Journal of Petrology Advance Access originally published online on March 14, 2006
Journal of Petrology 2006 47(7):1261-1315; doi:10.1093/petrology/egl008
| ||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||
Genesis of Ultramafic Lamprophyres and Carbonatites at Aillik Bay, Labrador: a Consequence of Incipient Lithospheric Thinning beneath the North Atlantic Craton
1 MAX-PLANCK-INSTITUT FÜR CHEMIE POSTFACH 3060, 55020 MAINZ, GERMANY
2 INSTITUT FÜR GEOWISSENSCHAFTEN, UNIVERSITÄT MAINZ BECHERWEG 21, 55099 MAINZ, GERMANY
3 DEPARTMENT OF EARTH SCIENCES, MEMORIAL UNIVERSITY ST. JOHN'S, NEWFOUNDLAND, CANADA A1B 3X5
4 DEPARTMENT OF EARTH AND ATMOSPHERIC SCIENCES, UNIVERSITY OF ALBERTA EDMONTON, ALBERTA, CANADA T6G 2E3
5 GEOLOGICAL SURVEY OF CANADA OTTAWA, ONTARIO, CANADA K1A 0E8
6 GEOFORSCHUNGSZENTRUM POTSDAM TELEGRAFENBERG, 14473 POTSDAM, GERMANY
7 GEOCHEMISCHES INSTITUT, UNIVERSITÄT GÖTTINGEN GOLDSCHMIDTSTRASSE 1, 37077 GÖTTINGEN, GERMANY
RECEIVED JULY 16, 2004; ACCEPTED FEBRUARY 2, 2006
 |
ABSTRACT |
|---|
TOP
ABSTRACT
INTRODUCTIONNumerous dykes of ultramafic lamprophyre (aillikite, mela-aillikite, damtjernite) and subordinate dolomite-bearing carbonatite with U–Pb perovskite emplacement ages of
590–555 Ma occur in the vicinity of Aillik Bay, coastal Labrador. The ultramafic lamprophyres principally consist of olivine and phlogopite phenocrysts in a carbonate- or clinopyroxene-dominated groundmass. Ti-rich primary garnet (kimzeyite and Ti-andradite) typically occurs at the aillikite type locality and is considered diagnostic for ultramafic lamprophyre–carbonatite suites. Titanian aluminous phlogopite and clinopyroxene, as well as comparatively Al-enriched but Cr–Mg-poor spinel (Cr-number < 0.85), are compositionally distinct from analogous minerals in kimberlites, orangeites and olivine lamproites, indicating different magma geneses. The Aillik Bay ultramafic lamprophyres and carbonatites have variable but overlapping 87Sr/86Sri ratios (0·70369–0·70662) and show a narrow range in initial 
Nd (+0·1 to +1·9) implying that they are related to a common type of parental magma with variable isotopic characteristics. Aillikite is closest to this primary magma composition in terms of MgO (15–20 wt %) and Ni (200–574 ppm) content; the abundant groundmass carbonate has 
13CPDB between –5·7 and –5
, similar to primary mantle-derived carbonates, and 18OSMOW from 9·4 to 11·6. Extensive melting of a garnet peridotite source region containing carbonate- and phlogopite-rich veins at 4–7 GPa triggered by enhanced lithospheric extension can account for the volatile-bearing, potassic, incompatible element enriched and MgO-rich nature of the proto-aillikite magma. It is argued that low-degree potassic silicate to carbonatitic melts from upwelling asthenosphere infiltrated the cold base of the stretched lithosphere and solidified as veins, thereby crystallizing calcite and phlogopite that were not in equilibrium with peridotite. Continued Late Neoproterozoic lithospheric thinning, with progressive upwelling of the asthenosphere beneath a developing rift branch in this part of the North Atlantic craton, caused further veining and successive remelting of veins plus volatile-fluxed melting of the host fertile garnet peridotite, giving rise to long-lasting hybrid ultramafic lamprophyre magma production in conjunction with the break-up of the Rodinia supercontinent. Proto-aillikite magma reached the surface only after coating the uppermost mantle conduits with glimmeritic material, which caused minor alkali loss. At intrusion level, carbonate separation from this aillikite magma resulted in fractionated dolomite-bearing carbonatites (13CPDB –3·7 to –2·7) and carbonate-poor mela-aillikite residues. Damtjernites may be explained by liquid exsolution from alkali-rich proto-aillikite magma batches that moved through previously reaction-lined conduits at uppermost mantle depths. KEY WORDS: liquid immiscibility; mantle-derived magmas; metasomatism, Sr–Nd isotopes; U–Pb geochronology
|
INTRODUCTION |
|---|
The ultramafic lamprophyres (UML; Rock, 1986
) are a widely recognized group of alkaline igneous rocks associated with continental extension; however, their origin is poorly understood. Although they are volumetrically minor components of continental magmatism, they are of fundamental significance for our understanding of deep melting events during the initial stages in the development of continental rifts. UML typically occur as dyke swarms and in central complexes (Rock, 1991), but their genesis has commonly been discussed as though they are varieties of kimberlite (Dawson, 1971; Griffin & Taylor, 1975; Raeside & Helmstaedt, 1982; Alibert & Albarède, 1988; Dalton & Presnall, 1998) mainly as a result of a similar macroscopic appearance and often problematic identification within existing classification schemes (Tappe et al., 2005a). However, compositional differences and the lack of spatial coexistence between contemporaneous UML and kimberlites (Rock, 1991; Mitchell, 1995) suggest that they are derived from distinct magma types. The occurrence of UML is largely confined to regions of lithospheric extension and they are commonly associated with carbonatite magmatism. Additionally, UML magmatism forms the earliest igneous activity in some flood basalt provinces (Queen et al., 1996; Leat et al., 2000; Riley et al., 2003) and also occurs on oceanic plateaux (Nixon et al., 1980; Neal & Davidson, 1989). In contrast, kimberlites occur exclusively within areas of stable Archean cratons or in the surrounding Proterozoic mobile belts (Mitchell, 1986; Janse & Sheahan, 1995). As for kimberlites, UML magmas may contain diamonds (Hamilton, 1992; Mitchell et al., 1999; Digonnet et al., 2000; Birkett et al., 2004), indicating that the depth of melting can be in excess of 150 km (>5 GPa), and, thus, may not be the crucial petrogenetic difference.
Large areas of Labrador, adjacent northeastern Quebec and western Greenland consist of Archean blocks surrounded by Paleoproterozoic mobile belts (Fig. 1) stabilized at 1900–1700 Ma (Wardle & Hall, 2002). Continental extension affected this cratonic area repeatedly during Mesoproterozoic (1350–1140 Ma; Romer et al., 1995; Upton et al., 2003), Neoproterozoic (620–550 Ma; Gower et al., 1986; Kamo et al., 1989; Larsen & Rex, 1992; Murthy et al., 1992) and Mesozoic times (Hansen, 1980; Keen et al., 1994; Chian et al., 1995; Larsen et al., 1999; Srivastava & Roest, 1999), eventually causing the break-up of the North Atlantic craton and opening of the Labrador Sea at 60 Ma (Chalmers & Laursen, 1995; Chalmers & Pulvertaft, 2001).
|
All of these extensional episodes were accompanied by volatile-rich alkaline igneous activity (Larsen & Rex, 1992
), but the most productive, in terms of UML magma generation, was the Late Neoproterozoic episode, which was distally associated with initiation of the Iapetus Ocean (Tappe et al., 2004). Neoproterozoic UML dykes related to this lithospheric stretching occur in the Sisimiut–Sarfartoq–Maniitsoq areas of western Greenland (Scott, 1981; Thy et al., 1987; Larsen & Rex, 1992; Mitchell et al., 1999; Heaman, 2005), the Torngat Mountains in northern Quebec and Labrador (Digonnet et al., 2000; Tappe et al., 2004), the Otish Mountains region in central–north Quebec (Heaman et al., 2004), single occurrences along the northern Labrador coast at Hebron, Saglek, Eclipse Harbour, Iselin Harbour and Killinek Island (Tappe et al., 2005b), and Aillik Bay in central-east Labrador (Malpas et al., 1986; Foley, 1989a). Here, we report the results of a petrological and geochemical study, combined with U–Pb perovskite and 40Ar/39Ar phlogopite age determinations, on a diverse suite of Neoproterozoic UML and associated carbonatites from the Aillik Bay area on the Labrador Sea coast. We discuss whether the large compositional diversity reflects mantle source heterogeneity or variability in the melting process, or relates to modification of a common parental UML magma by low-pressure processes, such as liquid immiscibility and devolatilization. Additionally, we specify and emphasize fundamental differences in the characteristics and genesis of UML and other ultramafic magma types, such as kimberlites. Late Neoproterozoic UML magma production occurred throughout the North Atlantic region attendant with widespread lithospheric stretching, thinning and the eventual break-up of the supercontinent Rodinia.
|
GEOLOGY OF THE AILLIK BAY AREA |
|---|
The southern North Atlantic craton margin
The Aillik Bay area is situated within the Paleoproterozoic Makkovik orogen at the southern edge of the Archean North Atlantic craton (NAC; Fig. 1). Reworked Archean orthogneisses (protolith ages 3260–2800 Ma) equivalent to the adjacent NAC are exposed along the western margin of the Makkovik Province (Fig. 2), whereas a juvenile high-grade magmatic arc crust dominates the central (Aillik Group) and eastern part close to the Grenville deformation front (Culshaw et al., 2000a
, 2000b; Sinclair et al., 2002). These supracrustal units formed during a sequence of subduction and accretion events between 1900 and 1700 Ma (Makkovikian Orogeny) and were later detached from the basement and thrust onto the edge of the NAC (Ketchum et al., 2002; Wardle & Hall, 2002). Seismic data and the unradiogenic initial Nd isotope composition of widespread post-orogenic granites (1720–1650 Ma; e.g. Strawberry granite in Fig. 2) clearly indicate that the western part of the Makkovik orogen, including the Aillik Bay area (Fig. 2), is underlain by the Archean crust of the NAC (Kerr & Fryer, 1994; Hall et al., 1995; Kerr & Wardle, 1997; Kerr et al., 1997).
|
Widespread lithospheric thinning occurred throughout eastern North America along the former Laurentian margin during the Late Neoproterozoic (Bond et al., 1984
; Kamo et al., 1995; Torsvik et al., 1996; Cawood et al., 2001; Puffer, 2002), resulting in continental break-up and subsequent opening of the Iapetus Ocean at 600 Ma. In central Labrador, this episode of continental stretching is recorded by remnant graben structures forming the eastward continuation of the prominent St. Lawrence valley rift system (Gower et al., 1986; Murthy et al., 1992).
Alkaline magmatism and previous age constraints
Late Neoproterozoic UML and carbonatite dykes occur in an area at least 30 km by 30 km around Aillik Bay (Fig. 2; Appendix A). These dykes are narrow (up to 3 m wide) and dominantly steeply dipping; subordinate flat-lying sheets also occur. Recognized UML types are aillikite, mela-aillikite and damtjernite, following the scheme devised by Tappe et al. (2005a). The subvertical dykes are roughly north–south-oriented and appear to converge towards a focus in the Labrador Sea (Fig. 2). Flow banding, back-veining and internal chill-bands are often seen, whereas fluidized globular ‘autolithic’ segregations are rare. Individual members of this dyke swarm cross-cut each other in a rather arbitrary manner. Some aillikite sheets or dykes grade laterally into carbonatite and/or mela-aillikite (Fig. 3). A weighted K–Ar mica age of 570 Ma, obtained for a poorly described ultramafic dyke rock from Aillik Bay (Leech et al., 1962), provided the only age constraint for the carbonate-rich magmatism when this study was initiated.
|
UML magmatism was preceded by Mesoproterozoic ultrapotassic magma production (
1374 Ma; Tappe et al., in preparation), represented by subvertical, 0·2–2 m wide, fine- to medium-grained olivine lamproite dykes within the same area. The youngest record of alkaline igneous activity around Aillik Bay is a Mesozoic suite of melilitite, nephelinite and basanite dykes (142 Ma; Tappe et al., in preparation), which appears to be related to the poorly exposed Ford's Bight alkaline intrusion (King & McMillan, 1975). |
GEOCHRONOLOGY |
|---|
U–Pb dating of ultramafic lamprophyres
Aillikite, mela-aillikite and damtjernite dykes were selected from all parts of the Aillik Bay area for U–Pb perovskite dating (Fig. 2). Analytical details can be found in Appendix B. Results are reported in Table 1 and displayed in concordia diagrams in Fig. 4. Two perovskite fractions of aillikite dyke ST123 from the east shore of Kaipokok Bay yielded similar 206Pb/238U dates of 560·7 ± 2·4 and 564·5 ± 3·0 Ma, respectively. Hence, a weighted average 206Pb/238U date of 562·2 ± 1·9 Ma is considered the best age estimate for emplacement of dyke ST123. A similar 206Pb/238U age of 569·2 ± 1·8 Ma was obtained from mela-aillikite dyke ST114A exposed on the west shore of Aillik Bay. The emplacement age of aillikite dyke ST228 from the southern shore of Makkovik Bay was determined to be 576·4 ± 6·5 Ma (weighted average 206Pb/238U date of two perovskite fractions 574·4 ± 1·8 Ma and 578·6 ± 2·0 Ma). The emplacement age of aillikite dyke ST220II from West Turnavik Island is 589·6 ± 1·3 Ma (weighted average of 589·4 ± 1·4 and 590·5 ± 2·8 Ma).
|
|
The youngest perovskite ages obtained from damtjernite dykes are 555·0 ± 1·8 Ma (ST211A; Main Turnavik Island) and 563·9 ± 2·5 Ma (ST256; east shore of Makkovik Bay). Perovskites from damtjernite dyke ST174 (Pigeon Island) yielded a 206Pb/238U date of 574·6 ± 1·6 Ma. Strikingly similar weighted average ages of 581·9 ± 2·3 Ma (582·5 ± 4·8 and 581·9 ± 2·6 Ma) and 582·5 ± 2·1 Ma (582·8 ± 3 and 582·1 ± 2·2 Ma) were obtained from the damtjernites ST140A (east shore of Aillik Bay) and ST188A (Red Island), respectively.
Taken together, the high-precision 206Pb/238U perovskite dates for four individual aillikite/mela-aillikite dykes and five damtjernite dykes cover a similar age range between 562–590 Ma and 555–583 Ma, respectively (Fig. 4). Hence, aillikite and damtjernite magmatism can be considered coeval over 30–35 Myr during Late Neoproterozoic extension at a craton margin. Although no carbonatites have been dated, their close association with the various dated UML types implies contemporaneous emplacement.
40Ar/39Ar dating of cognate inclusion ST162I
A clinopyroxene–phlogopite inclusion recovered from aillikite dyke ST162 on the west shore of Aillik Bay (Fig. 2) yielded a phlogopite 40Ar/39Ar plateau age of 573·3 ± 3·3 Ma (Table 2), which falls within the U–Pb perovskite age range of the four dated aillikite dykes.
|
The plateau was calculated over 10 consecutive steps, which contained 98% of the total released 39Ar. The gas release spectrum is shown in Fig. B1 (Appendix B). The 3·5%, 3·9% and 4·2% laser power steps significantly overlap the plateau but have slightly older apparent ages, indicating a potential presence of excess argon. However, on an inverse isochron diagram, a regression through these three data points and the seven others included in the plateau age calculation passes through a 40Ar/36Ar value of 433·1 ± 257·7, which is within error of the atmospheric value (295·5). As the three apparently older steps have no significant effect on the overall age, they have been included in the plateau age calculation.
|
PETROGRAPHY |
|---|
Rock types of the Neoproterozoic Aillik Bay UML suite include carbonatite, aillikite, mela-aillikite and damtjernite, listed in order of decreasing carbonate content. Additionally, aillikite dykes host a wide variety of micaceous cognate inclusions (Fig. 5). Modal mineral abundances are listed in Table 3 and mineral compositional data are given in Tables 4–11 (extended tables can be downloaded from the Journal of Petrology website at http://www.petrology.oupjournals.org as Electronic Appendix 1). It should be noted that descriptions of the following accessory minerals are solely provided as Electronic Appendix 2: ilmenite, rutile, perovskite, titanite, apatite, alkali feldspar, feldspathoids, pectolite.
|
|
|
|
|
|
|
|
|
|
Carbonatite dykes
Two distinct types of carbonatite can be distinguished: (1) a dolomite carbonatite devoid of any mafic silicates; (2) a mixed dolomite–calcite carbonatite containing minor amounts of clinopyroxene, phlogopite and olivine crystals. The dolomite carbonatite mainly consists of a mosaic of equigranular Fe-rich dolomite crystals (100–300 µm). Hydroxy-fluorapatite forms abundant euhedral microphenocrysts (50–150 µm). Interstices may be filled by barite, quartz, alkali feldspar (orthoclase and albite) and/or tiny rare earth element (REE)-carbonate crystals. Large rutile grains commonly occur (50–100 µm), whereas opaque phases including magnetite are comparatively rare.
