumed. The intersection with the HM buffer curve divides this curve into a metastable (high-temperature) and a stable (low-temperature) branch. Assuming a stable formation of aegirine, these intersections should reflect maximum crystallization temperatures of 260–280°C. If the assumed uncertainty is taken into account, the formation of aegirine took place at temperatures below 450°C.
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Popp & Gilbert (1972) showed that the solubility of jadeite in aegirine at constant pressure increases with falling temperature if aegirine coexists with albite and quartz (which is the case for the Puklen aegirines). They investigated the stability of aegirine–jadeite solid solutions at low pressures for temperatures between 300° and 600°C. Assuming a pressure of 1 kbar, temperatures below
350°C are indicated, which is in good agreement with the calculations above. Whereas Bonin (1986) proposed temperatures of 600–625°C for conversion of arfvedsonite into aegirine + quartz + haematite + H2O, Boily & Williams-Jones (1994) demonstrated, based on oxygen isotope data, that aegirine formed at significantly lower temperatures, in agreement with our estimate.
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RETENTION AND ALTERATION OF PRIMARY ISOTOPE SYSTEMATICS |
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Oxygen isotope fractionation
In this section, we discuss whether the measured oxygen isotope compositions of minerals represent primary magmatic values or the effects of hydrothermal alteration or re-equilibration during slow cooling.
In syenites, oxygen isotope fractionations between different minerals within the same sample are: qtz-aug = +2·6 to +2·8; qtz-am I = +3·5 to +4·2; qtz-am II = +5·1 to +5·6. The fractionations between quartz and augite indicate equilibrium at magmatic temperatures of 750–700°C using the fractionation factors of Zheng (1993a, 1993b). This is in good agreement with the results obtained from feldspar thermometry. The only exception is sample GM1616, with a remarkably large qtz-aug of +3·3, which is due to the exceptionally low 18O value of augite from this sample (+4·7). This low value can be explained by the fact that augite in this sample is strongly affected by late-stage peralkaline fluids. From early magmatic augite to later amphibole I and amphibole II, fractionation increases, giving rise to lower apparent equilibrium temperatures (640–570°C and 500–465°C, respectively).
In granophyres, the measured fractionations between quartz and amphibole I (+3·7 and +4·2) and between quartz and amphibole II (+5·0 and +5·4) are almost identical to those of the syenites, but this value is very small for quartz–augite (+1·0) in sample GM1593. This low value and the resulting high calculated equilibrium temperature of >1300°C indicate non-equilibrium conditions for the two minerals in this sample. Because fractionations between quartz, amphibole I and amphibole II show similar values to those in the syenites, it can be concluded either that the augites of sample GM1593 are xenocrysts or that the quartz and amphibole equilibrated with a distinctly different fluid or melt at lower temperatures, whereas the augite did not. Based on their major element and oxygen isotope composition, we assume that these augites are early crystals and were incorporated from the syenites.
In alkali granite and microgranites, quartz–amphibole fractionations vary between +3·6 and +5·3, yielding low apparent equilibrium temperatures of 640–480°C. It appears that quartz and amphibole are not in isotopic equilibrium at magmatic temperatures of 700°C. Possibly, a late closure to oxygen diffusion of quartz caused by low cooling rates may have caused an increase in qtz-amph. Differences in 18O values between late-stage aegirine and quartz are high (+4·6 to +6·1) and indicate low apparent equilibrium temperatures of <250°C (Zheng, 1993a).
Different effects on the Rb–Sr and the Sm–Nd systems during metasomatism
As mentioned above, the Rb–Sr isotope system has been disturbed and no longer reflects primary magmatic values. Disturbance of the Rb–Sr system is common because Rb and Sr are highly mobile in metasomatic fluids. This has been shown in many whole-rock studies (e.g. Stevenson et al., 1997; Ashwal et al., 2002). The extremely low calculated initial 87Sr/86Sr values indicate loss of radiogenic Sr and/or the addition of Rb. Trace element data for augite of sample GM1616 (Fig. 7) indicate that probably the latter was the case. Additionally, despite the high present-day 87Rb/86Sr ratio, a relatively low present-day 87Sr/86Sr ratio was measured for the augite of sample GM1616. This indicates that the 87Sr/86Sr ratio of the infiltrating late-stage fluid must have been significantly lower than that of the primary magmatic fluid and metasomatism changed both the 87Rb/86Sr and 87Sr/86Sr ratios of the system.
