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JOURNAL OF PETROLOGY VOLUME 43 NUMBER 5 PAGES 885–905 2002 Gabbroic Pegmatite Intrusions, Iberia Abyssal Plain, ODP Leg 173, Site 1070: Magmatism during a Transition from Non-volcanic Rifting to Sea-floor Spreading JAMES S. BEARD1∗, PAUL D. FULLAGAR2 AND A. KRISHA SINHA3 1 VIRGINIA MUSEUM OF NATURAL HISTORY, 1001 DOUGLAS AVENUE, MARTINSVILLE, VA 24112, USA 2 DEPARTMENT OF GEOLOGICAL SCIENCES, UNIVERSITY OF NORTH CAROLINA, CHAPEL HILL, NC 27599, USA 3 DEPARTMENT OF GEOLOGICAL SCIENCES, VIRGINIA TECH, BLACKSBURG, VA 24061, USA RECEIVED FEBRUARY 21, 2001; REVISED TYPESCRIPT ACCEPTED NOVEMBER 7, 2001 On the Iberia Abyssal Plain (Ocean Drilling Program Site 1070), gabbroic pegmatites and related rocks (127 ± 4 Ma, U–Pb zircon) intrude upper mantle that was subsequently exposed and serpentinized during Early Cretaceous non-volcanic rifting. The pegmatites include a 3–4 m dike or sill (the ‘main’ pegmatite), numerous dikelets of 1–5 cm thickness, and clasts within the overlying ophicalcite breccia. Exclusive of rodingitization, the main pegmatite contains 40–70% calcic andesine, 25–35% kaersutitic amphibole (Mg# 60–70), 5–25% augite (Mg# 70–80) and 1–2% ilmenite. The dikelets are more magnesian (Mg# up to 82 in kaersutite and 88 in augite). Most indications are that the high Mg#s in the dikelets reflect igneous compositions. Isotopic and elemental chemistry indicate that the pegmatite-forming melt was enriched in incompatible elements relative to normal mid-ocean ridge basalt, but not as enriched as Azores basalts. The amphibole-bearing plagioclase peridotites of the Iberia Abyssal Plain are an appropriate source for the pegmatite melts. A combination of decompression accompanying unroofing and heating from the upwelling asthenosphere beneath the developing rift caused P–T conditions in the amphibole-bearing lithosphere to exceed the dehydration-melting solidus (>1050°C), producing small-volume, enriched, hydrous melts. Pegmatite intrusion pre-dates unroofing at Site 1070 and post-dates syn-rift sedimentation and faulting of serpentinite seen inboard on the Iberia Abyssal Plain. Thus, serpentinization and unroofing were timetransgressive and the age of the non-volcanic sea floor formed by the unroofed mantle grows younger outboard, just as is the case for normal, volcanic sea floor. ∗Corresponding author. E-mail: [email protected] Oxford University Press 2002 KEY WORDS: abyssal plain; gabbro; Iberia; pegmatites; rifting INTRODUCTION AND GEOLOGIC SETTING The west Iberian ocean–continent transition is a welldocumented example of a non-volcanic (a more precise term might be magma-poor) rifted margin. Combined seismic studies and drilling on Ocean Drilling Program (ODP) Legs 103, 149, and 173 have revealed that the oceanic basement beneath much of the Iberia Abyssal Plain (IAP) consists of serpentinized peridotite, representing mantle rocks exhumed during continental rifting, and stranded fault blocks of continental crust (Boillot et al., 1987; Krawczyk et al., 1996; Pickup et al., 1996; Whitmarsh & Sawyer, 1996; Whitmarsh et al., 1998) (Figs 1 and 2). Rifting at the margin began with the thinning of the continental crust, probably in the latest Jurassic (Whitmarsh et al., 2000; Whitmarsh & Wallace, 2001). JOURNAL OF PETROLOGY VOLUME 43 NUMBER 5 MAY 2002 Fig. 1. Location map of the Iberia Abyssal Plain showing Site 1070 and other Ocean Drilling Program drill-hole sites from Legs 149 and 173. Fig. 2. Schematic cross-section based on drill-hole data along Ocean Drilling Program Legs 149 and 173. Hole 1070 is the farthest outboard site drilled to date on the Iberia Abyssal Plain. 