The dolomite–calcite carbonatites exhibit a granular to interlocking texture dominated by calcite grains and laths (150–300 µm). Calcite coexists with subordinate laths of Fe-rich dolomite (Fig. 6a). Zoned phlogopite plates (up to 0·5 mm) and olivine grains (up to 1·0 mm; replaced by carbonates) are observed, suggesting gradation into aillikites. However, the presence of diopside-rich clinopyroxene phenocrysts (up to 1·2 mm) contrasts with the aillikites. Fresh rutile grains and apatite prisms (up to 0·4 and 1·0 mm, respectively) are abundant, as are opaque oxides. Secondary interstitial barite and/or fluorite may occur.
|
Aillikite dykes
Aillikites are texturally heterogeneous (e.g. Fig. 5a and b); some exhibit an inequigranular texture with olivine and phlogopite macrocrysts up to 7 mm in diameter, whereas the majority are weakly inequigranular and have a porphyritic texture (Fig. 5b). Porphyritic aillikites are characterized by phenocrysts of euhedral to subhedral olivine (0·6–1·3 mm), phlogopite (0·25–0·5 mm), apatite and magnetite (0·2–0·4 mm) in a carbonate matrix. Mela-aillikites are distinguished from aillikites in containing more mafic silicate phases (>70 vol. %) and less carbonate (<10 vol. %). The change in the modal mineral proportions is gradational from aillikite to mela-aillikite. End-member mela-aillikite contains abundant clinopyroxene prisms in the groundmass (100–300 µm; Fig. 5c), which are rare in aillikite. Both rock types carry microphenocrysts of olivine (0·25–0·5 mm), phlogopite (<0·25 mm), apatite, opaque oxides (dominantly titanomagnetite and Mg-rich ilmenite) and perovskite or rutile (50–200 µm). Primary kimzeyitic garnet typically occurs (<100 µm).
A rare textural variety occurs locally in otherwise uniformly textured aillikite dykes and consists of globular aillikite segregations ‘cemented’ by primary calcite laths. The globular segregations are made up of olivine and/or glimmerite kernels surrounded by concentrically arranged fine-grained aillikite matrix (Fig. 5a). Their strong resemblance to nucleated autoliths suggests an origin by fluidization of partly solidified early magma fractions as a result of local near-surface devolatilization.
Damtjernite dykes
Damtjernites are medium- to fine-grained, porphyritic to intergranular rocks (Fig. 5d) containing rare macrocrysts (up to 2·0 cm) of virtually Cr-free diopside-rich clinopyroxene and/or Cr-free titanian aluminous phlogopite. Modal layering, internal chill zones, bounded felsic segregations, flow-alignment and rotation structures are common macroscopic features of these rocks, which were called ‘sannaites’ by Foley (1984) and Malpas et al. (1986), but have been renamed here following Tappe et al. (2005a).
The phenocryst assemblage of the damtjernites consists of olivine (up to 1 mm), phlogopite (up to 5 mm), rare clinopyroxene (250–800 µm) and apatite (up to 1 mm). The modal abundance of euhedral to subhedral olivine phenocrysts may vary even within dykes from 20 vol. % to only a few crystals. Phlogopite forms large plates typically enclosing clinopyroxene and apatite needles. The groundmass mainly consists of a mesh made up of clinopyroxene and apatite needles (up to 200 µm) and biotite flakes resembling the late rims on phlogopite. Sr-calcite, alkali feldspar (almost pure orthoclase and albite) and nepheline occur in variable but small modal proportions interstitial to the mica and clinopyroxene of the groundmass (Fig. 5d). Additional rare felsic phases are analcime and sodalite. Pectolite is observed as a fibrous replacement product of groundmass clinopyroxene or as primary crystals in interstices. Olivine is absent in the groundmass. Abundant titanomagnetite grains, ilmenite laths, and rutile or perovskite crystals are the principal groundmass oxide phases. Perovskite relicts may be enclosed by Zr-rich titanite crystals recording fluctuations in silica activity during magma evolution or slow cooling. Schorlomite and/or melanite garnet occurs rarely in the groundmass in association with perovskite (Fig. 6e). Felsic segregations (orthoclase, albite, nepheline, analcime, sodalite, calcite, Mg-ankerite) with fairly sharp contacts with the groundmass are a characteristic feature of the damtjernites (Foley, 1984).
Cognate inclusions
A suite of undeformed micaceous inclusions, exclusively hosted by aillikites, comprises (1) glimmerite, (2) clinopyroxene–phlogopite and (3) olivine–phlogopite nodules in order of decreasing abundance (Table 3).
Glimmerite nodules are typically oval and less than 2 cm in diameter (Fig. 5e); rare examples approach 5 cm. Most glimmerites consist of interlocking 20–100 µm phlogopite flakes (Fig. 5f); some contain larger isolated phlogopite ‘clasts’ (up to 500 µm). Fluorapatite and rare orthoclase fill interstices or form discontinuous bands (Fig. 5e and f) that may open into radiating patches. Tiny spinel and ilmenite grains (<100 µm; also composite) are scattered throughout the fine-grained matrix.
Clinopyroxene–phlogopite nodules (up to 8 cm across; Fig. 5e) are similar in shape to the glimmerites, but have more variable mineralogy and are coarser grained. Large poikilitic phlogopite plates and clinopyroxene prisms dominate (up to 2 mm), whereas lath-like to interstitial Mg-ilmenite, chromite–titanomagnetite and prismatic hydroxy-fluorapatite occur only as minor components (Fig. 5 g). Mica plates occasionally enclose carbonated subhedral olivine. A subtle grain-size layering was observed in larger nodules. Interstitial calcic amphibole occasionally replaces clinopyroxene and phlogopite (Figs 5g and 6f), presumably as a product of melt infiltration. Subhedral titanite occurs in a single nodule (ST250C).
Rare cumulate-textured olivine–phlogopite nodules contain irregular olivine grains (300–800 µm), which are typically enclosed by large phlogopite plates (0·5–1·0 mm; Fig. 5 h). Hydroxy-fluorapatite (200–400 µm), titanomagnetite, ilmenite (200–800 µm) and rare clinopyroxene (<300 µm) occur as intercumulus phases. Perovskite and zirconolite (<100 µm) are rare accessories associated with titanomagnetite surrounding olivine grains.
|
MINERAL COMPOSITIONS |
|---|
Olivine
Aillikite/mela-aillikite olivine phenocrysts or microphenocrysts exhibit a fairly large range in forsterite component (Fo91–80 mol %; Figs 6b and 7, and Table 4), NiO (0·5–0·05 wt %), CaO and MnO (maxima of 0·9 and 0·4 wt %, respectively). Contrasting Mg/Fe evolutionary trends may indicate that different olivine populations are present. Olivine phenocrysts with normal zoning have core compositions of Fo87–91, decreasing to Fo82–85 towards the rim; NiO decreases, whereas CaO and MnO typically increase (0·1–0·3 wt %). Repetition of a normal zoning pattern may occur (Fig. 6b). Reverse zoning was often observed, with core compositions of Fo82–84 (NiO 0·2–0·3 wt %) steadily increasing towards the rim (Fo87–88; NiO 0·4 wt %). A discrete Fe-enriched overgrowth (Fo82; NiO 0·1 wt %) with sharp contact to the inner phenocryst usually occurs on reverse-zoned crystals, and can be strongly CaO enriched (up to 0·9 wt %).
|
Olivine phenocrysts in damtjernites are normal zoned (Fo80–86·5) and contain 0·18–0·5 wt % NiO (Fig. 7). CaO and MnO approach 0·4 wt % at the rims. The most primitive olivine cores in damtjernites are more evolved than their most primitive counterparts from aillikite/mela-aillikite (Fo91) but have similar high NiO concentrations (Fig. 7).
Rare subhedral olivine crystals in clinopyroxene–phlogopite nodules are normally zoned with a fairly evolved composition (Fo77–86·6; <0·1 wt % NiO; <0·4 wt % CaO; Fig. 7) and a conspicuously high MnO content (0·3–0·7 wt %). The olivine compositions (Fo80–86) in olivine–phlogopite nodules overlap with the low Mg/Fe end of aillikite phenocrysts, but are richer in NiO (up to 0·4 wt %) at a given Fo content (Fig. 7).
In general, olivine phenocryst compositions in UML and associated cognate inclusions from the Aillik Bay area are less primitive (<Fo91) than those found in kimberlites and lamproites, which typically approach Fo93 (Mitchell, 1986; Mitchell & Bergman, 1991; Fedortchouk & Canil, 2004; Prelevic et al., 2005).
Phlogopite
Phlogopite phenocrysts from aillikite and dolomite–calcite carbonatite typically have (1) a resorbed core with 15–16 wt % Al2O3 and up to 5 wt % TiO2, (2) a broad inner rim with elevated Al2O3 (up to 18 wt %) and lower TiO2 (about 2 wt %) and (3) a narrow outer rim with Al2O3 and TiO2 falling below 10 and 1 wt %, respectively, at constantly high MgO. This trend may culminate in virtually Al- and Ti-free tetraferriphlogopite rims in the most carbonate-rich samples (Figs 6c and 8a; Table 5). Whereas tetraferriphlogopite is uncommon in kimberlites, the rims reported here are similar to those from orangeites but distinct from titanian tetraferriphlogopites in lamproites. Phlogopite from type aillikite/mela-aillikite and carbonatite is generally Ba poor with core compositions typically below 1 wt % BaO. Rim compositions only rarely approach 3 wt % BaO. This is in contrast to kimberlite phlogopites, which commonly show a strong Al and Ba enrichment and Ti depletion toward the rim (Fig. 8a). The inner zones of less extremely zoned phlogopite plates (0·5–1·0 mm) from mela-aillikite contain 13–15 wt % Al2O3 and 3–5 wt % TiO2, but a high-Al inner rim composition such as in aillikite is absent. As in aillikites, Al2O3 depletion (8–13 wt %) toward the rim is common, but TiO2 increases rimwards (up to 6 wt %; Fig. 8a), which contrasts with the decreasing TiO2 trend observed in aillikites. Furthermore, micas from mela-aillikite follow a different Mg/Fe evolutionary trend than aillikite micas, with a strong increase in Fe at the expense of Mg leading to discrete dark brown biotite rims.
|
Damtjernite phlogopite plates compositionally resemble the less extremely zoned Ba-poor phlogopite plates from mela-aillikites with inner zones containing 13–15 wt % Al2O3 and 3–5 wt % TiO2 (Fig. 8b). A few samples were found to contain micas approaching 8 wt % TiO2 in the core. In general, these micas lack the high-Al inner rim composition of aillikite micas, but show Al depletion toward the rim (8–13 wt % Al2O3; Fig. 8b). We noted both rimward TiO2 increase (up to 8·5 wt % as in mela-aillikite) and decrease (down to 1 wt % as in aillikite). Micas in damtjernite show a strong Fe increase at the expense of Mg. This culminates in broad dark brown biotite overgrowths (Fig. 5d).
The mica compositional range in the clinopyroxene–phlogopite and olivine–phlogopite nodules (Fig. 8c) is the same as in phenocrysts from aillikites/mela-aillikites and damtjernites with the characteristically high Al2O3 (13–15 wt %) and TiO2 (1–8 wt %), but low BaO concentrations (<1·0 wt %). Fluorine concentrations are as low as in UML micas (<1·3 wt %), but much lower than in glimmerite phlogopite. The phlogopite plates are distinct from primitive mica compositions reported for MARID nodules (typically <12 wt % Al2O3, Dawson & Smith, 1977; Smith et al., 1978) and do not show evolution toward either tetraferriphlogopite or biotite.
Glimmerite phlogopites are compositionally unlike any of the phlogopite phenocrysts, plates or groundmass flakes described above (Table 5). They are highly magnesian (Mg-number 70–90), Al2O3 and TiO2 poor (5–12 and 0·3–2·0 wt %, respectively; Fig. 8c), BaO depleted (<0·2 wt %) but enriched in F (1–3 wt %).
Clinopyroxene and amphibole
Phenocrystic clinopyroxene in dolomite–calcite carbonatite and damtjernite, as well as groundmass prisms in damtjernites and mela-aillikites, are diopside-rich, showing Al2O3 and TiO2 enrichment towards the rim (up to 10 and 6 wt %, respectively). However, the average atomic Al/Ti ratio of carbonatite clinopyroxene is 3, distinctively higher than in clinopyroxene from associated mela-aillikite and damtjernite (2; Fig. 9a; Table 6). Cr2O3 concentrations in all these diopsides are below 0·1 wt %. Aegirine-rich overgrowths (up to 46 mol %) occur around diopside-rich phenocrysts in damtjernite. Some of these phenocrysts contain rare resorbed green Fe-rich salitic clinopyroxene cores that are rich in Al2O3 (up to 9 wt %).
|
Clinopyroxene–phlogopite nodules also contain diopside-rich clinopyroxene enriched in Al2O3 and TiO2 (up to 8 and 4 wt %, respectively), with atomic Al/Ti of
2·5 similar to the UML clinopyroxenes (Fig. 9b). Slightly FeO- and Na2O-enriched salitic core compositions (up to 9 and 3 wt %, respectively) may occur with Cr2O3 below 0·3 wt %. Olivine–phlogopite nodules carry rare diopside (4 and 2 wt % Al2O3 and TiO2, respectively), which is the most Cr2O3-rich composition (0·1–0·6 wt %) of all the clinopyroxenes from the Aillik Bay UML suite. In general, the strong Al and Ti enrichment in clinopyroxene from the Aillik Bay UML and their cognate inclusions contrasts with diopsidic compositions typical for groundmass clinopyroxene in orangeites, lamproites and associated MARID-type inclusions (Mitchell & Bergman, 1991; Mitchell, 1995). The intercumulus calcic amphibole found in the clinopyroxene–phlogopite nodules (Fig. 6f) is generally MgO and TiO2 rich (Mg-number 73–90 and 1·9–5·0 wt % TiO2) and ranges from magnesiohastingsite through pargasite to rare magnesiokatophorite. Fluorine is <0·5 wt % and K2O does not exceed 1·9 wt %.
Spinel group
Spinel group minerals from Aillik Bay UMLs and related micaceous inclusions generally follow a ‘titanomagnetite trend’ (trend 2 of Mitchell, 1986) which is characterized by FeT2+/(FeT2+ + Mg) >0·7, increasing Fe and Ti but decreasing Mg, Al and Cr (Fig. 10). The most Mg-rich spinels were found in aillikites but do not exceed 13·5 wt % MgO (Table 7). By comparison, kimberlite spinels are more magnesian (12–20 wt % MgO) and follow a trend of increasing Ti at buffered Fe/Mg of 0·5 (Fig. 10b; trend 1 of Mitchell, 1986). Aillik Bay UML spinels have Cr/(Cr + Al) ratios <0·85, which is in marked distinction to Cr-rich spinels in lamproites and orangeites (Cr-number >0·85), which follow the ‘titanomagnetite trend’ (Mitchell & Bergman, 1991; Mitchell, 1995).
|
Early-stage spinels in aillikites are typically composed of chromite–spinel solid solutions (up to 43 wt % Cr2O3, 13 wt % MgO, 12 wt % Al2O3). Rims of zoned spinel microphenocrysts and individual grains are of ulvöspinel–magnesian ulvöspinel–magnetite composition (up to 11 wt % MgO). The titanomagnetites are enriched in Al2O3 (up to 11 wt %) with low atomic Cr/(Cr + Al) ratios (<0·3). Spinels from mela-aillikites may contain cores of titanian magnesiochromite–chromite solid solution (up to 12 wt % TiO2, 9 wt % MgO, 25 wt % Cr2O3) and of chromite–spinel solid solution, similar to their aillikite analogues. Individual titanomagnetite microphenocrysts or rims around zoned chromite grains contain less MgO and Al2O3 (<5 wt %) than in the aillikites (Fig. 10a).
Spinels in damtjernites are dominantly titanomagnetite, which rarely exhibits cores of titanian magnesiochromite–chromite solid solution (up to 14 wt % TiO2, 9 wt % MgO, 22 wt % Cr2O3). Titanomagnetite in damtjernites has the lowest MgO concentration (typically <1 wt %) of all the Aillik Bay UML spinels (Fig. 10a), and contains similar levels of Al2O3 to mela-aillikites, but significantly less Al2O3 than the aillikites. Individual grains have rims approaching magnetite end-member composition.
Rare spinels in glimmerites are similar to the most evolved aillikite spinels with MgO and Al2O3 typically below 5 wt % following a ‘titanomagnetite trend’ (Fig. 10b). Composite spinel grains in clinopyroxene–phlogopite nodules may contain cores of chromite (up to 43 wt % Cr2O3) and/or Cr-spinel (up to 20 wt % Cr2O3) typically mantled by titanomagnetite. Titanomagnetite–magnetite resembles late-stage spinels from pyroxene-rich mela-aillikites in being very close to magnetite end-member composition. They are much more depleted in MgO and Al2O3 (typically <2·0 and 3·0 wt %, respectively) than their analogues from pyroxene-free glimmerites and aillikites (Fig. 10b). Titanomagnetite in olivine–phlogopite nodules resembles evolved aillikite spinels. They contain more MgO than spinels from clinopyroxene-rich nodules and mela-aillikite dykes (up to 7 wt %; Fig. 10b). Cr-spinel with up to 25 wt % Cr2O3 rarely occurs as inclusion in olivine.
Carbonate and Ti-rich primary garnet
Fe-rich dolomite crystals in dolomite carbonatite contain between 2 and 9 wt % FeO; only rarely approaching 12 wt % towards the rim. MnO is elevated (0·2–1·2 wt %), whereas SrO and BaO are conspicuously low (<0·2 wt %). Rare interstitial REE-carbonate is probably bastnäsite and contains up to 56 wt % LREE2O3 (where LREE is light rare earth element). Calcite in mixed dolomite–calcite carbonatite coexists with subordinate laths of Fe-rich dolomite (Fig. 6a) which resembles its counterpart from the dolomite carbonatites (2–12 wt % FeO and 0·2–1·2 wt % MnO). It differs in that both calcite and dolomite contain up to 1·5 wt % SrO. The groundmass of aillikites is dominated by a mosaic of Sr-calcite (up to 2 wt % SrO), whereas dolomite containing up to 10 wt % FeO is rare. Fe-rich dolomite seems to dominate over calcite (<1 wt % SrO) in the generally carbonate-poor groundmass of mela-aillikites. Carbonate in the damtjernite groundmass is Sr-calcite with up to 4·8 wt % SrO.