In contrast to the Rb–Sr system, Sm is positively correlated with Nd and samples with high Sm/Nd and 147Sm/144Nd ratios consequently have high present-day 143Nd/144Nd ratios (Table 8). These observations suggest that the Sm–Nd system was not as strongly affected by metasomatism as the Rb–Sr system. As shown in Fig. 7c, the concentrations for both Sm and Nd in metasomatized augite increased by a factor of about two compared with unaffected core regions. This effect is relatively small compared with the Rb–Sr system, where the enrichment of Rb and Sr is between 10 and 100. Probably, the similar atomic sizes and resulting physico-chemical characteristics of Sm and Nd are responsible for the similar behaviour of Sm and Nd during late-magmatic processes compared with the drastic differences between Rb and Sr. Therefore, we conclude that the Nd isotopic compositions can be used for geological interpretations.
Figure 15 summarizes the Nd isotopic data in a conventional isochron diagram. The most obvious feature is that the data do not define a single isochron but a trend line at best. The two separates of aegirine from alkali granites show the most pronounced deviation from the general trend. Omitting the aegirine data, a late Gardar age of 1111 ± 95 Ma is obtained. This is in agreement with the assumption that the Puklen rocks intruded penecontemporaneously with the Nunarssuit complex (Pulvertaft, 1961; Finch et al., 2001). The large uncertainty and high MSWD value (mean of squared weighted deviations) of the isochron could result from the relatively low variation in 147Sm/144Nd ratios, from analytical errors, from a post-magmatic modification of the Sm–Nd system as mentioned above (e.g. Andersen, 1984) or from heterogeneous initial isotopic compositions of the samples as discussed below. However, in three syenitic samples (GM1590, GM1600 and GM1616), two minerals were analysed and in all three cases the calculated Nd values of these separates agree well within error (Table 8). Two-point isochrons defined by these samples agree with the above-mentioned age within error (Fig. 15). This supports the petrological and oxygen isotopic results that suggest an essentially closed-system behaviour for each syenite sample after contamination and during cooling (see below).
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MELT SOURCE AND CONTAMINATION |
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Mineral–melt fractionations allow the calculation of the magma oxygen isotopic compositions directly from measured values of minerals (Taylor & Sheppard, 1986). We used
quartz-melt and pyroxene-melt of Kalamarides (1986), Taylor & Sheppard (1986) and Harris (1995) to calculate 18Omelt values for the Puklen syenites. For granophyres, only quartz-melt was used, as the augite of sample GM1593 is not believed to be in equilibrium with quartz because of unreasonably low quartz-augite. For the coarse alkali granites and microgranites, only quartz-melt was used because aegirine is of hydrothermal origin.
The estimates of 18Omelt vs wt % SiO2, showing broadly a negative correlation, are plotted in Fig. 16. Syenites have 18Omelt values of +5·9 to +6·3 and granophyres have slightly lower values (+5·4 and +5·7). Coarse alkali granites (+5·1 and +5·3) are significantly lower in 18Omelt than the syenites. If it is considered that 18Omelt estimates for granites are maxima, as qtz-am in the alkali granites (see above) indicates a late closure of quartz to oxygen diffusion, the measured 18O values for quartz in granites are likely to be higher than values calculated at the crystallization temperatures of quartz. It is well known that the closure temperature for quartz is, depending on grain size, 500°C (e.g. Giletti & Yund, 1984; Jenkin et al., 1991). During closed-system cooling, the 18O value of quartz will increase relative to other minerals. Therefore, 18Omelt values for granites are possibly even lower than calculated. Microgranites yielded the lowest 18Omelt values (+4·7 and +5·2).
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The calculated
18Omelt value of +5·9 to +6·3 for syenites essentially supports a mantle derivation of the parent magma (Kyser, 1986). The analysed mineral separates from syenites span a large range in Nd values between –3·8 and –6·4 (Table 8). These differences probably indicate that different magma batches derived from the same source experienced variable amounts of crustal contamination with country rocks during emplacement. Consequently, sample GM1600 (Nd = –6·4) may represent a strongly contaminated magma, whereas sample GM1616 possibly reflects the least contaminated magma with an Nd value of –3·8 for primary augite. The high quartz content of sample GM1600 is therefore believed to be due to a high amount of contamination with silicic country rock in this sample and not due to the effects of fractional crystallization.