886 BEARD et al. GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN Studies of rocks from Sites 900, 1067, and 1068 on a single basement high where Permian continental crust is in fault contact with serpentinized peridotite (Figs 1 and 2) suggest a long period of rifting, dissection by several generations of faults, and tectonic exhumation. At this basement high (‘Hobby High’; Whitmarsh et al., 2000) active tectonic dissection of the continental crust and exhumation of mantle rocks was under way by 137 Ma and continued for 10–20 my (Whitmarsh & Wallace, 2001; Manatschal et al., 2002). Site 1070 lies >100 km outboard of Hobby High, with much of the intervening oceanic basement being serpentinized peridotite (Fig. 2). Linear magnetic anomalies in the vicinity of Site 1070 suggest the presence of ocean crust of 124–127 Ma age (Whitmarsh & Miles, 1995), but no basalt or diabase was found. Instead, drilling revealed pegmatoidal gabbro intrusive into partially serpentinized peridotite. The basement rocks at Site 1070 were exhumed by late Aptian (112–119 Ma) time (Whitmarsh & Wallace, 2001). Site 1070 lies 30 km east of the ‘J’ anomaly, a strong, linear magnetic high found on both sides of the Atlantic (Rabinowitz et al., 1979; Tucholke & Ludwig, 1982; Masson & Miles, 1984; Fig. 1). The ‘J’ anomaly is generally considered to mark an episode of voluminous basaltic magmatism dating from >120 Ma (i.e. between anomalies M0 and M1), just after the eruption of mid-ocean ridge-type basalts (coincident with anomaly M2, >122 Ma) off the Grand Banks (Tucholke & Ludwig, 1982; Vogt & Tucholke, 1989). Most workers (e.g. Tucholke & Ludwig, 1982; Whitmarsh et al., 1998) suggest that ocean crust formation in the Grand Banks–Iberia segment began at >125 Ma (anomaly M4 on the Southern Iberia Abyssal Plain), but, to date, no basalts of this age have been recovered there. Some syn-rift magmatism has been documented at the west Iberian margin, especially at its northern and southern ends. An undeformed gabbro–troctolite–diorite lopolith on the Gorringe Bank, south of the IAP, has yielded U–Pb zircon ages of 136–138 Ma, and younger (122 Ma, U–Pb zircon) gabbros occur near the northern end of the IAP on the Galicia Bank (Scharer et al., 2000; Fig. 1). However, very few unequivocal syn-rift igneous rocks have been recovered during drilling on the Iberia Abyssal Plain proper. In particular, metamorphosed gabbroic rocks drilled at Sites 900, 1067, and 1068 on Legs 149 and 173 have now been shown to be Paleozoic in age (Rubenach, 1999) and not syn-rift gabbro as originally thought (e.g. Feraud et al., 1996). Leg 149 recovered extremely weathered basalt and diabase clasts from serpentinite breccias in Hole 899B. These rocks have chemical characteristics similar to enriched mid-ocean ridge basalt (EMORB) and probably reflect syn-rift sea-floor magmatism (Seifert et al., 1997). However, the age and source of these rocks remains unknown, and their exceptionally severe and pervasive alteration makes their further study (as igneous rocks) problematic. Basement at Site 1070 consists of a calcite-cemented ‘jigsaw’ breccia reminiscent of Alpine ophicalcites (e.g. Bernoulli & Weissert, 1985) and apparently formed moreor-less in place during low-temperature alteration and serpentinization (Skelton & Valley, 2000). This is immediately underlain by a thick (3–4 m recovered thickness over a 10 m interval), relatively fresh, gabbro pegmatite dike intruding the underlying serpentinite (Whitmarsh et al., 1998). Beneath this is relatively massive, partially (>70–95%) serpentinized peridotite. In addition to the large dike, pegmatoidal gabbro also occurs as clasts in the upper breccia and as strongly rodingitized dikelets in the underlying serpentinite. Apart from serpentinite, the only other rock type found in the Hole 1070A basement is albitite, which, like the pegmatoidal gabbro, occurs as small dikes in serpentinite and clasts in the upper breccia. The gabbro pegmatite yields Ar–Ar cooling ages of 119–121 Ma on igneous amphibole (kaersutite) (Turrin, 1999; Manatschal et al., 2002). We report here a new U–Pb (ion probe on zircon from associated albitites) age of 127 ± 4 Ma, which we interpret as an igneous age. The Site 1070 gabbro pegmatites are the only igneous rocks that are demonstrably coeval with rifting in the