Small kimzeyitic garnets (Zr-rich andradite) are restricted to aillikites and have a fairly constant TiO2 content (9–11 wt %), whereas ZrO2 spans a wide range between 10 and 17 wt % (Table 8). Core compositions are generally richer in Zr than the rims. Schorlomite and/or melanite garnet is rare but characteristic for damtjernites, and observed zoning patterns are typically from Ti-rich core compositions to more Fe-rich rims (1·8–18 wt % TiO2; 15·7–21·6 wt % FeO). Zirconian schorlomite with up to 5 wt % ZrO2 in the core was rarely found. The presence of Ti-rich andradites and kimzeyitic garnets reflects the high Ca and Ti but low Al concentration of the UML magma and can therefore be regarded as characteristic for UML–carbonatite associations (Platt & Mitchell, 1979; Rock, 1986; Tappe et al., 2005a). These garnets do not occur in kimberlites and lamproites (Mitchell & Bergman, 1991; Mitchell, 1995).
|
PRESSURE ESTIMATES FOR COGNATE INCLUSIONS |
|---|
The clinopyroxenes and rare calcic amphibole of the clinopyroxene–phlogopite nodules allow qualitative pressure estimates. The clinopyroxene barometer of Nimis & Ulmer (1998)
requires an independent temperature estimate, which we obtained using the clinopyroxene thermometer of Kretz (1982). The uncertainty in temperature is 60 °C (1
) and results in large errors in pressure estimates (0·3 GPa, 1). Nevertheless, the crystallization pressure of clinopyroxenes from several clinopyroxene–phlogopite nodules can be bracketed between 0·8 and 1·5 GPa, corresponding to 25–45 km depth. Rare clinopyroxene from an olivine–phlogopite nodule gives a similar pressure estimate of 0·9–1·7 GPa.
Calcic amphibole in clinopyroxene–phlogopite nodules yielded the lowest crystallization pressures of 0·4–0·7 GPa (Al-in-hornblende barometer of Hammarstrom & Zen, 1986), corresponding to 10–20 km depth. This agrees with textural relations indicating late melt/fluid infiltration into the nodule material (Fig. 6f). No pressure estimate can be given for the glimmerite nodules, but the low-Ba mica compositions may be a reflection of comparably high crystallization pressures (Guo & Green, 1990).
|
MINERALOGICAL CONSTRAINTS ON CRYSTALLIZATION CONDITIONS AND THEIR IMPLICATIONS FOR MANTLE SOURCE CHARACTERISTICS |
|---|
Oxygen fugacity estimates from olivine–spinel and ilmenite–magnetite pairs
Olivine and Cr-spinel are the earliest phases crystallized from aillikite magma and may be used to constrain the oxygen fugacity conditions during early stages in UML magma evolution. We applied the FeMg–1 exchange thermometer of O'Neill & Wall (1987)
and the olivine–orthopyroxene–spinel oxybarometer of Ballhaus et al. (1991). Ferric iron in spinel was calculated assuming stoichiometry (Ballhaus et al., 1990). Because ultramafic lamprophyres are not saturated in orthopyroxene, the oxybarometer of Ballhaus et al. (1991) yields maximum fO2 values, which can be corrected for the appropriate silica activity of the melt as outlined by Fedortchouk & Canil (2004). The perovskite–titanite reaction (Nicholls et al., 1971) rather than the monticellite–diopside reaction as chosen by Fedortchouk & Canil (2004) for kimberlite was considered as the upper limit of silica activity controlling UML magma evolution at Aillik Bay. This assumption is consistent with the observation that perovskite and diopside-rich clinopyroxene frequently occur in the groundmass of these rocks. Some damtjernites contain perovskite and titanite in reaction relationship, indicating that crystallization occurred along this silica activity buffer. We assumed an equilibration pressure for olivine–spinel pairs of 1 GPa: pressure has only a minor influence on the calculation of the equilibrium olivine–spinel crystallization temperature (20 °C/GPa) and oxygen fugacity (0·03 log-bar units/GPa). The olivine–spinel equilibration temperatures for the aillikite magma range from 912 °C to 1300 °C (Fig. 11). The oxygen fugacity varies from FMQ –0·03 to FMQ +2·43 (log-bar unit deviation from fayalite–magnetite–quartz buffer) with most pairs recording fO2 slightly above the FMQ buffer. An olivine–spinel pair from a damtjernite (1253 °C; FMQ +1·84), and from an olivine–phlogopite cognate inclusion (1002 °C; FMQ +1·83) fall within the fO2–T range calculated for aillikites.
|
The oxygen fugacity during UML groundmass crystallization was estimated from late ilmenite–magnetite pairs (0·2 GPa; QUILF-95 program; Andersen et al., 1993
). Results scatter around the FMQ buffer (FMQ –1·81 to FMQ –0·03 for aillikite; FMQ –1·98 to FMQ +1·89 for cognate inclusions; Fig. 11). Lower equilibration temperatures calculated for these ilmenite–magnetite pairs than for olivine–spinel phenocrysts are in keeping with the distinct crystallization stages. Ilmenite–magnetite temperatures for aillikites and micaceous cognate inclusions range from 868 °C to 641 °C, and 849 °C to 556 °C, respectively. Temperatures for damtjernite are at the lower end of this range (662–525 °C) and the few fO2 values for damtjernites show the greatest negative deviation from FMQ (FMQ –3·35 to FMQ +1·37) presumably reflecting sub-solidus re-equilibration (Fig. 11).
The oxygen fugacity values for UML from Aillik Bay are significantly higher than those for diamondiferous Slave craton kimberlites (Fig. 11; <FMQ –2·0; Fedortchouk & Canil, 2004), which lie close to the D/GCO (diamond/graphite–CO) buffer. Highly reduced crystallization conditions were also calculated for kimberlites from the Kaapvaal craton (Mitchell, 1973). Because the redox state of a primitive mafic magma has the potential to preserve that of its source (Carmichael, 1991), it can be inferred that the UML magma was derived from a fairly oxidized mantle region beneath the stretched North Atlantic craton in contrast to comparatively reduced sources for kimberlites within a stable cratonic mantle.
Hydrogen fluoride fugacity estimates from phlogopite-apatite pairs
Estimates of crystallization temperature and relative hydrogen fluoride (HF) fugacity in the UML magma were obtained from coexisting apatite and phlogopite crystals (Fig. 12). Equilibrium is assumed based on textures and mode of occurrence, e.g. apatite inclusions in phlogopite (Fig. 5f and g), and equilibrium pressure was set at a minimum of 0·2 GPa.
|
The biotite–apatite geothermometer recalibrated by Zhu & Sverjensky (1992)
yields rough temperature estimates in the range 1250–500 °C for aillikites and damtjernites from the Aillik Bay area (mean error = ±50 °C at 1). The micaceous cognate inclusions from aillikites and macrocrysts from damtjernites fall within this temperature range. High temperatures of the UML phlogopites at low crystallization pressures are also indicated by their high TiO2 concentrations (Forbes & Flower, 1974; Robert, 1976; Righter & Carmichael, 1996). Apatite inclusions in olivine phenocrysts indicate early crystallization of apatite.
Estimates of the relative HF fugacity, log(fHF/fH2O), based on Andersen & Austrheim (1991), using the apatite–phlogopite equilibrium temperature yield values between –4 and –7 ± 0·15 for Aillik Bay UML and associated cognate inclusions (Fig. 12). The correlation between HF fugacity and temperature displays a relatively shallow (slow) cooling trend internally buffered by phlogopite, consistent with mica being the dominant phase.
There is considerable overlap between relative HF fugacities recorded by Aillik Bay aillikite and damtjernite phenocrysts, and cognate clinopyroxene–phlogopite and olivine–phlogopite inclusions also fall within this range. Interestingly, the glimmerite inclusions record higher relative HF fugacities than their aillikite hosts and the remainder of the cognate inclusions at a given equilibration temperature (Fig. 12).
In summary, apatite and phlogopite equilibrated throughout the crystallization sequence of the UML magma (near liquidus to near solidus). Aillik Bay aillikites, damtjernites and cognate clinopyroxene–phlogopite/olivine–phlogopite inclusions experienced a similar evolution in terms of volatile fugacities. However, glimmerite nodules crystallized under higher relative HF fugacity conditions over a similar wide temperature range. The estimated relative HF fugacity of the Aillik Bay area UML is considerably lower than reported for carbonatites (Fig. 12; Fen complex; Andersen & Austrheim, 1991), but compares well with carbonate-rich UML from the Torngat Mountains (own data) and the Delitzsch complex, Germany (Seifert et al., 2000). The generally F-poor, OH-rich nature of phlogopites in Aillik Bay UML (<1 wt % F) may result from highly oxidizing crystallization conditions (Foley, 1989b), which is in marked contrast to the high F content in lamproitic micas (1–7 wt %; Mitchell & Bergman, 1991), which were experimentally shown to be derived from F-rich reduced mantle sources (Foley et al., 1986). Oxidizing crystallization conditions might also be responsible for the elevated Al content of the UML micas (atomic K/Al <1 vs >1 in lamproite micas at similarly strong bulk-rock Al depletion). The pronounced replacement of Mg by Ti leading to octahedral site vacancies is also in keeping with a redox control (Arima & Edgar, 1981; Foley, 1989b).
|
GEOCHEMISTRY AND ISOTOPIC COMPOSITION |
|---|
Major and compatible trace elements
Extreme SiO2 undersaturation, Al depletion, strong Ca enrichment and a potassic character is the hallmark of all members of the Aillik Bay UML suite (Fig. 13; Table 12 and Electronic Appendix 3). The aillikites contain 17–29·4 wt % SiO2 and 15–20·4 wt % MgO (Fig. 14). Mg-number ranges between 60 and 77, and Al2O3 concentrations are below 3·5 wt %. CaO is high but variable (13–24·7 wt %) and TiO2 is elevated (2·5–3·8 wt %) compared with kimberlites (Fig. 13). Moderately high K2O concentrations (1·3–2·4 wt %) but extreme Na2O depletion (<0·7 wt %) are characteristic for aillikites. P2O5 (0·9–3·2) and CO2 (10·1–20·8 wt %) concentrations are higher in aillikites than in mela-aillikites (P2O5 = 0·2–1·3 wt %; CO2 = 2–9·9 wt %; Figs 13 and 15). The latter have elevated SiO2 (30·7–36 wt %), TiO2 (3·7–5·8 wt %), Al2O3 (3·7–4·8 wt %) and K2O (up to 3·1 wt %) but much lower CaO (7·4–12·4 wt %) concentrations. Cr and Ni contents are high in aillikites (210–574 ppm Ni, 322–857 ppm Cr) and even higher in mela-aillikites (572–787 ppm Ni, 705-1150 ppm Cr; Fig. 14).
|
|
|
Damtjernites have higher SiO2 (29·3–38 wt %), TiO2 (3·5–6·9 wt %), Al2O3 (3·6–9·9 wt %) and Na2O (0·2–4·3 wt %) than aillikites (Fig. 13), reflecting less primary carbonate and the occurrence of clinopyroxene plus a felsic mineral in the groundmass. MgO varies considerably between 5·5 and 16 wt % (Fig. 14), which translates into a wide Mg-number range (40–67). Damtjernites have Ni and Cr concentrations that range from values typical for mantle-derived primitive magmas to low values close to the detection limit (22–517 ppm Ni; <10–586 ppm Cr; Fig. 14). There is some overlap with the major element composition of mela-aillikites (which lack the felsic groundmass component), but evolved damtjernites approach significantly higher levels of SiO2, Al2O3 and Na2O (Fig. 13). K2O and P2O5 concentrations can be high (0·9–3·5 and 0·7–3·6 wt %, respectively), as is the amount of CO2 (0·2–10·9 wt %).
Carbonatites typically contain <18 wt % SiO2 but may approach 22 wt % SiO2 (Fig. 13). They classify chemically as magnesiocarbonatite but a few straddle the boundary to ferruginous calciocarbonatite (Gittins & Harmer, 1997). In the Aillik Bay UML suite, they are the rock types with lowest concentrations of TiO2 and Al2O3 (typically <2·6 and 3·3 wt %, respectively). CaO (17–38·4 wt %), MgO (6·3–16·8 wt %), Fe2O3 (5·3–11·9) and MnO (up to 0·6 wt %) concentrations are high in both dolomite–calcite carbonatite and dolomite carbonatite. The CO2 content of the dolomite carbonatite is higher (32·2–39·5 wt %) than in the mafic silicate-bearing calcitic carbonatites (17·3–30·5 wt %; Fig. 13). K2O and P2O5 contents approach 2·5 and 4·5 wt %, respectively. Dolomite carbonatite is Ni and Cr depleted (<60 ppm), but concentrations in dolomite–calcite carbonatite may be up to 300 and 750 ppm, respectively.
Incompatible trace elements
Primitive mantle-normalized incompatible element abundances are displayed in Fig. 16. Aillikites are strongly enriched in Cs, Ba, Th, U, Nb, Ta and LREE with normalized concentrations of up to 600 x primitive mantle. Relative depletions are apparent at Rb, K, Pb, Sr, P, Zr, Hf and the heavy REE (HREE), which can be as low as 3 x primitive mantle (Fig. 16a). REE fractionation is extreme, with LaN/YbN between 70 and 136 (Fig. 13) and SmN/YbN between 18 and 30. The strong relative Zr–Hf depletion in aillikites, quantified by a (Zr + Hf)/(Zr + Hf)* ratio of 0·15–0·41, whereby (Zr + Hf)* is interpolated between neighbouring Nd and Sm, is surpassed only by the carbonatites. Low Zr/Nb ratios (1·2–4) are an expression of the strong incompatible trace element enrichment. Mela-aillikites have lower incompatible element abundances than aillikites but exhibit a similar pattern (Fig. 16a). However, the relative Zr–Hf depletion (0·47–0·87) and LREE/HREE fractionation (LaN/YbN = 33–84) are less pronounced.
Damtjernites exhibit similar enrichments in Th, U, low field strength elements (LFSE) and LREE to aillikites. However, marked differences between the two UML variants are apparent in the elevated Nb, Ta, Zr, Hf and HREE concentrations of the damtjernites (Figs 15a and 16b). Interestingly, the intra-HFSE (high field strength element) fractionation is similar in the two UML types, as exemplified by broadly overlapping Zr/Nb ratios of damtjernites (2–6) and aillikites (1·2–4; Fig. 13). The relative Zr–Hf depletion (0·33–0·55) of damtjernites is often less pronounced than in aillikites (<0·41). The REE are less fractionated in the damtjernites than in aillikites, as indicated by their lower LaN/YbN (45–62; Fig. 13) and SmN/YbN (11–17) ratios.
|
|
Incompatible elements in the carbonatites are extremely fractionated with strong Ba, Th and REE enrichment but Rb, K, Zr, Hf and Ti depletion (Fig. 16c). Whereas the patterns of the mixed dolomite–calcite carbonatites (LaN/YbN = 13–89) have some resemblance to those of aillikites, those for dolomite carbonatites (LaN/YbN = 36–377) are different, with LREE concentrations >1000 x primitive mantle, coupled with strong Zr–Hf depletions (0·01–0·02).
Sr–Nd isotope composition
Bulk-rock Sr and Nd isotope compositions are listed in Table 13. Carbonatites, aillikites, mela-aillikites and damtjernites show overlapping compositions forming a horizontal array in Sr–Nd isotope space close to Bulk Earth (Fig. 17). Age-corrected Nd(582) values (+0·1 to +1·9) fall within a narrow range, and 87Sr/86Sr(i) ratios vary from 0·70369 to 0·70466. Two samples, the dolomite–calcite carbonatite ST126 and the damtjernite ST256, have a more radiogenic Sr isotope composition (0·70579 and 0·70662, respectively) defining the high 87Sr/86Sr(582) end of the horizontal array (Fig. 17). As leached fractions of ST126 and ST256 yielded a similarly radiogenic 87Sr/86Sr(582) (0·70568 and 0·70526, respectively) and these samples do not show any other anomalous features compared with the other samples, we consider their Sr isotope composition to be primary. The clinopyroxene–phlogopite nodule ST162I falls within the Sr–Nd compositional range of the UML and carbonatites (87Sr/86Sr(i) = 0·70393; Nd(i) = +0·25).
|
|
Carbon and oxygen isotope composition of the bulk-rock carbonate fraction
The bulk-rock carbonate fraction of aillikites varies relatively little in
13CPDB (–5·7 to –5) and 18OSMOW (9·4–11·6), with sample ST220II being exceptional in approaching 13C of –4 (Fig. 18; Table 14). The dolomite–calcite carbonatites have a heavier carbon isotope composition (–3·7 to –3·3 13C) than the associated aillikites, with only sample ST199 being less 13C enriched (–4·8), close to values typical for aillikites. The oxygen isotope composition (9·6–10·8 18O) is within the range of the aillikites. Only sample ST231A (13·0 18O), which visibly experienced carbonate recrystallization, falls outside this narrow range. The dolomite carbonatites have the isotopically heaviest carbon composition, with 13C between –2·8 and –2·7 (n = 2). The oxygen isotope composition (10·8–11·5 18O) covers the higher end of the aillikite range, slightly elevated in comparison with the calcite-bearing carbonatites. There is no correlation between C and O isotope composition of carbonates from aillikite and carbonatite (Fig. 18), or between stable isotope and Sr–Nd isotope ratios. Damtjernites contain a composite carbonate fraction (groundmass and segregations), which is highly variable in its carbon isotope composition (–7 to –3 13C) at fairly constant 18O between 9·9 and 11·4 (within the range of aillikites).
|
|
|
DISCUSSION |
|---|
Field relations, age determinations and radiogenic isotope signatures imply that the various UML and carbonatite types as well as their micaceous cognate inclusions are broadly coeval and appear to be related to a common mode of magma production that persisted over a period of 30–35 Myr. To evaluate the petrogenesis of UML in the Aillik Bay area, it is necessary to unravel the effects of low-pressure modification of the primary magma, thus clearing the way for constraining the nature of the magma source, assessing whether all UML types can be related to a common magma type, and considering the geodynamic conditions under which melting occurred.