The differences in inferred 18Omelt values between the syenites and the other Puklen rocks cannot be due to closed-system fractionation processes (Kalamarides, 1986; Baker et al., 2000). Moreover, the data obtained during this study indicate at least two isotopically different sources and possibly a mixing between these two. The low 18Omelt values of the coarse alkali granites and granophyres can be explained only if the magma was derived from a source rock with a low 18O value, or contamination with material of low 18O has occurred. Three whole-rock 18O analyses for basement rocks of the Isortoq area (Fig. 1b) have a mean of +7·8 ± 0·5 (R. Halama, unpublished data, 2002). Consequently, the upper granitic crust of the Gardar Province is not a significant contaminant for the Puklen granites. Other contaminants or sources with low 18O—for example, lower-crustal rocks (Valley, 1986) or sources that were already altered by meteoric fluids before melting (Harris & Ashwal, 2002)—are necessary to explain the oxygen isotopic differences between the syenites and the other Puklen rocks.
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EFFECTS OF LATE-STAGE PROCESSES ON OXYGEN ISOTOPE COMPOSITIONS |
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Late-stage minerals from the syenites (amphibole II) and alkali granites (aegirine) have much lower
18O values than the primary magmatic minerals. Petrological criteria indicate that amphibole II in the syenites formed at temperatures below 550°C and aegirine at even lower temperatures, possibly <300°C (see above). Principally, two contrasting explanations for these low 18O values are possible: (1) the late-stage minerals formed in the presence of a low-18O fluid; (2) the low values reflect increasing fractionation at lower temperatures between minerals open to oxygen exchange.
To evaluate these two possibilities, changes in the 18O values of the minerals and the coexisting fluid phase during cooling were modelled for both rock types. For the calculation of the closed-system evolution in the syenites, the following assumptions were made:
- quartz, augite and amphibole I were in isotopic equilibrium at magmatic temperatures of 700–800°C. Using the fractionation factors of Bottinga & Javoy (1975) and Zheng (1993a), the calculated feldspar composition at the magmatic stage would have been between +6·5 and +7·5.
- The calculated oxygen isotope composition of coexisting water (which is used as an approximation for the water-dominated fluid phase in amphibole II-bearing syenites) at these temperatures ranges from +6·3 to +7·5 based on fractionation factors of Zheng (1993a, 1993b).
- During cooling and amphibole II formation, oxygen exchange was possible only between amphibole II, feldspar and the coexisting fluid phase. Quartz, augite and amphibole I closed to oxygen diffusion at temperatures below 500°C (Giletti et al., 1978; Giletti & Yund, 1984; Farver & Giletti, 1985; Farver, 1989).
- Fluid/rock ratios are small enough that the fluid does not represent an inexhaustible reservoir that would retain a constant oxygen isotopic composition, independent of temperature or minerals it is in equilibrium with.
- As a simplification, the oxygen contents of amphibole and feldspar are assumed to be equal.
Given the above assumptions, the 18O value of the system open to exchange at any temperature is defined by
 | (9) |
18Ofsp is given by
 | (10) |
18Oam and 18Ofsp. Figure 17 illustrates the results of these calculations. Calculated and measured 18O values for amphibole II are equal at temperatures of 400°C. This fits well with the petrological results. Hence, the low 18O values measured for amphibole II in syenites can be explained by a closed-system cooling model. Feldspar should have an oxygen isotope composition of +6·5 to +7·5 at a temperature of 400°C, which is in agreement with a published 18O value of about +6·5 for fresh feldspar from the Puklen complex (Sheppard, 1986).