southern IAP. As such, they help to document the transition from non-volcanic margin to normal ocean crust formation in this part of the North Atlantic Basin. Because the pegmatites represent magmas formed at a time and place close to the onset of true mid-ocean ridge magmatism, they can also provide information on the processes involved in the inception of magmatism in oceanic rift systems. METHODS 887 The major element chemistry of all minerals was determined by electron microprobe in the Department of Geological Sciences at Virginia Tech. Separates for trace element and isotopic analysis were prepared by hand selecting pegmatite fragments rich in a particular mineral, coarsely crushing the fragment and hand picking a rough mineral separate. Green and zoned amphibole fragments were excluded from the rough separate. The separates were crushed to 100 mesh in an agate mortar under acetone, run through a magnetic separation and inspected for obvious impurities. Feldspar separates were then leached in dilute HF for 30 min and rinsed in deionized water. One split of all separates was analyzed by Actlabs using inductively coupled plasma (ICP) (for most transition metals) and ICP-mass spectrometry (ICP-MS) for all other elements. JOURNAL OF PETROLOGY VOLUME 43 A second split was analyzed for Sr, Nd, and Pb isotopes. Isotopic tracer solutions (either 150Nd, 147Sm, 87Rb, and 84 Sr spikes, or a mixed 205Pb/233–236U spike) were added to samples before dissolution. Samples were dissolved in Savillex vials using HF and HNO3. Except for plagioclase samples, AG50Wx8 cation exchange resin was used to separate Rb, Sr, and bulk rare earth element (REE) fractions from solutions using HCl. The bulk REE fraction was adsorbed onto AG50Wx4 resin equilibrated with alpha-HIBA, and Nd and Sm were selectively eluted with alpha-HIBA. The Sr fraction was purified using EiChrom SrSpec resin. Separation of Pb and U from plagioclase samples employed standard HBr and HCl anion exchange chemistry. The mass spectrometer used for the analyses is a VG (Micromass) Sector 54 thermal ionization mass spectrometer in the Department of Geological Sciences, University of North Carolina at Chapel Hill (UNC-CH). Nd was measured as the metal species using a dynamic multicollector mode. Samarium was analyzed in static multicollector mode. Mass fractionation was corrected using 146Nd/144Nd = 0·7219 and 149Sm/152Sm = 0·51685. Recent results for standard isotopic reference materials: La Jolla Nd standard, 143Nd/144Nd = 0·511855; Ames Nd metal, 143Nd/144Nd = 0·512140; BCR-1, 143Nd/144Nd = 0·512631. Sr was analyzed in dynamic multicollector mode. 87Sr/86Sr ratios were corrected for mass fractionation using an exponential mass fractionation law. All 87Sr/86Sr ratios from the UNC-CH laboratory are reported relative to a value of 0·710250 for this NBS 987 standard (typically, four standards were analyzed with each turret of samples; if the 87Sr/86Sr ratio for the four standards averaged 0·710225, a value of 0·000025 was added to the ratio for each sample from that turret). Internal precision (standard error) for Sr analyses is typically 0·000006–0·000010. Isotopic compositions of Pb and U are measured in multicollector static mode. Uncertainties (2) for measured Pb isotopic ratios are typically below 0·1%. Recent replicate analyses of SRM 981 Pb yield 206Pb/204Pb = 16·897 ± 9, 207Pb/ 206 Pb = 0·91352 ± 15, and 208Pb/206Pb = 2·1617 ± 9. Typical total procedural blanks are insignificant. U and Pb isotopic abundances in zircon (Table 1) were measured using the Cameca IMS 1270 ion microprobe in the Department of Earth and Space Sciences at UCLA following the method of Schumacher et al. (1994). Because of reasonable counting statistics associated with 206 Pb (as opposed to 207Pb), we base our age estimates on 206 Pb/238U. ZIRCON U–Pb AGES A large (1 mm) zircon was hand picked from a partially altered albitite clast sampled from the autochthonous NUMBER 5 MAY 2002 breccia above the main pegmatite at Site 