Modification of the parental UML magma
Role of contamination and fractionation processes
Extensive interaction of the UML magma with continental crustal material can be ruled out by the strong Si and Al undersaturation, high Ce/Pb ratios (>25), positive Nd(582) values and fairly unradiogenic Sr isotope composition (87Sr/86Sr(582) typically <0·7045). The continental crust of the North Atlantic craton is characterized by extremely unradiogenic Nd and radiogenic Sr isotope compositions and therefore cannot have modified the primary UML magma (e.g. Makkovik Province gneisses Nd(582) –30; Saglek block gneisses 87Sr/86Sr(582) >0·73; Collerson et al., 1989; Kerr & Fryer, 1993). There is no correlation between UML/carbonatite isotope composition and chemical parameters such as Si, Al, K and Pb, which are typically elevated in continental crustal rocks.
Fine-grained microporphyritic aillikites with high Mg-numbers, and high Ni and Cr contents meet the criteria for near-primary mantle-derived magmas (Frey et al., 1978); Mg-number >68 and Ni >320 ppm. Furthermore, they resemble experimentally produced ‘higher-degree’ melts of synthetic carbonated garnet peridotite in the simple CMAS–CO2 system close to 5 GPa (Fig. 19; Gudfinnsson & Presnall, 2005). It should be noted that the accumulation of olivine macrocrysts typical of kimberlites (Mitchell, 1986) is seen in only few aillikite samples (L2, ST225) and is not considered significant.
|
Our field observation that carbonatites, aillikites and mela-aillikites are intimately related is supported by the strong compositional continuum in Si, Ti, Al, Ca, Mn, P, CO2, Cr, Ni, Nb and LREE contents. On variation diagrams, they define well-correlated trends from which damtjernites are typically displaced (Figs 13–15). These trends are controlled principally by a relative separation of carbonate and a 2 : 1 to 1 : 2 mix of olivine and phlogopite (Figs 13 and 19a). The most likely process is devolatilization of CO2-rich liquids/fluids from aillikite magma, which caused the compositional shift to either side of these trends with carbonatite and mela-aillikite being close to the separated end-members (see section on devolatilization).
Damtjernites cannot be interpreted unequivocally as near-primary magmas using the criteria of Frey et al. (1978), although these are based on melts being in equilibrium with dry peridotite and not with metasomatized source materials. The most primitive olivine-rich members approach ‘primary’ Mg-number, Ni and Cr contents, but more evolved types are almost olivine-free, resulting in a conspicuous bimodal distribution in terms of MgO, Ni and Cr content (Fig. 14). The cumulate textures formed by large subhedral olivine and rare composite clinopyroxene crystals within the fine-grained groundmass imply that these may represent cognate high-pressure phases. Clinopyroxene crystallization pressure ranges between 0·8 and 1·5 GPa (Nimis & Ulmer, 1998), similar to the estimates for the clinopyroxene-bearing cognate inclusions. An unrealistically high amount (30–40 wt %) of olivine removal (DMgO = 2·6; DNi = 4·9–6·5; following Herzberg & O'Hara, 2002), together with minor amounts of Cr-spinel, would be required to relate these two groups of damtjernites (Fig. 14b), which have an identical olivine-free groundmass. On the basis of the large damtjernite sample suite examined (n = 35), we can rule out a continuous fractionation series, leading us to favour a role for a silicate–carbonate unmixing process in their petrogenesis (see section on immiscibility).
Major and selected trace element variation diagrams show no gradation between aillikites and damtjernites (Figs 13 and 14), implying that they experienced different evolutionary histories (Fig. 19). There is no petrologically sound mineral assemblage that could relate the two contrasting rock types by crystal fractionation. Fractional crystallization of olivine and Cr-spinel, however, can account for some of the intra-group variation in aillikites and damtjernites, as illustrated by the distinct Ni and Cr vs MgO trends (Fig. 14).
Linking aillikites, mela-aillikites and carbonatites by a devolatilization process
Several lines of evidence indicate that the two distinct types of carbonatite occurring at Aillik Bay do not represent primary magma compositions. The dolomite carbonatite would be the more likely candidate for a primary magma composition (Sweeney, 1994; Harmer & Gittins, 1997; Lee & Wyllie, 1998), but its Fe-rich dolomite is highly evolved in terms of Fe/Mg distribution. Furthermore, the conspicuous lack of mafic silicate phases such as olivine and phlogopite, in combination with the granular texture and the extremely fractionated incompatible element distributions, indicate that the dolomitic carbonatites cannot represent liquids (Wyllie & Tuttle, 1960), but are more probably the product of a carbonate fluid extraction process.
In contrast, the presence of calcite and dolomite laths in the mixed carbonatites indicates that they crystallized from a liquid. Although the analyzed dolomite–calcite carbonatite and aillikite samples were collected from discrete subvertical dykes, lateral gradations between both rock types within single sills (Fig. 3) have been described from Cape Makkovik (Malpas et al., 1986; Foley, 1989a). Field relations imply that the carbonatite phase of the flat-lying sheets may represent a separated carbonatite liquid with associated fluid that was moving ahead of the more viscous silicate-rich aillikite magma, opening fissures during dyke or sill emplacement. Additional support for such a devolatilization process comes from the fluidized globular aillikite segregations (Fig. 5a), which clearly demonstrate that a CO2-rich liquid/fluid was expelled from the aillikite magma. The subordinate phlogopite with ‘aillikitic’ composition in dolomite–calcite carbonatites is best interpreted as being incorporated from the aillikite parent magma during flowage and disruption (Fig. 8a). However, the Al2O3-rich clinopyroxene phenocrysts (Fig. 9a), which do not occur in aillikites, seem to have crystallized from the separated carbonatite liquid, as carbonatites show a high potential to promote the Ca-Tschermaks component in clinopyroxene (Blundy & Dalton, 2000).
Our stable isotope results are consistent with a fractionation relationship between aillikite and carbonatite. Groundmass carbonate from type aillikite is isotopically the most ‘primitive’, straddling the compositional fields for primary mantle-derived carbonatite (Taylor et al., 1967; Clarke et al., 1994) in a conventional 13CPDB–18OSMOW diagram (Fig. 18). The carbonatites contain isotopically heavier carbon than the aillikites with an increase in 13C from aillikite through dolomite–calcite carbonatite to dolomite carbonatite of the order of 3, so that they fall outside the fields for primary carbonatite (e.g. Keller & Hoefs, 1995). Post-magmatic processes have been shown to have only little effect on the C isotopic composition, but may produce major changes in O isotopes (Deines, 1989; Santos & Clayton, 1995), which is not seen in the Aillik Bay samples (except for carbonatite ST231A).
The identical ‘mantle-like’ Sr–Nd isotope composition of aillikites and carbonatites (Fig. 17) suggests that their distinct carbon isotope composition (Fig. 18) does not reflect derivation from different sources, and also that hydrothermal alteration did not play a major role, as most of the Sr resides in the carbonates. Rayleigh fractionation of a common parent magma best explains the strong 13C enrichment seen from aillikites towards dolomitic carbonatites. This may be caused by release of a CO2-rich liquid/fluid (enriched in 13C; Mattey et al., 1990; Chacko et al., 1991; Deines, 2004) by diffusive separation from carbonate-rich aillikite magma as a result of near-surface decompression at high temperature (>600 °C). This liquid plus associated fluid moved ahead of the crystal- and inclusion-laden aillikite magma and opened cracks for UML dyke emplacement. It eventually crystallized to form 13C-enriched dolomite–calcite carbonatite and rare dolomite carbonatite preferentially at UML dyke terminations. Co-precipitation of calcite and dolomite has been experimentally demonstrated to occur from 880 °C down to 650 °C at 0·2 GPa in the CMS–CO2–H2O system (Otto & Wyllie, 1993) with an expanded dolomite stability field at low temperatures.
Quantifying this Rayleigh fractionation process [the fractionation factor at 1200° C was calculated using equation (4) of Deines (2004)], a typical aillikite sample (e.g. ST164) would have to release <70 vol. % of its carbonate fraction to cause an increase in 13C by >2, approaching observed values for the carbonatites. This amount of separated CO2-rich liquid plus fluid would leave a residual magma (carrying early olivine and phlogopite phenocrysts) that on further cooling crystallizes a groundmass assemblage resembling mela-aillikite (Fig. 5b and Fig. 5c), which contains abundant clinopyroxene, biotite, and Mg- and Mn-depleted spinels/ilmenites (Figs 8a, 9a and 10a), but <10 vol. % carbonate. This model is consistent with the occurrence of composite aillikite/carbonatite and aillikite/mela-aillikite dykes, further implying that this fractionation process operated at shallow intrusion levels.
A role for liquid immiscibility in the genesis of damtjernites?
Fractional crystallization and partial melting cannot relate the distinct mineral assemblage and chemical composition of contemporaneous and isotopically indistinguishable aillikites and damtjernites, leading us to invoke an unmixing process. Although immiscible silicate–carbonate liquids have been reported from natural mantle-derived xenoliths (Amundsen, 1987; Kogarko et al., 1995; Chalot-Prat & Arnold, 1999; van Achterbergh et al., 2004), they have been experimentally produced predominantly at crustal pressures. (Freestone & Hamilton, 1980; Kjarsgaard & Hamilton, 1989; Kjarsgaard et al., 1989; Lee & Wyllie, 1997a, Lee & Wyllie, 1997b). One of the major findings of these experiments was that primitive CO2-bearing mantle-derived alkaline magmas do not intersect the miscibility gap and that only their evolved derivatives approach the silicate limb of the solvus.
Aillikite and damtjernite are unlikely to represent conjugate liquids given their rather small compositional differences (Fig. 13). It is more reasonable to consider an intermediate role for aillikite magma as the carbonated silicate parent that underwent immiscibility, but aillikite compositions do not fall in any published experimentally determined silicate–carbonate miscibility gaps (Fig. 20a), which mainly deal with nephelinitic systems (e.g. Lee & Wyllie, 1997a). However, fairly large miscibility gaps have been found in a CO2-saturated SNAC–CO2 system between 1225–1325 °C at 1·5–2·5 GPa (Fig. 20a; Brooker, 1998; Brooker & Kjarsgaard, in preparation). Although damtjernites fall along the 1275 °C/1·5 GPa immiscibility solvus, aillikite compositions seen at the surface are too alkali poor to enter the miscibility gap. However, it is reasonable to assume that a ‘transitional’ magma type such as aillikite has reacted with uppermost mantle wall-rocks (Moore & Wood, 1998), thereby losing alkalis; thus, the high-pressure equivalents may lie in the miscibility gap (Fig. 20a). The glimmerite nodules, which are present in every aillikite dyke examined, may provide a link to a more alkaline proto-aillikite magma that was able to exsolve a damtjernite liquid and a conjugate calciocarbonatite liquid.
|
Although sövitic carbonatites have not been found during our recent field work, two calciocarbonatite dykes were reported by Hawkins (1977)
and Foley (1982) from Aillik Bay. Their apparent scarcity is probably due to the fact that access to the radial dyke swarm is limited to its southern periphery and the centre of a hypothetical complex, where sövites may be more common, is covered by the sea (Fig. 2). Examples for UML–carbonatite complexes with such a central sövitic core are the Fen (Andersen, 1988; Andersen & Austrheim, 1991; Dahlgren, 1994) and Alnö complexes (Kresten, 1980; Vuorinen & Skelton, 2004) in Scandinavia, and the Callander complex in Ontario (Ferguson & Currie, 1971).
Proto-aillikite magma may have started to line the conduits with glimmerite material causing a loss of alkalis, Si, Al and Mg thereby lowering Mg/Ca but elevating Si/Al in the melt (Figs 19b and 20a). These subtle but critical compositional modifications (see Lee & Wyllie, 1997a) may have prevented the derivative magma from approaching the miscibility gap (Fig. 20a), allowing its direct ascent to the surface as glimmerite-laden aillikite. However, proto-aillikite magma that moved through previously reaction-lined conduits did not lose appreciable alkalis, so that it could not escape the unmixing process described above (Fig. 21a). Unmixing can explain the observation that phlogopite phenocrysts in aillikite and damtjernite have similar core compositions derived from common magmas, but later follow contrasting evolutionary paths (Fig. 8a and b). Moreover, the high NiO content (up to 0·5 wt %; Fig. 7) of early damtjernite olivines (<Fo86) indicates that immiscible damtjernite separation from proto-aillikite magma occurred when the latter was already slightly evolved but still alkali-rich, because olivines crystallized from strongly alkaline melts have exceptionally high DNi (Foley & Jenner, 2004).
|
Trace element modelling is also consistent with the proposed relation between damtjernite and aillikite by liquid immiscibility. Using partitioning data determined for nephelinitic systems at pressures of up to 1 GPa (Hamilton et al., 1989
; Jones et al., 1995; Veksler et al., 1998), we have calculated the budget of crucial trace elements for a hypothetical carbonatite magma conjugate to damtjernite (Fig. 20b). These trace element distributions resemble those of immiscible carbonatites (e.g. Oldoinyo Lengai; Bizimis et al., 2003), and the trace element abundances of aillikites are indeed transitional between the damtjernite and carbonatite conjugate pairs. This demonstrates that a proto-aillikite magma could have been the parental liquid to damtjernite, and that trace element partitioning between immiscible carbonate–silicate liquids can fully account for the observed differences in the large ion lithophile element (LILE), HFSE and HREE budget of rock types (Figs 16b and 20b).
Petrogenetic significance of cognate inclusions in aillikites
The mineralogy of clinopyroxene–phlogopite and olivine–phlogopite inclusions in aillikites resembles both aillikites (e.g. Mg-ilmenite, Cr-spinel) and damtjernites (e.g. Ti-Al-rich clinopyroxene, titanite), but also shows individual peculiarities (e.g. high Mn-olivine, zirconolite). These inclusions may be genetically related, representing different components in a continuum of compositions. The constituent minerals seem to have crystallized from a highly alkaline, carbonated silicate melt (isotopically similar to aillikite and damtjernite) at uppermost mantle to lower crustal depths (25–45 km). The presence of a CO2-rich phase is witnessed by the common replacement of olivine by carbonate. A cumulate origin is also indicated by the banded structure of larger nodules. It is likely that this material lines the conduits of the alkaline intrusions beneath Aillik Bay (Fig. 21a) and was disrupted by later batches of carbonate-rich aillikite magma. The parental magma to these nodules presumably never reached the surface, but clearly shows UML affinity. This multi-stage model, which is broadly similar to the suggested origin for phlogopite clinopyroxenite nodules in the high-K lavas of western Italy (Giannetti & Luhr, 1990), is consistent with the 35 Myr time span (590–555 Ma) of Aillik Bay area UML magmatism.
Glimmerites are probably the product of a similar wall-rock coating process, but there are strong indications that they are genetically related to carbonate-rich proto-aillikite. Whereas clinopyroxene–phlogopite and olivine–phlogopite nodules are in equilibrium with their host aillikite magma regarding volatile fugacities (see also similar mica compositions; Fig. 8), the distinctively higher relative HF fugacity of phlogopite–apatite pairs in glimmerites (Fig. 12), combined with the observation that glimmerite phlogopite coexists with pure orthoclase, suggests that glimmerites crystallized in CO2-rich, but H2O-poor conditions.
Mantle source characteristics and melting processes
Identification of the parental UML magma
Amongst the UML rock types from Aillik Bay, aillikite is considered to be closest to a primary magma composition. In contrast, damtjernites have many non-primary features, so we do not consider them during subsequent discussions on mantle processes. Aillikite is the only rock type carrying centimetre-sized cognate inclusions (Fig. 5e) testifying to rapid magma ascent, although mantle peridotite xenoliths have not been found.
Compositional modification by crystallization of glimmerite presumably at uppermost mantle levels caused only small deviations from primary compositions (Fig. 19b). Recalculation of aillikite sample ST164 to correct for a loss of up to 10 wt % glimmerite material (95 :5 ; phlogopite:apatite) would lower SiO2 (1·5 wt %), K2O (0·8), Al2O3 (0·5) and MgO (0·5) but elevate CaO (1·5) and FeO (0·6) relative to the hypothetical proto-aillikite magma.
Constraints on the mineralogy and location of the source region
The high MgO and CO2 content, as well as the potassic nature of aillikite, has to be explained in terms of partial melting of a carbonate-rich peridotitic source with an essential contribution from a K-bearing phase. Partial melting can be expected to have occurred under pressures greater than 5 GPa, but presumably not much greater than 7 GPa, as indicated by the rare occurrence of diamonds in aillikites from the North Atlantic region and by the equilibration pressures of entrained lithospheric mantle peridotite xenoliths (Larsen & Rønsbo, 1993; Mitchell et al., 1999; Digonnet et al., 2000; Bizzarro & Stevenson, 2003; Griffin et al., 2004). High-pressure melting experiments by Dalton & Presnall (1998) and Gudfinnsson & Presnall (2005), conducted on synthetic carbonated peridotite material (CMAS–CO2 system), are the only studies available that may be applicable to the genesis of aillikite, although important components such as alkalis, H2O, FeO and TiO2 are lacking. Indeed, those workers have produced near-solidus carbonate-rich liquids between 3 and 8 GPa that resemble primary carbonatites, and carbonate-rich ultramafic magmas such as aillikite and kimberlite at slightly higher degrees of melting. The most primitive aillikites from Aillik Bay resemble those melts that segregated significantly above the carbonate-bearing solidus close to 5 GPa in the Gudfinnsson & Presnall (2005) experiments (Fig. 19). However, it must be stressed that the co-occurrence of aillikite and carbonatite at Aillik Bay cannot be interpreted in terms of a melting continuum; instead they are related by a low-pressure fractionation process (Fig. 19).