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However, measured values of altered perthitic feldspar are much lower, between about +3 and +4
, indicating a late hydrothermal alteration in the presence of low-18O fluids. As the dominant mineral in the syenites is alkali feldspar, this mineral dominates the whole-rock oxygen isotopic composition. A rough estimate, assuming 15 vol. % quartz (with constant values of about +8), 15% mafic minerals (average of about +5) and 70 vol. % feldspar indicates that the inhomogeneous whole-rock 18O values measured for syenites (+4·8 to +6·9) must largely be due to variation of the oxygen isotope composition of the alkali feldspar. Calculated 18O values of the altered feldspar thus vary between about +2·8 and +5·0. This is in agreement with two measured values of altered feldspar (+3·5 and +3·9, Table 7). Whether the 18O value of a rock (mineral) increases or decreases during fluid–rock interaction is a function of temperature of alteration, the 18O value of the fluid, and of the fluid–rock ratio. At temperatures below 100°C it is possible that the whole-rock 18O value actually increases even if fluids with very low 18O values (-10 or lower) are involved, simply because the mineral–water O-isotope fractionation factors are very high at these temperatures. However, at higher temperatures this is not the case, as mineral–water fractionation factors decrease with rising temperatures. It is therefore suggested that the heterogeneity of whole-rock 18O values is mainly caused by alteration of feldspars by low-18O fluids, probably dominated by meteoric water at temperatures in excess of 100°C. Such alteration has been described from a number of plutonic rocks (e.g. Criss & Taylor, 1986; Lutz et al., 1988).
Applying a closed-system model similar to that described above to the alkali granites for possible feldspar–aegirine exchange, the low 18O values of aegirine cannot be explained. If a closed system is assumed, the calculated 18O values of aegirine are much higher than the measured values. Thus, an influx of low-18O fluids is necessary to explain the isotopic compositions of late-stage aegirine in alkali granites (Fig. 17b). The measured aegirines are in oxygen isotopic equilibrium with late-stage quartz veins at temperatures of 250°C, which is in good agreement with the petrological results for aegirine formation. The calculated fluid oxygen isotopic composition at these low temperatures is about –3, which is more than 7 lower than the late-stage, closed-system fluid modelled to be in equilibrium with the syenites. This low-18O fluid could be of meteoric origin. The inferred palaeolatitude for South Greenland during late Gardar times is 30–60°N (Piper, 1992). By comparison with the present-day distribution of isotopic compositions of meteoric waters (e.g. Rozanski et al., 1993), it is likely that values for the local meteoric waters were even lower than our estimate of -3. The chemical and isotopic composition of this meteoric fluid may well have been buffered by interaction with the country rocks. During this process it may have also picked up significant quantities of Nd or at least changed its Nd-isotopic composition significantly, resulting in the remarkably low Nd values for aegirines. Additionally, this meteoric fluid may also be responsible for the inferred low-temperature alteration of alkali feldspar in the syenites and alkali granites. Although not strictly required, it is also possible that such a fluid has contributed to the decrease of the 18O values of amphibole II, a possibility that cannot be excluded on the basis of the present data.
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SUMMARY AND CONCLUSIONS |
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In the alkaline to peralkaline rocks of the Puklen complex two phase assemblages can be distinguished: a primary magmatic and a secondary late- to post-magmatic assemblage. The primary magmatic assemblage in the syenites consists of alkfsp ± qtz + aug ± ol + Fe–Ti oxides and interstitial Na–Ca amphibole. Solvus thermometry of early ternary feldspar phenocrysts indicates minimum crystallization temperatures of
950–750°C. Oxygen fugacity during this stage was low and ranged from 0·8 to 2·3 log units below the FMQ buffer. Mineral- and whole-rock geochemical data suggest that fractional crystallization of feldspar, olivine, augite, magnetite–ilmenite and apatite from a silica-undersaturated parental magma led to an increase of silica activity from 0·7 in quartz-free samples towards unity in quartz-bearing ones. The magmatic assemblage in the granites (alkfsp + qtz + am + ilm) crystallized at temperatures >750°C and redox conditions around the FMQ buffer. A number of processes including fractional crystallization, assimilation of country rocks, post-magmatic alteration and sub-solidus re-equilibration played a significant role in the complex geochemical and isotopic evolution of the Puklen rocks. The O, Sr and Nd isotopic data presented here for the Puklen complex can be used to distinguish between magmatic and post-magmatic processes. This is particularly caused by the fact that the three isotope systems investigated here (O, Rb/Sr, Sm/Nd) are affected to variable degrees during late-stage metasomatism and alteration. Isotopic compositions of different minerals from the same sample indicate a multi-source genesis for the Puklen rocks. Whole-rock analyses in similar cases would be inadequate and lead to complex results and incorrect geological interpretations. Even mineral separates may be problematic if adequate conclusions on magma genesis and late-stage history are desired, as some of the processes mentioned above may overlap each other.