1070 (Whitmarsh et al., 1998). The zircon was hosted by a large (>1 cm) single grain of largely unaltered sodic plagioclase. Similar zircons have been found in several other albitite clasts and dikes. Six spot analyses were attempted on the zircon; four of these yielded statistically significant results. Three of these spots yielded 206Pb/238U ages of 124·1, 126·8, and 130·2 Ma (all ±7 Ma, Table 1); the fourth yielded an age of 112 Ma. The three older dates (average 127 Ma, range consistent with all analyses, 123–131 Ma) we interpret as reflecting the igneous age of the gabbroic pegmatite. This age is similar to the ocean crust age predicted from magnetic anomalies in the vicinity of Site 1070 (Whitmarsh & Wallace, 2001) and also intermediate between the Gorringe Bank (137 Ma) and Galicia Bank (122 Ma) gabbro ages. This is consistent with the northto-south progression that has been proposed for the onset of magmatism at the west Iberian margin (Scharer et al., 2000). We interpret the 112 Ma age as reflecting lead loss, possibly incurred during serpentinization close to the time of exhumation (e.g. Whitmarsh & Wallace, 2001). PETROGRAPHY AND MINERAL CHEMISTRY Overall The Site 1070 gabbro pegmatites include a large (recovered thickness 3–4 m) dike or sill (the ‘main’ pegmatite), numerous small (1–5 cm thick) dikelets intrusive into the serpentinite basement, and fragments within the overlying parautochthonous serpentinite breccia. Exclusive of alteration products, the main pegmatite consists of 40–70% plagioclase, 25–35% red–brown amphibole, 5–25% augite, and 1–2% ilmenite (Whitmarsh et al., 1998). Overall, the main pegmatite has igneous textures resembling an extremely coarse-grained (grain sizes up to 7 cm) amphibole gabbro. Conspicuous red–brown amphibole poikolitically encloses equant to tabular plagioclase and replaces augite. Rodingite-like (e.g. zoisite, prehnite, analcime) alteration is widespread. In the main pegmatite, rodingite zones and patches are evident in all thin sections and hand samples, with plagioclase being particularly strongly affected. In the clasts, and especially the thin dikelets, rodingitization may be pervasive or spare only remnant pyroxene, kaersutite, and ilmenite. Fresh plagioclase was not found in any of the small gabbroic dikelets. In addition to low-temperature alteration, the main pegmatite is locally deformed in discrete, high-temperature, shear zones. Within the shear zones, igneous composition plagioclase and amphibole 888 BEARD et al. GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN Table 1: U–Pb ion probe data from ODP Site 1070 zircon [ODP sample 1070A-8R4, 30–33 cm, piece 3 (albitite)] Spot no. 207 Pb/235U Abs. SE 206 Pb/238U 206 Abs. SE Pb/238U Error 207 Pb/235U Error age (Ma) (Ma) age (Ma) (Ma) 119·9 1 0·1124 0·1314 0·0204 0·001202 130·2 7·591 108·2 2 0·1393 0·07255 0·01987 0·001104 126·8 6·979 132·4 3 –0·04555 0·1274 0·01765 0·001199 112·8 7·596 –47·34 4 0·1732 0·106 0·01944 0·001059 124·1 6·695 162·2 64·66 135·6 91·71 SE, standard error. (i.e. andesine–labradorite and kaersutite) were recrystallized, suggesting deformation occurred at or near granulite-facies conditions (Whitmarsh et al., 1998). In addition to the gabbroic rocks, several albitite breccia clasts and dike fragments were recovered. The albitites consist of 95% coarsely crystalline, mildly to pervasively altered, weakly zoned sodic plagioclase with minor altered biotite, altered amphibole (tremolite), and accessory zircon. They contain no quartz. within large amphibole crystals. Within the remnants, patchy replacement of augite by amphibole along fracture or cleavage planes is common. Crystallographically oriented ilmenite plates, presumably formed during subsolidus exsolution, occur in many augite grains. The augite is moderately aluminous (2–3% Al2O3) and moderately titanian (0·5–1% TiO2) (Table 2). In the main pegmatite, augite has Mg# ranging from 72 to 77, whereas that in breccia clasts and small pegmatite dikelets is substantially