The high MgO and Ni content of aillikite requires a major contribution from an olivine-rich peridotitic mantle. Furthermore, the strong LREE/HREE fractionation of aillikites at very low HREE concentration levels can be explained only in terms of melting in the presence of residual garnet. The abundant carbonate in aillikite is isotopically consistent with mantle derivation (13C –5, Deines, 2002) and bulk-rock CO2 correlates positively with CaO (Fig. 15b), but not with MgO, an observation also made for many kimberlites (Bailey, 1984). This implies that the mantle carbonate that contributed to the aillikite magma was CaCO3 rather than dolomite or magnesite, although peridotite melting at >5 GPa must have occurred in the magnesite stability field (Brey et al., 1983). Given that the most primitive aillikites have MgO/CaO ratios between 1 and 1·3 (Fig. 19b), much lower than in experimentally produced near-solidus melts of magnesite-bearing peridotite (up to 25; Brey et al., 1983), calcite is the most likely source of the carbonate at the P–T conditions of cratonic mantle lithosphere (Biellmann et al., 1993), but only as part of a non-peridotitic vein assemblage at 4–7 GPa.
Aillikite magma clearly segregated from a mantle source that contained an early melting hydrous K-bearing phase. Phlogopite and K-richterite are known from metasomatized mantle assemblages (Dawson & Smith, 1977; Waters, 1987; Ionov & Hofmann, 1995; Grégoire et al., 2002, 2003), both being stable to pressures above 7 GPa (Sudo & Tatsumi, 1990; Foley, 1991; Konzett et al., 1997; Konzett & Ulmer, 1999). Phlogopite is considered the most likely K-bearing phase in the melting assemblage because of its potential to produce silica-undersaturated melts with extremely high K/Na ratios. In contrast, K-richterite was demonstrated to melt out close to the solidus of ultramafic assemblages, yielding SiO2-rich melt compositions that are more akin to lamproites (Foley et al., 1999). The impact of residual source phlogopite on the incompatible element patterns of aillikites (Fig. 16a) is clearly seen at the pronounced troughs at K and Rb, but not for Cs owing to a DCs (0·6) that is an order of magnitude lower than DRb (5·2; Foley et al., 1996). Ba is strongly scattered and less meaningful for evaluating the role of phlogopite in the presence of mantle carbonate. Because K is a stoichiometric component in phlogopite, its content in the melt will be near-constant as long as phlogopite is residual, varying mainly as a function of the proportion to which phlogopite enters the melt. The K content of a melt in equilibrium with mantle phlogopite that melts to an extent of 20–50% (Greenough, 1988) is 1·6–5 wt %. The average K content of aillikite (1·8 ± 0·3 wt %) is at the lower end of this range, which is consistent with lower K-saturation levels in undersaturated melts produced under CO2-rich conditions (Rogers et al., 1992). Alternatively, proto-aillikite might have lost up to 1 wt % K2O by wall-rock coating with glimmerite material.
Apatite is an essential constituent of the source region given the high P2O5 concentrations in aillikite (2·2 ± 0·7 wt %; Fig. 13). The imprint of residual apatite on the incompatible element patterns may be seen at the Sr–P trough and probably at the U spike (Fig. 16a) given its potential for fractionating Th (DTh 0·12) from U (DU 0·001; Klemme & Dalpe, 2003).
Although the comparably high TiO2 (3·4 ± 0·5 wt %) concentrations of aillikite can be explained by melting Ti-rich phlogopite, it is more likely that an early melting Ti-rich oxide phase, probably ilmenite, controlled the HFSE budget during melting given the extremely high Nb–Ta abundances (181 ± 28 ppm Nb). Melting experiments on non-peridotitic ultramafic vein assemblages showed that ilmenite melts out quickly, whereas rutile persists, as does apatite, to higher temperatures (Foley et al., 1999). This observation makes ilmenite the more likely titanate that was present in the aillikite magma source region. Interestingly, the relative Zr–Hf depletion in aillikites is decoupled from Ti and Nb–Ta (Fig. 16a) and, therefore, unlikely to be the effect of a residual titanate. The relatively low but variable Zr–Hf concentrations in aillikite are not caused by mantle source carbonate, which is typically depleted in Zr and Hf (Ionov, 1998; Moine et al., 2004), as there is no negative correlation between absolute Zr–Hf abundances and aillikite/carbonatite CO2 content (Fig. 15a). A residual Ti-free oxide phase such as baddeleyite, which was found as mantle xenocrysts in the Ile Bizard alnöite (Heaman & LeCheminant, 2001) and has been synthesized as part of a carbonate-bearing metasomatic vein assemblage under upper mantle conditions by Meen et al. (1989), can account for the observed Zr–Hf troughs in the aillikite trace element spectra (Fig. 16a). The likely presence of residual baddeleyite in the melting assemblage, buffering Zr and Hf in the melt, is not expected to have significantly affected the melt REE distribution given its relatively low REE concentrations (Reischmann et al., 1995) and baddeleyite/carbonatite melt LREE to middle REE (MREE) partition coefficients close to unity (Klemme & Meyer, 2003).
An important constraint on the location of the mantle source region comes from the thermal stability of the required source mineralogy. We have pointed out that phlogopite and carbonate are essential in the melting assemblage and both phases are not stable at the temperatures of convecting upper mantle (1480 °C; McKenzie & Bickle, 1988). They are stable at the P–T conditions of the cold mantle lithosphere (Wendlandt & Eggler, 1980; Mengel & Green, 1989; Sweeney et al., 1993; Dalton & Wood, 1995; Ulmer & Sweeney, 2002) and this restricts aillikite generation to lithospheric portions of the cratonic mantle (<1400 °C; McKenzie et al., 2005). It does not, however, rule out a contribution from the convecting asthenospheric upper mantle in the form of metasomatizing agents (Fig. 21a). We favour a veined lithospheric mantle source over pervasively metasomatized peridotite for the genesis of aillikite, as it can account for the occurrence of minerals that are not in equilibrium with peridotite.
Isotopic constraints on the age and style of mantle metasomatism
The Nd isotope composition of aillikites, close to present-day Bulk Earth (Fig. 17), does not advocate a long-term enrichment of the UML mantle source in incompatible trace elements. LREE enrichment must have occurred shortly prior to melting, without subsequent aging to produce negative Nd values. Calculated Nd model ages of 1·0–1·1 Ga (based on Sm/Nd in aillikites) indicate the maximum LREE extraction age from a depleted mantle reservoir, because the Sm/Nd of a melt is generally lower than that of its source. Thus, it seems likely that the enrichment event did not precede Late Neoproterozoic UML magmatism by more than 400 Myr (probably much less), which is at least 300 Myr after the Mesoproterozoic lamproite magmatism had occurred in this region (1374 Ma; Tappe et al., in preparation).
The Sr–Nd data from Aillik Bay UML samples define a horizontal array (Fig. 17), which may be explained by enhanced radiogenic Sr in-growth in the source region because of the extremely high Rb/Sr ratios of phlogopite concentrated in veins, whereas there is no phase present that could cause such a rapid change of the Nd isotope composition. If representative Rb and Sr concentrations for mantle phlogopites (45–300 and 14–175 ppm, respectively; Grégoire et al., 2003) are considered, then only 50–200 Myr would be needed to alter an initial asthenospheric 87Sr/86Sr ratio of 0·7029 (Zindler & Hart, 1986) to the maximum values measured for the Aillik Bay UML (0·7066). Such a ‘phlogopite signature’ can, therefore, be produced within only a few tens of million years prior to magmatism and has been reported from other UML and olivine melilitite occurrences (Rogers et al., 1992; Andronikov & Foley, 2001; Riley et al., 2003). The fact that carbonatites follow the Nd–Hf mantle array led Bizimis et al. (2003) to conclude that their carbonated mantle source regions underwent rapid remelting given the potential of mantle carbonate to produce radiogenic Hf while leaving Nd isotopes unaffected. Taken together, this reinforces our argument that the metasomatic carbonate–phlogopite assemblage that gave rise to the production of Late Neoproterozoic UML magmas was short-lived and presumably formed a vein network at the base of the cratonic lithosphere (Fig. 21).
|
PETROGENESIS OF PARENTAL AILLIK BAY UML MAGMA AS PART OF THE NORTH ATLANTIC ALKALINE PROVINCE |
|---|
A number of rifting episodes affected the cratonic North Atlantic region during Middle to Late Proterozoic times and eventually led to the opening of the Iapetus Ocean and the break-up of the supercontinent Rodinia. Associated alkaline and carbonatitic igneous activity occurred from the St. Lawrence Valley Rift system (Gittins et al., 1967
, 1975; Doig & Barton, 1968; Ferguson & Currie, 1971) to Scandinavia (Griffin & Taylor, 1975; Brueckner & Rex, 1980; Kresten, 1980; Dahlgren, 1994; Meert et al., 1998; O'Brien et al., 2005), with the Aillik Bay UML suite forming an integral part of this North Atlantic alkaline province (Doig, 1970). Deep melting events at around 1400–1200 Ma produced lamproites, and can be placed within intact long-term enriched subcontinental lithospheric mantle (strongly negative Nd and highly unradiogenic Pb; Nelson, 1989; Tappe et al., in preparation), whereas the widespread Late Neoproterozoic UML magmas (600–550 Ma), such as those from Aillik Bay, show an imprint from juvenile asthenosphere-derived material (positive Nd coupled to incompatible trace element enrichment). This indicates that progressive continental stretching resulted in an incipient protrusion of hotter asthenosphere to shallow levels beneath the fractured continental margins (Fig. 21a). As a result of this lithospheric thinning, the cratonic geotherm was displaced to higher temperatures more similar to that of rift margins (Thompson & Gibson, 1994). Additionally, a depression of the former cratonic lithosphere solidus by volatile fluxing and oxidation (Foley, 1988; Taylor & Green, 1988) within the rifted mantle triggered small-degree melting under CO2-bearing conditions (Fig. 21b). The depth interval of initial CO2-present melting in the upwelling asthenosphere is considered to have been in excess of 5 GPa given the potential presence of diamonds in the overlying cratonic lithosphere. The small melt fraction produced had carbonatite-like characteristics (Wyllie, 1980; Dalton & Presnall, 1998) and during ascent quickly encountered the cold base of the cratonic lithosphere where the melts solidified because of their low heat capacity (Spera, 1984, 1987; McKenzie, 1989; Meen et al., 1989) producing carbonate-phlogopite dominated veins with minor apatite, ilmenite and baddeleyite. Such veining is to be expected at the transition from the porous to a channelized flow regime, which may roughly coincide with the asthenosphere–lithosphere boundary (McKenzie, 1985; Foley, 1988, 1992). Continued lithospheric extension further moved the asthenosphere–lithosphere boundary upwards and outwards (indicated as steps 1–2–3 in Fig. 21a), so that a newly adjusted geotherm (from geotherm 1 to 2 in Fig. 21b) allowed remelting of the vein assemblage. These potassic silicate–carbonate vein melts infiltrated the garnet peridotite wall-rock, which caused its extensive volatile-fluxed melting and resulted in a hybrid carbonate-rich UML magma type (proto-aillikite). Whereas the high incompatible trace element contents and, in turn, Sr–Nd isotope compositions of aillikites are dominated by the vein melt, the high MgO and compatible trace element concentrations are controlled by melting in the surrounding peridotite.
A multi-stage veined mantle melting model for UML magma production beneath an incipiently rifted cratonic area (Fig. 21) not only accounts for the composition of aillikites, but also explains the relatively large magma volumes (compared with kimberlite clusters; see Tappe et al., 2004) and the long time span of continuous UML magmatism in the Aillik Bay area (35 Myr) and elsewhere along the borders of the Labrador Sea (Heaman, 2005; Tappe et al., 2005b). However, primary carbonate-rich UML magmas (proto-aillikite) seem to reach upper crustal levels only rarely, although they may be abundant early components of extension-related continental magmatism.
|
SUPPLEMENTARY DATA |
|---|
Supplementary data for this paper are available at Journal of Petrology online.
|
APPENDIX A: SAMPLE LIST FOR AILLIK BAY AREA UML AND CARBONATITES |
|---|
|
|
APPENDIX B: ANALYTICAL TECHNIQUES |
|---|
U–Pb perovskite geochronology
Nine ultramafic lamprophyre hand specimens were processed through standard crushing and mineral separation procedures (Wilfley table, methylene iodide, Frantz isodynamic separator) at the University of Alberta, Edmonton, following the techniques described by Heaman & Kjarsgaard (2000)
. Perovskite recovery was best in the 40–120 µm range. Fresh euhedral crystals devoid of inclusions were individually selected, collected as morphological fractions and dissolved as such in a mix of HF and HNO3 (usually more than 100 grains per fraction to obtain >50 µg perovskite). Uranium and lead were isolated and concentrated from perovskite using a HBr anion exchange chromatography technique. The isotopic composition of these elements was measured on a VG354 thermal ionization mass spectrometer operating in single Faraday or analogue Daly mode.
All isotopic data reported in Table 1 were corrected for mass discrimination (+0.09%/a.m.u. Pb and +0.16%/a.m.u. U), tracer and blank contribution; uncertainties are reported at 1. Furthermore, the presence of initial common lead was corrected using the crustal lead evolution model of Stacey & Kramers (1975). The 206Pb/238U perovskite ages were shown to be most robust because they are least sensitive to this initial common lead correction (Heaman, 1989; Heaman & Kjarsgaard, 2000). All perovskite analyses are concordant and thus allow for the calculation of multi-fraction ages using a weighted mean approach (Ludwig, 1998).
40Ar/39Ar mica thermochronology
The clinopyroxene–phlogopite nodule ST162I was processed for 40Ar/39Ar analysis of phlogopite plates by standard mineral separation techniques, including hand-picking of inclusion-free unaltered crystals in the size range 0·5–1 mm. The phlogopite crystals were loaded into an aluminum foil packet and arranged radially in an aluminum canister (40 mm x 19 mm), which contained the flux monitor PP-20 hornblende (Hb3gr equivalent) with an apparent age of 1072 Ma (Roddick, 1983). The canister was irradiated for 120 h in position 5c at the research reactor of McMaster University (Hamilton, Ontario) in a fast neutron flux (3 x 1016 neutrons/cm2).
Laser 40Ar/39Ar step-heating analysis of the irradiated sample was carried out at the Geological Survey of Canada, Ottawa. The sample was loaded into a 1·5 mm diameter hole in a copper planchet and stepwise heated under vacuum using a Merchantek MIR10 10 W CO2 laser equipped with a 2 mm x 2 mm flat-field lens. The released Ar gas was cleaned over getters for 10 min before isotope analysis using a VG3600 gas source mass spectrometer. Error calculation on individual steps follows the numerical error analysis routines outlined by Scaillet (2000), whereas error analysis on grouped data follows the algebraic methods of Renne et al. (1998). Neutron flux gradients were evaluated by analyzing the PP-20 flux monitors, which were interspersed among the sample packets throughout the sample canister, and by interpolating a linear fit against calculated J-factor and sample position. The error on the J-factor value reported in Table 2 is conservatively estimated at ±0.6% (2).
Blanks were measured before and after the sample analysis and levels varied in the range 40Ar = (1·4–1·5) x 10–6 nmol, 39Ar = (1·2–1·4) x 10–9 nmol, 38Ar = (0·7–1·2) x 10–9 nmol, 37Ar = (0·4–0·5) x 10–9 nmol, 36Ar = (4·6–5·7) x 10–9 nmol, all at ±20% uncertainty. Nucleogenic interference corrections are (40Ar/39Ar)K = 0·025 ± 0·005, (38Ar/39Ar)K = 0·011 ± 0·010, (40Ar/37Ar)Ca = 0·002 ± 0·002, (39Ar/37Ar)Ca = 0·00068 ± 0·00004, (38Ar/37Ar)Ca = 0·00003 ± 0·00003, (36Ar/37Ar)Ca = 0·00028 ± 0·00016. All errors are quoted at the 2 level of uncertainty (Fig. B1).
|
Mineral chemistry
Mineral chemistry data were obtained using a JEOL JXA 8900 RL electron microprobe at Mainz University. Operating voltage for most silicates and carbonates was 15 kV with a beam current of 12 nA and 8 nA, respectively. Opaque oxides, perovskite, rutile, titanite and garnet were analysed with an accelerating voltage of 20 kV and a beam current of 20 nA. The beam diameter varied between 1 and 10 µm depending on the volatile abundance in the mineral of interest. Counting time for common silicates was between 15 and 20 s on the peak, whereas trace element-rich accessories were measured for up to 50 s on the peak. International standards of natural materials were used for calibration and all data were reduced with a CITZAF procedure, except for the carbonates, for which a ZAF correction was applied.
The JEOL JXA 8900 RL electron microprobe at Göttingen University was used for high-precision halogen determination in apatite–phlogopite pairs, to calculate an equilibrium fluorine distribution. Fluorine was calibrated against a natural topaz standard under 15 kV and 15 nA operating conditions with a 10 µm beam spot. Hexagonal apatite cross-sections perpendicular to the crystallographic c-axis were avoided because of the variation of F and Cl X-ray intensity owing to anisotropic diffusion preferably along this crystallographic direction (Stormer et al., 1993).