Oxygen isotope compositions indicate that the magmas parental to the syenitic melts of the Puklen complex are compatible with derivation from a mantle source. During ascent of the syenite magmas, variable degrees of contamination with upper-crustal material occurred, which is shown by the large range of Nd values in the syenites. In contrast, the oxygen isotopic compositions of the primary minerals and estimated magma compositions are homogeneous (Figs 9 and 16). This may be a consequence of mass balance considerations: as oxygen is a major constituent of silicate rocks, a large contrast in oxygen isotopic composition between primary melt and contaminant is needed to produce significant differences in oxygen isotopic composition. The similar oxygen isotopic composition and the very different Nd isotope characteristics of the granitic upper crust of the Gardar Province fit well with a model in which the Puklen syenites are explained by variable amounts of contamination of a mantle-derived melt by granitic upper crust.
Calculated 18Omelt values for the alkali granites are significantly lower than those of the syenites (Figs 9 and 18). These low 18O values are believed to be a source feature and not an effect of low-temperature oxygen diffusion. However, the Nd values for two separates of magmatic amphibole from the alkali granites fall well within the range of Nd values from the syenites. Based on the Nd isotopic data, a common source for the syenites and alkali granites may be possible, but the low oxygen isotopic composition of the granites compared with the syenites requires a different contaminant or source with low 18O composition, which was not involved in the genesis of the syenites.
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Late-stage fluids, which were retained by the syenites, caused the replacement of the primary assemblage by secondary silicates: augite was replaced by aegirine–augite, aenigmatite formed at the expense of ilmenite and, as a result of the release of Ca during this process, secondary autometasomatic minerals such as ferro-actinolite, carbonates, hydroandradite and titanite formed. The formation of these secondary Ca-minerals took place over a wide temperature range between about 700° and 300°C. During this cooling, oxygen fugacity rose in most samples to values around the FMQ buffer. It is remarkable that the latest minerals in the syenites are not of a typical peralkaline composition, but are Ca-rich like the primary assemblage. The restriction to the syenites and their absence in the granites implies either a short transport capacity of the fluid phase for these elements or that late-stage autometasomatism in syenites and granites worked differently. Oxygen and neodymium isotope data for these secondary Ca phases indicate that the syenites generally experienced closed-system behaviour during cooling only. A few syenitic samples (GM1611, GM1615 and GM1616) reflect fO2 conditions and temperatures similar to those for the granites. This may indicate that fluids from the granites influenced these samples, as they were collected close to the granite body.
The late-stage formation of aegirine at the expense of arfvedsonite in the granites is shown to be a low-temperature process, which took place at conditions around the HM buffer at temperatures possibly around 300°C. Thus, the granites became more oxidized than the syenites, which possibly implies a different fluid source for the granites. The large oxygen and neodymium isotopic differences between primary amphibole and late-stage aegirine in the granites indicate that the source of the aegirine-forming fluids was isotopically different from that which formed the primary amphiboles. A major influx of a second, meteoric fluid, with low oxygen and neodymium isotope composition, was probably responsible for this. The widespread alteration of feldspar in all rock types of the Puklen complex, which is reflected in the low 18O isotopic composition of the whole rocks, may correlate with this latest meteoric fluid circulation.
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APPENDIX: MINERAL ABBREVIATIONS USED IN FIGURES AND TEXT |
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ACKNOWLEDGEMENTS |
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LA–ICP-MS measurements were carried out at the Large Scale Geochemical Facility supported by the European Community–Access to Research Infrastructure action of the Improving Human Potential Programme, contract HPRI-CT-1999-00008 awarded to Professor B. J. Wood (University of Bristol), which is gratefully acknowledged. Bruce Paterson provided invaluable help during these measurements. M. Westphal is thanked for his help during microprobe measurements, Elmar Reitter for his careful help during Sr and Nd isotope measurements, Gaby Stoschek for her help with sample preparation for mass spectrometry and oxygen isotope analysis, and Jasmin Köhler for patient hand picking of mineral separates. Thomas Wenzel helped to improve an earlier version of this manuscript. Extremely thorough reviews and constructive comments by D. Baker, C. Harris, I. Parsons, R. Trumbull and M. Wilson (editor) improved the quality of this work substantially. Financial support for this work was provided by the Deutsche Forschungsgemeinschaft (grant Ma-2135/1-2).
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