Whole-rock geochemistry
Major and selected trace elements were measured on fused discs by standard X-ray fluorescence (XRF) spectrometry at the University of Greifswald (Tappe, 2004). A wide range of trace elements and REE were analysed by a combination of inductively coupled plasma atomic emission spectrometry (ICP-AES) and ICP-mass spectrometry (ICP-MS) after a fusion digestion and acid dissolution procedure, respectively (Activation Laboratories, Ancaster, Canada). Concentrations for elements determined by both the ICP-MS and XRF technique are comparable within analytical error (e.g. Sr, Ce, Y, Zr, U, Th, Pb). Volatiles were measured by a combination of loss on ignition and direct determination of CO2 using a C-S analyser.
Sr–Nd isotope composition
Sr–Nd isotope compositions were determined on the same powders as used for the major and trace element contents. We selected a Savillex beaker dissolution after tests on a variety of rock types, dissolved by both Teflon bomb and beaker technique, had confirmed that the isotope composition was identical within analytical error. The bomb dissolution was carried out in microcapsules that were placed together in an external Teflon-lined steel vessel heated in an oven at 160 °C for 7 days. Powders dissolved in Savillex beakers were attacked in a HF–HNO3 mixture on a hotplate for 3 days. After slow evaporation to near dryness, the samples were taken up in 6N HCl and heated again for 1 day, repeating this step up to three times until a clear solution was obtained. Sr and Nd were separated and concentrated using Biorad AG50W cation and Eichrom Ln-Spec anion exchange resin, respectively. Sr and Nd isotope compositions were measured on a VG 54-30 Sector (Ta single filaments) and Finnigan MAT 262 (Re double filaments) thermal ionization mass spectrometer, respectively, both operating in dynamic mode (GFZ Potsdam). During the measurement period, the NBS-987 Sr reference material yielded an average value for 87Sr/86Sr of 0·710265 ± 12 and the La Jolla standard yielded a 143Nd/144Nd value of 0·511850 ± 7 (2 of 11 measurements). The initial isotopic composition was calculated for an intrusion age of 582 Ma (U–Pb perovskite age of damtjernite ST140A), using the decay constants 1·42 x 10–11 year–1 and 6·54 x 10–12 year–1 for 87Rb and 147Sm, respectively.
Oxygen and carbon isotope composition
The oxygen and carbon isotope composition of bulk-rock carbonate fractions was measured at Göttingen University. Rock powders (<20 µm) were reacted with anhydrous H3PO4 under vacuum at 25 °C for 24 h to liberate the CO2 of the carbonates. The volume of collected CO2 gas was close to a 100% of the theoretical yield so that no isotope fractionation during dissolution of dolomite- and calcite-bearing samples is expected to have occurred (Al-Aasm et al., 1990). The purified CO2 was analysed using a Finnigan MAT-251 gas source mass spectrometer and measured isotope ratios are expressed as 13C and 18O relative to PDB (Pee Dee Belemnite) and SMOW (Standard Mean Ocean Water), respectively. Reproducibility was better than 0·1 for 13C and 0·2 for 18O as determined by repeated measurements (n = 5, 2) of an in-house limestone standard.
|
ACKNOWLEDGEMENTS |
|---|
We are indebted to Dejan Prelevic, Bruce Ryan, Mike Villeneuve and Zoran Jovanovic for helpful discussions that ensued during this study. Randy Edmunds, Lori Dyson and Mike Bishop are thanked for logistical support in the wilderness of Labrador; and Ralf Tappert and Michelle Goryniuk for their hospitality in Edmonton. Stacey Hagen, Judy Schultz and Kendra Siemens ably assisted the geochronology work. Burkhard Schulz-Dobrick is thanked for keeping the electron microprobe facility in Mainz in excellent running condition. Thoughtful comments by Lotte Larsen, Roger Mitchell, Teal Riley, Sally Gibson and Marjorie Wilson on this manuscript are gratefully acknowledged. This project is funded by the German Research Foundation (DFG grant Fo181/15) and NSERC operating grants to G.A.J. and L.M.H. This publication is Geological Survey of Canada Contribution 2005682.
*Corresponding author. Present address: Department of Earth and Atmospheric Sciences, University of Alberta, Edmonton, Alberta, Canada T6G 2E3. Telephone: +49-06131-39-24759. Fax: +49-06131-39-23070. E-mail: stappe{at}ualberta.ca
|
REFERENCES |
|---|
Al-Aasm I. S., Taylor B. E., South B. (1990) Stable isotope analysis of multiple carbonate samples using selective acid extraction. Chemical Geology 80:119–125.[Web of Science]
Alibert C. and Albarède F. (1988) Relationships between mineralogical, chemical, and isotopic properties of some North American kimberlites. Journal of Geophysical Research 93:7643–7671.
Amundsen H.E.F. (1987) Evidence for liquid immiscibility in the upper mantle. Nature 327:692–695.[CrossRef]
Andersen D. J., Lindsley D. H., Davidson P. M. (1993) QUILF: a Pascal program to assess equilibria among Fe–Mg–Mn–Ti oxides, pyroxenes, olivine, and quartz. Computers and Geosciences 19:1333–1350.[CrossRef]
Andersen T. (1988) Evolution of peralkaline calcite carbonatite magma in the Fen Complex, Southeast Norway. Lithos 22:99–112.[CrossRef][Web of Science]
Andersen T. and Austrheim H. (1991) Temperature–HF fugacity trends during crystallization of calcite carbonatite magma in the Fen Complex, Norway. Mineralogical Magazine 55:81–94.[Web of Science]
Andronikov A. V. and Foley S. F. (2001) Trace element and Nd–Sr isotopic composition of ultramafic lamprophyres from the East Antarctic Beaver Lake area. Chemical Geology 175:291–305.[CrossRef][Web of Science]
Arima M. and Edgar A. D. (1981) Substitution mechanisms and solubility of titanium in phlogopites from rocks of probable mantle origin. Contributions to Mineralogy and Petrology 77:288–295.[CrossRef][Web of Science]
Bailey D. K. (1984) Kimberlite: ‘the mantle sample’ formed by ultrametasomatism. In Kornprobst J. (Ed.). Kimberlites: Kimberlites and Related Rocks (Elsevier, Amsterdam)–333.
Ballhaus C., Berry R. F., Green D. H. (1990) Oxygen fugacity controls in the Earth's upper mantle. Nature 348:437–440.[CrossRef]
Ballhaus C., Berry R. F., Green D. H. (1991) High pressure experimental calibration of the olivine–orthopyroxene–spinel oxygen geobarometer: implications for the oxidation state of the upper mantle. Contributions to Mineralogy and Petrology 107:27–40.[CrossRef][Web of Science]
Beard A. D., Downes H., Vetrin V. R., Kempton P. D., Maluski H. (1996) Petrogenesis of Devonian lamprophyre and carbonatite minor intrusions, Kandalaksha Gulf (Kola Peninsula, Russia). Lithos 39:93–119.[CrossRef][Web of Science]
Biellmann C., Guyot F., Gillet P., Reynard B. (1993) High-pressure stability of carbonates: quenching of calcite-II, high-pressure polymorph of CaCO3. European Journal of Mineralogy 5:503–510.
Birkett T. C., McCandless T. E., Hood C. T. (2004) Petrology of the Renard igneous bodies: host rocks for diamond in the northern Otish Mountains region, Quebec. Lithos 76:475–490.[CrossRef][Web of Science]
Bizimis M., Salters V.J.M., Dawson J. B. (2003) The brevity of carbonatite sources in the mantle: evidence from Hf isotopes. Contributions to Mineralogy and Petrology 145:281–300.[CrossRef][Web of Science]
Bizzarro M. and Stevenson R. (2003) Major element composition of the lithospheric mantle under the North Atlantic craton: evidence from peridotite xenoliths of the Sarfartoq area, southwestern Greenland. Contributions to Mineralogy and Petrology 146:223–240.[CrossRef][Web of Science]
Blundy J. and Dalton J. A. (2000) Experimental comparison of trace element partitioning between clinopyroxene and melt in carbonate and silicate systems, and implications for mantle metasomatism. Contributions to Mineralogy and Petrology 139:356–371.[CrossRef][Web of Science]
Bond G. C., Nickeson P. A., Kominz M. A. (1984) Breakup of a supercontinent between 625 Ma and 555 Ma: new evidence and implications for continental histories. Earth and Planetary Science Letters 70:325–345.[CrossRef][Web of Science]
Brey G. P., Brice W. R., Ellis D. J., Green D. H., Harris K. L., Ryabchikov I.D. (1983) Pyroxene–carbonate reactions in the upper mantle. Earth and Planetary Science Letters 62:63–74.[CrossRef][Web of Science]
Brooker R. A. (1998) The effect of CO2 saturation on immiscibility between silicate and carbonate liquids: an experimental study. Journal of Petrology 39:1905–1915.[CrossRef][Web of Science]
Brueckner H. K. and Rex D. C. (1980) K–Ar and Rb–Sr geochronology and Sr isotopic study of the Alnö alkaline complex, northeastern Sweden. Lithos 13:111–119.[CrossRef][Web of Science]
Carmichael I.S.E. (1991) The redox states of basic and silicic magmas: a reflection of their source regions? Contributions to Mineralogy and Petrology 106:129–141.[CrossRef][Web of Science]
Cawood P. A., McCausland P.J.A., Dunning G. R. (2001) Opening Iapetus: constraints from the Laurentian margin in Newfoundland. Geological Society of America Bulletin 113:443–453.
Chacko T., Mayeda T. K., Clayton R. N., Goldsmith J. R. (1991) Oxygen and carbon isotope fractionations between CO2 and calcite. Geochimica et Cosmochimica Acta 55:2867–2882.[CrossRef][Web of Science]
Chalmers J. A. and Laursen K. H. (1995) Labrador Sea: the extent of continental and oceanic crust and the timing of the onset of sea-floor spreading. Marine and Petroleum Geology 12:205–217.[CrossRef][Web of Science]
Chalmers J. A. and Pulvertaft T.C.R. (2001) Development of the continental margins of the Labrador Sea: a review. In Wilson R.C.L., Whitmarsh R.B., Taylor B., Froitzheim N. (Eds.). Non-volcanic Rifting of Continental Margins: a Comparison of Evidence from Land and Sea. Geological Society, London, Special Publications 187:77–105.
Chalot-Prat F. and Arnold M. (1999) Immiscibility between calciocarbonatitic and silicate melts and related wall rock reactions in the upper mantle: a natural case study from Romanian mantle xenoliths. Lithos 46:627–659.[CrossRef][Web of Science]
Chian D., Louden K. E., Keen C. E., Reid I. (1995) Evolution of nonvolcanic rifted margins: new results from the conjugate margins of the Labrador Sea. Geology 23:589–592.
Clarke L.B., Le Bas M.J., Spiro B. (1994) Rare earth, trace element and stable isotope fractionation of carbonatites at Kruidfontein, Transvaal, South Africa. In Meyer H.O.A. and Leonardos O.H. (Eds.). Kimberlites, Related Rocks and Mantle Xenoliths (Companhia de Pesquisa de Recursos Minerais, Rio de Janeiro) pp. 236–251.
Collerson K. D., McCulloch M. T., Nutman A. P. (1989) Sr and Nd isotope systematics of polymetamorphic Archean gneisses from southern West Greenland and northern Labrador. Canadian Journal of Earth Sciences 26:446–466.
Connelly J. N., van Gool J.A.M., Mengel F. C. (2000) Temporal evolution of a deeply eroded orogen: the Nagssugtoqidian Orogen, West Greenland. Canadian Journal of Earth Sciences 37:1121–1142.
Culshaw N. G., Ketchum J., Barr S. (2000a) Structural evolution of the Makkovik Province, Labrador, Canada: tectonic processes during 200 Myr at a Paleoproterozoic active margin. Tectonics 19:961–977.[CrossRef][Web of Science]
Culshaw N. G., Brown T., Reynolds P. H., Ketchum J.W.F. (2000b) Kanairiktok shear zone: the boundary between the Paleoproterozoic Makkovik Province and the Archean Nain Province, Labrador, Canada. Canadian Journal of Earth Sciences 37:1245–1257.
Dahlgren S. (1994) Late Proterozoic and Carboniferous ultramafic magmatism of carbonatitic affinity in Southern Norway. Lithos 31:141–154.[CrossRef][Web of Science]
Dalton J. A. and Presnall D. C. (1998) The continuum of primary carbonatitic–kimberlitic melt compositions in equilibrium with lherzolite: data from the system CaO–MgO–Al2O3–SiO2–CO2 at 6 GPa. Journal of Petrology 39:1953–1964.[CrossRef][Web of Science]
Dalton J. A. and Wood B. J. (1995) The stability of carbonate under upper-mantle conditions as a function of temperature and oxygen fugacity. European Journal of Mineralogy 7:883–891.
Dawson J. B. (1971) Advances in kimberlite geology. Earth-Science Reviews 7:187–214.
Dawson J. B. and Smith J. V. (1977) The MARID (mica–amphibole–rutile–ilmenite–diopside) suite of xenoliths in kimberlite. Geochimica et Cosmochimica Acta 41:309–323.[CrossRef][Web of Science]
Deines P. (1989) Stable isotope variations in carbonatites. In Bell K. (Ed.). Carbonatites: Genesis and Evolution (Unwin Hyman, London) pp. 301–359.
Deines P. (2002) The carbon isotope geochemistry of mantle xenoliths. Earth-Science Reviews 58:247–278.
Deines P. (2004) Carbon isotope effects in carbonate systems. Geochimica et Cosmochimica Acta 68:2659–2679.[CrossRef][Web of Science]
Digonnet S., Goulet N., Bourne J., Stevenson R., Archibald D. (2000) Petrology of the Abloviak aillikite dykes, New Québec: evidence for a Cambrian diamondiferous alkaline province in northeastern North America. Canadian Journal of Earth Sciences 37:517–533.
Doig R. (1970) An alkaline rock province linking Europe and North America. Canadian Journal of Earth Sciences 7:22–28.
Doig R. and Barton J. M. Jr. (1968) Ages of carbonatites and other alkaline rocks in Quebec. Canadian Journal of Earth Sciences 5:1401–1407.
Fedortchouk Y. and Canil D. (2004) Intensive variables in kimberlite magmas, Lac de Gras, Canada and implications for diamond survival. Journal of Petrology 45:1725–1745.
Ferguson J. and Currie K. L. (1971) Evidence of liquid immiscibility in alkaline ultrabasic dikes at Callander Bay, Ontario. Journal of Petrology 12:561–585.
Foley S.F. (1982) Mineralogy, geochemistry, petrogenesis and structural relationships of the Aillik Bay alkaline intrusive suite, Labrador, Canada. M.Sc. thesis, Memorial University of Newfoundland, St. John's. 210.
Foley S. F. (1984) Liquid immiscibility and melt segregation in alkaline lamprophyres from Labrador. Lithos 17:127–137.[CrossRef][Web of Science]
Foley S. F. (1988) The genesis of continental basic alkaline magmas: an interpretation in terms of redox melting. Journal of Petrology, Special Lithosphere Issue 139–161.
Foley S. F. (1989a) Emplacement features of lamprophyre and carbonatitic lamprophyre dykes at Aillik Bay, Labrador. Geological Magazine 126:29–42.[Abstract]
Foley S. F. (1989b) Experimental constraints on phlogopite chemistry in lamproites: 1. The effect of water activity and oxygen fugacity. European Journal of Mineralogy 1:411–426.
Foley S. F. (1991) High-pressure stability of the fluor- and hydroxy-endmembers of pargasite and K-richterite. Geochimica et Cosmochimica Acta 55:2689–2694.[CrossRef][Web of Science]
Foley S. F. (1992) Vein-plus-wall-rock melting mechanism in the lithosphere and the origin of potassic alkaline magmas. Lithos 28:435–453.[CrossRef][Web of Science]
Foley S. F. and Jenner G. A. (2004) Trace element partitioning in lamproitic magmas—the Gaussberg olivine leucitite. Lithos 75:19–38.[CrossRef][Web of Science]
Foley S. F., Taylor W. R., Green D. H. (1986) The role of fluorine and oxygen fugacity in the genesis of the ultrapotassic rocks. Contributions to Mineralogy and Petrology 94:183–192.[CrossRef][Web of Science]
Foley S. F., Jackson S. E., Fryer B. J., Greenough J. D., Jenner G. A. (1996) Trace element partition coefficients for clinopyroxene and phlogopite in an alkaline lamprophyre from Newfoundland by LAM-ICP-MS. Geochimica et Cosmochimica Acta 60:629–638.[CrossRef][Web of Science]
Foley S. F., Musselwhite D. S., van der Laan S. R. (1999) Melt compositions from ultramafic vein assemblages in the lithospheric mantle: a comparison of cratonic and non-cratonic settings. In Gurney J. J., Gurney J. L., Pascoe M. D., Richardson S. H. (Eds.). Proceedings of the 7th International Kimberlite Conference (Red Roof Design, Cape Town) pp. 238–246.
Foley S. F., Andronikov A. V., Melzer S. (2002) Petrology of ultramafic lamprophyres from the Beaver Lake area of Eastern Antarctica and their relation to the breakup of Gondwanaland. Mineralogy and Petrology 74:361–384.[CrossRef][Web of Science]
Forbes W. C. and Flower M.F.J. (1974) Phase relations of titan-phlogopite K2Mg4TiAl2Si6O20(OH)2: a refractory phase in the upper mantle? Earth and Planetary Science Letters 22:60–66.[CrossRef][Web of Science]
Fraser K. J., Hawkesworth C. J., Erlank A. J., Mitchell R. H., Scott-Smith B. H. (1985) Sr, Nd and Pb isotope and minor element geochemistry of lamproites and kimberlites. Earth and Planetary Science Letters 76:57–70.[CrossRef][Web of Science]
Freestone I. C. and Hamilton D. L. (1980) The role of liquid immiscibility in the genesis of carbonatites: an experimental study. Contributions to Mineralogy and Petrology 73:105–117.[CrossRef][Web of Science]
Frey F. A., Green D. H., Roy S. D. (1978) Integrated models of basalt petrogenesis: a study of quartz tholeiites to olivine melilitites from south eastern Australia utilizing geochemical and experimental petrological data. Journal of Petrology 19:463–513.
Frost B. R. (1991) Introduction to oxygen fugacity and its petrologic importance. In Lindsley D. H. (Ed.). Oxide Minerals: Petrologic and Magnetic Significance. Mineralogical Society of America, Reviews in Mineralogy 25:1–9.
Frost D. J. and Wood B. J. (1997) Experimental measurements of the fugacity of CO2 and graphite/diamond stability from 35 to 77 kbar at 925 to 1650 °C. Geochimica et Cosmochimica Acta 61:1565–1574.[CrossRef][Web of Science]
Giannetti B. and Luhr J. F. (1990) Phlogopite-clinopyroxenite nodules from high-K magmas, Roccamonfina Volcano, Italy: evidence for a low-pressure metasomatic origin. Earth and Planetary Science Letters 101:404–424.[CrossRef][Web of Science]
Gittins J. and Harmer R. E. (1997) What is ferrocarbonatite? a revised classification. Journal of African Earth Sciences 25:159–168.[CrossRef]
Gittins J., Macintyre R. M., York D. (1967) The ages of carbonatite complexes in eastern Canada. Canadian Journal of Earth Sciences 4:651–655.
Gittins J., Hewins R. H., Laurin A. F. (1975) Kimberlitic–carbonatitic dikes of the Saguenay River valley, Quebec, Canada. Physics and Chemistry of the Earth 9:137–148.[CrossRef]
Gower C. F., Erdmer P., Wardle R. J. (1986) The Double Mer Formation and the Lake Melville rift system, eastern Labrador. Canadian Journal of Earth Sciences 23:359–368.
Green D. H. and Falloon T. J. (1998) Pyrolite: a Ringwood concept and its current expression. In Jackson I. (Ed.). The Earth's Mantle (Cambridge University Press, Cambridge) pp. 311–378.
Greenough J. D. (1988) Minor phases in the Earth's mantle: evidence from trace- and minor-element patterns in primitive alkaline magmas. Chemical Geology 69:177–192.[CrossRef][Web of Science]
Grégoire M., Bell D. R., Le Roex A. P. (2002) Trace element geochemistry of phlogopite-rich mafic mantle xenoliths: their classification and their relationship to phlogopite-bearing peridotites and kimberlites revisited. Contributions to Mineralogy and Petrology 142:603–625.[Web of Science]
Grégoire M., Bell D. R., Le Roex A. P. (2003) Garnet lherzolites from the Kaapvaal Craton (South Africa): trace element evidence for a metasomatic history. Journal of Petrology 44:629–657.
Griffin W. L. and Taylor P. N. (1975) The Fen damkjernite: petrology of a ‘central-complex kimberlite’. Physics and Chemistry of the Earth 9:163–177.[CrossRef]
Griffin W. L., O'Reilly S. Y., Doyle B. J., Pearson N. J., Coopersmith H., Kivi K., Malkovets V. G., Pokhilenko N. P. (2004) Lithosphere mapping beneath the North American plate. Lithos 77:873–922.[CrossRef][Web of Science]
Gudfinnsson G. H. and Presnall D. C. (2005) Continuous gradations among primary carbonatitic, kimberlitic, melilititic, basaltic, picritic, and komatiitic melts in equilibrium with garnet lherzolite at 3–8 GPa. Journal of Petrology 46:1645–1659.
Guo J. and Green T. H. (1990) Experimental study of barium partitioning between phlogopite and silicate liquid at upper-mantle pressure and temperature. Lithos 24:83–95.[CrossRef][Web of Science]
Hall J., Wardle R. J., Gower C. F., Kerr A., Coflin K., Keen C. E., Carroll P. (1995) Proterozoic orogens of the northeastern Canadian Shield: new information from the Lithoprobe ECSOOT crustal reflection seismic survey. Canadian Journal of Earth Sciences 32:1119–1131.
Hamilton D. L., Bedson P., Esson J. (1989) The behaviour of trace elements in the evolution of carbonatites. In Bell K. (Ed.). Carbonatites: Genesis and Evolution (Unwin Hyman, London) pp. 405–427.
Hamilton R. (1992) Geology and structural setting of ultramafic lamprophyres from Bulljah Pool, central Western Australia. Journal of the Royal Society of Western Australia 75:51–56.
Hammarstrom J. M. and Zen E.-A. (1986) Aluminum in hornblende: an empirical igneous geobarometer. American Mineralogist 71:1297–1313.[Abstract]
Hansen K. (1980) Lamprophyres and carbonatitic lamprophyres related to rifting in the Labrador Sea. Lithos 13:145–152.[CrossRef][Web of Science]
Harmer R. E. and Gittins J. (1997) The origin of dolomitic carbonatites: field and experimental constraints. Journal of African Earth Sciences 25:5–28.[CrossRef]
Hawkins D. W. (1977) Emplacement, petrology and geochemistry of ultrabasic to basic intrusives at Aillik Bay, Labrador. M.Sc. thesis, Memorial University of Newfoundland, St. John's, Canada. 236.
Heaman L. M. (1989) The nature of the subcontinental mantle from Sr–Nd–Pb isotopic studies on kimberlitic perovskite. Earth and Planetary Science Letters 92:323–334.[CrossRef][Web of Science]
Heaman L. M. (2005) Patterns of kimberlite emplacement: the importance of robust geochronology. Report–Geological Survey of Denmark and Greenland 68:25.
Heaman L. M. and Kjarsgaard B. A. (2000) Timing of eastern North American kimberlite magmatism: continental extension of the Great Meteor hotspot track? Earth and Planetary Science Letters 178:253–268.[CrossRef][Web of Science]
Heaman L. M. and LeCheminant A. N. (2001) Anomalous U–Pb systematics in mantle-derived baddeleyite xenocrysts from Ile Bizard: evidence for high temperature radon diffusion? Chemical Geology 172:77–93.[CrossRef][Web of Science]
Heaman L. M., Kjarsgaard B. A., Creaser R. A. (2004) The temporal evolution of North American kimberlites. Lithos 76:377–397.[CrossRef][Web of Science]
Herzberg C. and O'Hara M. J. (2002) Plume-associated ultramafic magmas of Phanerozoic age. Journal of Petrology 43:1857–1883.
Ionov D. A. (1998) Trace element composition of mantle-derived carbonates and coexisting phases in peridotite xenoliths from alkali basalts. Journal of Petrology 39:1931–1941.[CrossRef][Web of Science]
Ionov D. A. and Hofmann A. W. (1995) Nb–Ta-rich mantle amphiboles and micas: implications for subduction-related metasomatic trace element fractionations. Earth and Planetary Science Letters 131:341–356.[CrossRef][Web of Science]
Janse A.J.A. and Sheahan P. A. (1995) Catalogue of world wide diamond and kimberlite occurrences: a selective and annotative approach. Journal of Geochemical Exploration 53:73–111.[CrossRef][Web of Science]
Jones J. H., Walker D., Pickett D. A., Murrell M. T., Beattie P. (1995) Experimental investigations of the partitioning of Nb, Mo, Ba, Ce, Pb, Ra, Th, Pa, and U between immiscible carbonate and silicate liquids. Geochimica et Cosmochimica Acta 59:1307–1320.[CrossRef][Web of Science]
Kamo S. L., Gower C. F., Krogh T. E. (1989) Birthdate for the Iapetus Ocean? A precise U–Pb zircon and baddeleyite age for the Long Range dikes, Southeast Labrador. Geology 17:602–605.
Kamo S. L., Krogh T. E., Kumarapeli P. S. (1995) Age of the Grenville dyke swarm, Ontario–Quebec: implications for the timing of Iapetan rifting. Canadian Journal of Earth Sciences 32:273–280.
Keen C. E., Courtney R. C., Dehler S. A., Williamson M. C. (1994) Decompression melting at rifted margins: comparison of model predictions with the distribution of igneous rocks on the eastern Canadian margin. Earth and Planetary Science Letters 121:403–416.[CrossRef][Web of Science]
Keller J. and Hoefs J. (1995) Stable isotope characteristics of recent natrocarbonatite from Oldoinyo Lengai. In Bell K. and Keller J. (Eds.). Carbonatite Volcanism: Oldoinyo Lengai and the Petrogenesis of Natrocarbonatites (Springer, Berlin) pp. 113–123.
Kennedy C. S. and Kennedy G. C. (1976) The equilibrium boundary between graphite and diamond. Journal of Geophysical Research 81:2467–2470.
Kerr A. and Fryer B. J. (1993) Nd isotope evidence for crust–mantle interaction in the generation of A-type granitoid suites in Labrador, Canada. Chemical Geology 104:39–60.[CrossRef][Web of Science]
Kerr A. and Fryer B. J. (1994) The importance of late- and post-orogenic crustal growth in the early Proterozoic: evidence from Sm–Nd isotopic studies of igneous rocks in the Makkovik Province, Canada. Earth and Planetary Science Letters 125:71–88.[CrossRef][Web of Science]
Kerr A. and Wardle R. J. (1997) Definition of an Archean–Proterozoic crustal suture by isotopic studies of basement intersections from offshore wells in the southern Labrador Sea. Canadian Journal of Earth Sciences 34:209–214.
Kerr A., Hall J., Wardle R. J., Gower C. F., Ryan B. (1997) New reflections on the structure and evolution of the Makkovikian–Ketilidian Orogen in Labrador and southern Greenland. Tectonics 16:942–965.[CrossRef][Web of Science]
Ketchum J.W.F., Culshaw N. G., Barr S. M. (2002) Anatomy and orogenic history of a Paleoproterozoic accretionary belt: the Makkovik Province, Labrador, Canada. Canadian Journal of Earth Sciences 39:711–730.
King A. F. and McMillan N. J. (1975) A Mid-Mesozoic breccia from the coast of Labrador. Canadian Journal of Earth Sciences 12:44–51.
Kjarsgaard B. A. and Hamilton D. L. (1989) The genesis of carbonatites by immiscibility. In Bell K. (Ed.). Carbonatites: Genesis and Evolution (Unwin Hyman, London Kingdom) pp. 388–404.
Kjarsgaard B. A., Hamilton D. L., Gittins J. (1989) Carbonatite origin and diversity: discussion and reply. Nature 338:547–548.[Web of Science]
Klemme S. and Dalpe C. (2003) Trace-element partitioning between apatite and carbonatite melt. American Mineralogist 88:639–646.
Klemme S. and Meyer H. P. (2003) Trace element partitioning between baddeleyite and carbonatite melt at high pressures and high temperatures. Chemical Geology 199:233–242.[Web of Science]
Kogarko L. N., Henderson C.M.B., Pacheco H. (1995) Primary Ca-rich carbonatite magma and carbonate–silicate–sulphide liquid immiscibility in the upper mantle. Contributions to Mineralogy and Petrology 121:267–274.[Web of Science]
Konzett J. and Ulmer P. (1999) The stability of hydrous potassic phases in lherzolite mantle: an experimental study to 9·5 GPa in simplified and natural bulk compositions. Journal of Petrology 40:629–652.[CrossRef][Web of Science]
Konzett J., Sweeney R. J., Thompson A. B., Ulmer P. (1997) Potassium amphibole stability in the upper mantle: an experimental study in a peralkaline KNCMASH system to 8·5 GPa. Journal of Petrology 38:537–568.[CrossRef][Web of Science]
Kresten P. (1980) The Alnö complex: tectonics of dyke emplacement. Lithos 13:153–158.[CrossRef][Web of Science]
Kretz R. (1982) Transfer and exchange equilibria in a portion of the pyroxene quadrilateral as deduced from natural and experimental data. Geochimica et Cosmochimica Acta 46:411–421.[CrossRef][Web of Science]
Kretz R. (1983) Symbols for rock-forming minerals. American Mineralogist 68:277–279.[Abstract]
Larsen L. M. and Rex D. C. (1992) A review of the 2500 Ma span of alkaline–ultramafic, potassic and carbonatitic magmatism in West Greenland. Lithos 28:367–402.[CrossRef][Web of Science]
Larsen L. M. and Rønsbo J. (1993) Conditions of origin of kimberlites in West Greenland: new evidence from the Sarfartoq and Sukkertoppen regions. Report–Geological Survey of Greenland 159:115–120.
Larsen L. M., Rex D. C., Watt W. S., Guise P. G. (1999) 40Ar–39Ar dating of alkali basaltic dykes along the southwest coast of Greenland: Cretaceous and Tertiary igneous activity along the eastern margin of the Labrador Sea. Geology of Greenland Survey Bulletin 184:19–29.
Leat P. T., Riley T. R., Storey B. C., Kelley S. P., Millar I. L. (2000) Middle Jurassic ultramafic lamprophyre dyke within the Ferrar magmatic province, Pensacola Mountains, Antarctica. Mineralogical Magazine 64:95–111.
Lee W. J. and Wyllie P. J. (1997a) Liquid immiscibility between nephelinite and carbonatite from 1·0 to 2·5 GPa compared with mantle melt compositions. Contributions to Mineralogy and Petrology 127:1–16.[CrossRef][Web of Science]
Lee W. J. and Wyllie P. J. (1997b) Liquid immiscibility in the join NaAlSiO4–NaAlSi3O8–CaCO3 at 1 GPa: implications for crustal carbonatites. Journal of Petrology 38:1113–1135.[CrossRef][Web of Science]
Lee W. J. and Wyllie P. J. (1998) Petrogenesis of carbonatite magmas from mantle to crust, constrained by the system CaO–(MgO + FeO*)–(Na2O + K2O)–(SiO2 + Al2O3 + TiO2)–CO2. Journal of Petrology 39:495–517.[CrossRef][Web of Science]
Leech G. B., Lowdon J. A., Stockwell C. H., Wanless R. K. (1962) Age determinations and geological studies. Reports Geological Survey of Canada 4:116–117.
Le Roex A. P. and Lanyon R. (1998) Isotope and trace element geochemistry of Cretaceous Damaraland lamprophyres and carbonatites, northwestern Namibia: evidence for plume–lithosphere interactions. Journal of Petrology 39:1117–1146.
Le Roex A. P., Bell D. R., Davis P. (2003) Petrogenesis of Group I kimberlites from Kimberley, South Africa: evidence from bulk-rock geochemistry. Journal of Petrology 44:2261–2286.
Ludwig K. R. (1998) On the treatment of concordant uranium–lead ages. Geochimica et Cosmochimica Acta 62:665–676.[CrossRef][Web of Science]
Malpas J., Foley S. F., King A. F. (1986) Alkaline mafic and ultramafic lamprophyres from the Aillik Bay area, Labrador. Canadian Journal of Earth Sciences 23:1902–1918.
Mattey D. P., Taylor W. R., Green D. H., Pillinger C. T. (1990) Carbon isotopic fractionation between CO2 vapour, silicate and carbonate melts: an experimental study to 30 kbar. Contributions to Mineralogy and Petrology 104:492–505.[CrossRef][Web of Science]
McKenzie D. (1985) The extraction of magma from the crust and mantle. Earth and Planetary Science Letters 74:81–91.[CrossRef][Web of Science]
McKenzie D. (1989) Some remarks on the movement of small melt fractions in the mantle. Earth and Planetary Science Letters 95:53–72.[CrossRef][Web of Science]
McKenzie D. and Bickle M. J. (1988) The volume and composition of melt generated by extension of the lithosphere. Journal of Petrology 29:625–679.
McKenzie D., Jackson J., Priestley K. (2005) Thermal structure of oceanic and continental lithosphere. Earth and Planetary Science Letters 233:337–349.[CrossRef][Web of Science]
Meen J. K., Ayers J. C., Fregeau E. J. (1989) A model of mantle metasomatism by carbonated alkaline melts: trace-element and isotopic compositions of mantle source regions of carbonatite and other continental igneous rocks. In Bell K. (Ed.). Carbonatites: Genesis and Evolution (Unwin Hyman, London) pp. 464–499.
Meert J. G., Torsvik T. H., Eide E. A., Dahlgren S. (1998) Tectonic significance of the Fen Province, S.Norway: constraints from geochronology and paleomagnetism. Journal of Geology 106:553–564.[Web of Science]
Mengel K. and Green D. H. (1989) Stability of amphibole and phlogopite in metasomatized peridotite under water-saturated and water-undersaturated conditions. In Ross J. (Ed.). Kimberlites and Related Rocks (Geological Society of Australia, Sydney N. S. W.) pp. 571–581.
Mitchell R. H. (1973) Composition of olivine, silica activity and oxygen fugacity in kimberlite. Lithos 6:65–81.[CrossRef]
Mitchell R. H. (1986) Kimberlites: Mineralogy, Geochemistry and Petrology (Plenum, New York) pp. 442.
Mitchell R. H. (1995) Kimberlites, Orangeites, and Related Rocks (Plenum, New York) pp. 410.
Mitchell R. H. and Bergman S. C. (1991) Petrology of Lamproites (Plenum, New York) pp. 447.
Mitchell R. H., Scott-Smith B. H., Larsen L. M. (1999) Mineralogy of ultramafic dikes from the Sarfartoq, Sisimiut and Maniitsoq areas, West Greenland. In Gurney J. J., Gurney J. L., Pascoe M. D., Richardson S. H. (Eds.). Proceedings of the 7th International Kimberlite Conference (Red Roof Design, Cape Town) pp. 574–583.
Moine B. N., Grégoire M., O'Reilly S. Y., Delpech G., Sheppard S.M.F., Lorand J. P., Renac C., Giret A., Cottin J. Y. (2004) Carbonatite melt in oceanic upper mantle beneath the Kerguelen Archipelago. Lithos 75:239–252.[CrossRef][Web of Science]
Moore K. R. and Wood B. J. (1998) The transition from carbonate to silicate melts in the CaO–MgO–SiO2–CO2 system. Journal of Petrology 39:1943–1951.[CrossRef][Web of Science]
Munno R., Rossi G., Tadini C. (1980) Crystal chemistry of kimzeyite from Stromboli, Aeolian Islands, Italy. American Mineralogist 65:188–191.[Abstract]
Murphy D. T., Collerson K. D., Kamber B. S. (2002) Lamproites from Gaussberg, Antarctica: possible transition zone melts of Archaean subducted sediments. Journal of Petrology 43:981–1001.
Murthy G., Gower C. F., Tubrett M., Paetzold R. (1992) Paleomagnetism of Eocambrian Long Range dykes and Double Mer Formation from Labrador, Canada. Canadian Journal of Earth Sciences 29:1224–1234.
Neal C. R. and Davidson J. P. (1989) An unmetasomatized source for the Malaitan alnöite (Solomon Islands): petrogenesis involving zone refining, megacryst fractionation, and assimilation of oceanic lithosphere. Geochimica et Cosmochimica Acta 53:1975–1990.[CrossRef][Web of Science]
Nelson D. R. (1989) Isotopic characteristics and petrogenesis of the lamproites and kimberlites of central West Greenland. Lithos 22:265–274.[CrossRef][Web of Science]
Nicholls J., Carmichael I.S.E., Stormer J. C. Jr. (1971) Silica activity and Ptotal in igneous rocks. Contributions to Mineralogy and Petrology 33:1–20.[CrossRef][Web of Science]
Nimis P. and Ulmer P. (1998) Clinopyroxene geobarometry of magmatic rocks: Part 1: an expanded structural geobarometer for anhydrous and hydrous, basic and ultrabasic systems. Contributions to Mineralogy and Petrology 133:122–135.[CrossRef][Web of Science]
Nixon P. H., Mitchell R. H., Rogers N. W. (1980) Petrogenesis of alnöitic rocks from Malaita, Solomon Islands, Melanesia. Mineralogical Magazine 43:587–596.[Web of Science]
Nowell G. M., Pearson D. G., Bell D. R., Carlson R. W., Smith C. B., Kempton P. D., Noble S. R. (2004) Hf isotope systematics of kimberlites and their megacrysts: new constraints on their source regions. Journal of Petrology 45:1583–1612.
O'Brien H. E., Peltonen P., Vartiainen H. (2005) Kimberlites, carbonatites, and alkaline rocks. In Lehtinen M., Nurmi P. A., Rämö O. T. (Eds.). Precambrian Geology of Finland—Key to the Evolution of the Fennoscandian Shield (Elsevier, Amsterdam) pp. 605–644.
O'Neill H.S.C. and Wall V. J. (1987) The olivine–orthopyroxene–spinel oxygen geobarometer, the nickel precipitation curve, and the oxygen fugacity of the Earth's upper mantle. Journal of Petrology 28:1169–1191.
Otto J. W. and Wyllie P. J. (1993) Relationships between silicate melts and carbonate-precipitating melts in CaO–MgO–SiO2–CO2–H2O at 2 kbar. Mineralogy and Petrology 48:343–365.[CrossRef][Web of Science]
Pearce N.J.G. and Leng M. J. (1996) The origin of carbonatites and related rocks from the Igaliko dyke swarm, Gardar Province, South Greenland: field, geochemical and C–O–Sr–Nd isotope evidence. Lithos 39:21–40.[CrossRef][Web of Science]
Platt R. G. and Mitchell R. H. (1979) The Marathon Dikes.I: Zirconium-rich titanian garnets and manganoan magnesian ulvöspinel–magnetite spinels. American Mineralogist 64:546–550.[Abstract]
Pollack H. N. and Chapman D. S. (1977) On the regional variation of heat flow, geotherms, and lithospheric thickness. Tectonophysics 38:279–296.[CrossRef][Web of Science]
Prelevic D., Foley S. F., Romer R. L., Cvetkovic V., Downes H. (2005) Tertiary ultrapotassic volcanism in Serbia: constraints on petrogenesis and mantle source characteristics. Journal of Petrology 46:1443–1487.
Price S. E., Russell J. K., Kopylova M. G. (2000) Primitive magma from the Jericho Pipe,N.W.T, Canada: constraints on primary kimberlite melt chemistry. Journal of Petrology 41:789–808.
Puffer J. H. (2002) A late Neoproterozoic eastern Laurentian superplume: location, size, chemical composition, and environmental impact. American Journal of Science 302:1–27.
Queen M., Heaman L. M., Hanes J. A., Archibald D. A., Farrar E. (1996) 40Ar/39Ar phlogopite and U–Pb perovskite dating of lamprophyre dykes from the eastern Lake Superior region: evidence for a 1·14 Ga magmatic precursor to Midcontinent Rift volcanism. Canadian Journal of Earth Sciences 33:958–965.
Raeside R. P. and Helmstaedt H. (1982) The Ile Bizard Intrusion, Montreal, Quebec: kimberlite or lamprophyre? Canadian Journal of Earth Sciences 19:1996–2011.
Reischmann T., Brügmann G. E., Jochum K. P., Todt W. A. (1995) Trace element and isotopic composition of baddeleyite. Mineralogy and Petrology 53:155–164.[CrossRef][Web of Science]
Renne P. R., Swisher C. C., Deino A. L., Karner D. B., Owens T. L., DePaolo D.J. (1998) Intercalibration of standards, absolute ages and uncertainties in 40Ar–39Ar dating. Chemical Geology 145:117–152.[CrossRef][Web of Science]
Righter K. and Carmichael I.S.E. (1996) Phase equilibria of phlogopite lamprophyres from western Mexico: biotite–liquid equilibria and P–T estimates for biotite-bearing igneous rocks. Contributions to Mineralogy and Petrology 123:1–21.[CrossRef][Web of Science]
Riley T. R., Leat P. T., Storey B. C., Parkinson I. J., Millar I. L. (2003) Ultramafic lamprophyres of the Ferrar large igneous province: evidence for a HIMU mantle component. Lithos 66:63–76.[CrossRef][Web of Science]
Robert J. L. (1976) Titanium solubility in synthetic phlogopite solid solutions. Chemical Geology 17:213–227.[CrossRef][Web of Science]
Rock N.M.S. (1986) The nature and origin of ultramafic lamprophyres: alnöites and allied rocks. Journal of Petrology 27:155–196.
Rock N. M. S. (1991) Lamprophyres (Blackie, Glasgow) pp. 285.
Roddick J. C. (1983) High precision intercalibration of 40Ar–39Ar standards. Geochimica et Cosmochimica Acta 47:887–898.[CrossRef][Web of Science]
Rogers N. W., Hawkesworth C. J., Palacz Z. A. (1992) Phlogopite in the generation of olivine melilitites from Namaqualand, South Africa and implications for element fractionation processes in the upper mantle. Lithos 28:347–365.[CrossRef][Web of Science]
Romer R. L., Schärer U., Wardle R. J., Wilton D. H. C. (1995) U–Pb age of the Seal Lake Group, Labrador: relationship to Mesoproterozoic extension-related magmatism of Laurasia. Canadian Journal of Earth Sciences 32:1401–1410.
Santos R. V. and Clayton R. N. (1995) Variations of oxygen and carbon isotopes in carbonatites: a study of Brazilian alkaline complexes. Geochimica et Cosmochimica Acta 59:1339–1352.[CrossRef][Web of Science]
Scaillet S. (2000) Numerical error analysis in 40Ar/39Ar dating. Chemical Geology 162:269–298.[CrossRef][Web of Science]
Scott B. H. (1981) Kimberlite and lamproite dykes from Holsteinsborg, West Greenland. Meddelelser om Grønland, Geoscience 4:3–24.
Seifert W., Kämpf H., Wasternack J. (2000) Compositional variation in apatite, phlogopite and other accessory minerals of the ultramafic Delitzsch Complex, Germany: implication for cooling history of carbonatites. Lithos 53:81–100.[CrossRef][Web of Science]
Sinclair G. S., Barr S. M., Culshaw N. G., Ketchum J. W. F. (2002) Geochemistry and age of the Aillik Group and associated plutonic rocks, Makkovik Bay area, Labrador: implications for tectonic development of the Makkovik Province. Canadian Journal of Earth Sciences 39:731–748.
Smith J. V., Brennesholtz R., Dawson J. B. (1978) Chemistry of micas from kimberlites and xenoliths: I. Micaceous kimberlites. Geochimica et Cosmochimica Acta 42:959–972.[CrossRef][Web of Science]
Spera F. J. (1984) Carbon dioxide in petrogenesis III: role of volatiles in the ascent of alkaline magma with special reference to xenolith-bearing mafic lavas. Contributions to Mineralogy and Petrology 88:217–232.[CrossRef][Web of Science]
Spera F. J. (1987) Dynamics of translithospheric migration of metasomatic fluid and alkaline magma. In Menzies M. A. and Hawkesworth C.J. (Eds.). Mantle Metasomatism (Academic Press, London) pp. 1–20.
Srivastava S. P. and Roest W. R. (1999) Extent of oceanic crust in the Labrador Sea. Marine and Petroleum Geology 16:65–84.[CrossRef][Web of Science]
Stacey J. S. and Kramers J. D. (1975) Approximation of terrestrial lead isotope evolution by a two-stage model. Earth and Planetary Science Letters 26:207–221.[CrossRef][Web of Science]
Stormer J. C., Pierson M. L., Tacker R. C. (1993) Variation of F and Cl X-ray intensity due to anisotropic diffusion in apatite during electron microprobe analysis. American Mineralogist 78:641–648.[Abstract]
Sudo A. and Tatsumi Y. (1990) Phlogopite and K-amphibole in the upper mantle: implication for magma genesis in subduction zones. Geophysical Research Letters 17:29–32.[Web of Science]
Sun S.-S. and McDonough W. F. (1989) Chemical and isotopic systematics of ocean basalts: implications for mantle composition and processes. In Saunders A. D. and Norry M. J. (Eds.). Magmatism in the Ocean Basins. Geological Society (Special Publications, London) 42: pp. 313–345.
Sweeney R. J. (1994) Carbonatite melt compositions in the Earth's mantle. Earth and Planetary Science Letters 128:259–270.[CrossRef][Web of Science]
Sweeney R. J., Thompson A. B., Ulmer P. (1993) Phase relations of a natural MARID composition and implications for MARID genesis, lithospheric melting and mantle metasomatism. Contributions to Mineralogy and Petrology 115:225–241.[CrossRef][Web of Science]
Tappe S. (2004) Mesozoic mafic alkaline magmatism of southern Scandinavia. Contributions to Mineralogy and Petrology 148:312–334.[CrossRef][Web of Science]
Tappe S., Jenner G. A., Foley S. F., Heaman L. M., Besserer D., Kjarsgaard B. A., Ryan A. B. (2004) Torngat ultramafic lamprophyres and their relation to the North Atlantic Alkaline Province. Lithos 76:491–518.[CrossRef][Web of Science]
Tappe S., Foley S. F., Jenner G. A., Kjarsgaard B. A. (2005a) Integrating ultramafic lamprophyres into the IUGS classification of igneous rocks: rationale and implications. Journal of Petrology 46:1893–1900.
Tappe S., Foley S. F., Kjarsgaard B. A., Heaman L. M., Jenner G. A., Stracke A., Romer R. L. (2005b) Ultramafic lamprophyres and carbonatites of Labrador and New Quebec: towards a genetic model for Neoproterozoic rift-related alkaline magmatism in the North Atlantic region. Report–Geological Survey of Denmark and Greenland 68:115–117.
Taylor H. P., Frechen J., Degens E. T. (1967) Oxygen and carbon isotope studies of carbonatites from the Laacher See District, west Germany and the Alnö District, Sweden. Geochimica et Cosmochimica Acta 31:407–430.[CrossRef][Web of Science]
Taylor W. R. and Green D. H. (1988) Measurement of reduced peridotite–C–O–H solidus and implications for redox melting of the mantle. Nature 332:349–352.[CrossRef]
Thompson R. N. and Gibson S. A. (1994) Magmatic expression of lithospheric thinning across continental rifts. Tectonophysics 233:41–68.[CrossRef][Web of Science]
Thy P., Stecher O., Korstgard J. A. (1987) Mineral chemistry and crystallization sequences in kimberlite and lamproite dikes from the Sisimiut area, central West Greenland. Lithos 20:391–417.[CrossRef][Web of Science]
Torsvik T. H., Smethurst M. A., Meert J. G., Van der Voo R., McKerrow W. S., Brasier M. D., Sturt B. A., Walderhaug H. J. (1996) Continental break-up and collision in the Neoproterozoic and Palaeozoic: a tale of Baltica and Laurentia. Earth-Science Reviews 40:229–258.
Ulmer P. and Sweeney R. J. (2002) Generation and differentiation of Group II kimberlites: constraints from a high-pressure experimental study to 10 GPa. Geochimica et Cosmochimica Acta 66:2139–2153.[CrossRef][Web of Science]
Upton B. G. J., Emeleus C. H., Heaman L. M., Goodenough K. M., Finch A. A. (2003) Magmatism of the mid-Proterozoic Gardar Province, South Greenland: chronology, petrogenesis and geological setting. Lithos 68:43–65.[CrossRef][Web of Science]
van Achterbergh E., Griffin B. J., Ryan C. G., O'Reilly S. Y., Pearson N. J., Kivi K., Doyle B. J. (2004) Melt inclusions from the deep Slave lithosphere: implications for the origin and evolution of mantle-derived carbonatite and kimberlite. Lithos 76:461–474.[CrossRef][Web of Science]
Veksler I. V., Petibon C. M., Jenner G. A., Dorfman A. M., Dingwell D. B. (1998) Trace element partitioning in immiscible silicate–carbonate liquid systems: an initial experimental study using a centrifuge autoclave. Journal of Petrology 39:2095–2104.[CrossRef][Web of Science]
Vollmer R., Ogden P., Schilling J. G., Kingsley R. H., Waggoner D. G. (1984) Nd and Sr isotopes in ultrapotassic volcanic rocks from the Leucite Hills, Wyoming. Contributions to Mineralogy and Petrology 87:359–368.[CrossRef][Web of Science]
Vuorinen J. H. and Skelton A. D. L. (2004) Origin of silicate minerals in carbonatites from Alnö Island, Sweden: magmatic crystallization or wall rock assimilation? Terra Nova 16:210–215.[CrossRef][Web of Science]
Wardle R. J. and Hall J. (2002) Proterozoic evolution of the northeastern Canadian Shield: Lithoprobe Eastern Canadian Shield Onshore–Offshore Transect (ECSOOT), introduction and summary. Canadian Journal of Earth Sciences 39:563–567.
Waters F. G. (1987) A suggested origin of MARID xenoliths in kimberlites by high pressure crystallization of an ultrapotassic rock such as lamproite. Contributions to Mineralogy and Petrology 95:523–533.[CrossRef][Web of Science]
Wendlandt R. F. and Eggler D. H. (1980) The origins of potassic magmas: 2. Stability of phlogopite in natural spinel lherzolite and in the system KAlSiO4–MgO–SiO2–H2O–CO2 at high pressures and high temperatures. American Journal of Science 280:421–458.
Wyllie P. J. (1980) The origin of kimberlite. Journal of Geophysical Research 85:6902–6910.
Wyllie P. J. and Tuttle O. F. (1960) The system CaO–CO2–H2O and the origin of carbonatites. Journal of Petrology 1:1–46.[Web of Science]
Zhu C. and Sverjensky D. A. (1992) F–Cl–OH partitioning between biotite and apatite. Geochimica et Cosmochimica Acta 56:3435–3467.[CrossRef][Web of Science]
Zindler A. and Hart S. (1986) Chemical geodynamics. Annual Review of Earth and Planetary Sciences 14:493–571.[CrossRef][Web of Science]
CiteULike 
Connotea 
Del.icio.us What's this?
This article has been cited by other articles:
|
|
 |
|
  A. R. Woolley and B. A. Kjarsgaard PARAGENETIC TYPES OF CARBONATITE AS INDICATED BY THE DIVERSITY AND RELATIVE ABUNDANCES OF ASSOCIATED SILICATE ROCKS: EVIDENCE FROM A GLOBAL DATABASE Can Mineral, August 1, 2008; 46(4): 741 - 752. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
A. F. Cooper and L. A. Paterson CARBONATITES FROM A LAMPROPHYRIC DYKE-SWARM, SOUTH WESTLAND, NEW ZEALAND Can Mineral, August 1, 2008; 46(4): 753 - 777. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
T. F.D. Nielsen and K. K. Sand THE MAJUAGAA KIMBERLITE DIKE, MANIITSOQ REGION, WEST GREENLAND: CONSTRAINTS ON AN Mg-RICH SILICOCARBONATITIC MELT COMPOSITION FROM GROUNDMASS MINERALOGY AND BULK COMPOSITIONS Can Mineral, August 1, 2008; 46(4): 1043 - 1061. [Abstract] [Full Text] [PDF]
|
|
|
|
 |
|
 P. L. Roeder and D. J. Schulze Crystallization of Groundmass Spinel in Kimberlite J. Petrology, August 1, 2008; 49(8): 1473 - 1495. [Abstract] [Full Text] [PDF]
|
|
|
|
 |
|
 C. Bonadiman, M. Coltorti, S. Duggen, L. Paludetti, F. Siena, M. F. Thirlwall, and B. G. J. Upton Palaeozoic subduction-related and kimberlite or carbonatite metasomatism in the Scottish lithospheric mantle Geological Society, London, Special Publications, January 1, 2008; 293(1): 303 - 333. [Abstract] [Full Text] [PDF]
|
|
| ||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||