more magnesian (Mg# 82–88) (Table 2, Fig. 3). Feldspar Within the main pegmatite body, plagioclase occurs as large (up to 5 cm), weakly zoned, euhedral to subhedral crystals. Alteration ranges from minor recrystallization along cleavage planes to complete replacement. Alteration minerals include prehnite, zoisite, zeolites(?), and analcime, but not albite. Potassium feldspar (Or98, rare) is probably an alteration phase, but late crystallization as an igneous phase cannot be ruled out. Plagioclase in breccia clasts and in the smaller gabbroic dikelets is usually pervasively altered. In the main pegmatite, plagioclase compositions (average of all analyses; An46·5Ab52·2Or1·3) range from An28 to An53, with most rims An35–47 and most cores An43–51. The orthoclase component of the plagioclase typically ranges from 1 to 2 mol % (Table 2). Plagioclase in the albitites is coarse-grained (commonly >1 cm), albite–sodic oligoclase (An6–12; Table 2). The sodic plagioclase in the albitites appears to be somewhat more resistant to low-temperature alteration than the calcic plagioclase in the gabbros. The only fresh plagioclase preserved in the breccia clasts and dikelets is in the albitites. Amphibole Bright red–brown kaersutite (pargasitic hornblende with up to 5·3 wt % TiO2; Fig. 4, Table 2) occurs as thick mantles about augite and as large (up to 7 cm) anhedral, interstitial to poikoblastic crystals enclosing all other igneous phases. The kaersutites are F-bearing hydroxylamphiboles (Table 2). Many amphibole grains are weakly zoned with Al- and Ti-poor, magnesian hornblende at grain boundaries and along fractures. This type of zoning probably represents subsolidus re-equilibration. A few grains, however, are spectacularly zoned from reddish kaersutite cores to deep green (sometimes nearly black), Na-, Fe-, Al-rich, Ti-poor hastingsite or ferropargasite rims (Fig. 5, Table 2). This zoning, also characterized by a drop in Mg# from 65 in the kaersutite core to 29 at the rim, appears to be an igneous feature. Amphibole in the smaller dikes and clasts is more magnesian than that in the main pegmatite (Mg# up to 82). Unlike secondary amphiboles that are clearly related to alteration, however, the amphibole in the clasts and dikes retains its Ti-rich and pargasitic composition (Fig. 4). Clinopyroxene Ilmenite Pinkish brown clinopyroxene (augite) is an early crystallizing phase that is partially to wholly replaced by late igneous amphibole. It occurs as irregular, remnant cores Rounded grains of ilmenite [il/(il + hm) = 88–92] up to 1 cm in diameter occur in all samples, commonly as inclusions in amphibole, augite, and plagioclase. The 889 JOURNAL OF PETROLOGY VOLUME 43 NUMBER 5 MAY 2002 Table 2: Major-element analyses of minerals Amphibole Clinopyroxene Main pegmatite dikelet albitite main pegmatite dikelet 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 9R-2 9R-2 9R-2 9R-1 9R-1 14R-3 8R-4 9R-2 9R-2 9R-1 14R-3 22 cm 22 cm 64 cm 75 cm 82 cm 85 cm 30 cm 22 cm 64 cm 45 cm 85 cm Fe-parg meta. meta. SiO2 42·72 39·80 42·81 41·65 49·54 44·87 53·29 53·52 52·24 53·05 53·31 TiO2 5·19 0·27 4·64 4·28 0·32 4·30 1·26 0·37 0·63 0·63 0·87 Al2O3 11·87 15·59 11·24 11·78 6·82 11·40 3·85 1·21 2·25 2·35 2·70 Cr2O3 0·04 0·03 0·06 0·05 0·07 0·20 0·08 0·01 0·04 0·04 0·17 MgO 12·72 5·78 12·41 13·47 15·42 16·18 21·37 14·86 14·10 14·36 16·50 CaO 11·17 11·68 11·38 10·95 12·63 11·78 11·39 21·13 21·62 21·36 21·74 MnO 0·24 0·11 0·23 0·29 0·17 0·16 0·04 0·28 0·33 0·32 0·19 FeO 11·62 21·99 12·89 12·11 11·14 6·77 3·67 8·93 8·31 8·71 4·69 Na2O 2·94 3·13 2·77 2·71 1·88 2·70 1·77 0·48 0·57 0·55 0·66 K2O 0·58 0·25 0·62 0·63 0·10 0·64 0·09 0·00 0·02 0·00 0·04 F 0·43 0·00 0·32 0·61 0·20 0·09 0·00 Total 99·53 98·63 99·37 98·51 98·29 99·08 96·81 100·78 100·10 101·36 100·85 Mg# 66·1 31·9 63·2 66·5 71·2 81·0 91·2 En 42·3 41·0 41·4 47·3 Fe2O3 A-site 0·67 0·89 0·68 0·63 0·49 0·65 0·21 Fs 14·3 13·6 14·1 7·6 AlIV 1·80 1·91 1·74 1·85 0·86 1·64 0·64 Wo 43·3 45·3 44·4 45·0 Ilmenite Plagioclase main pegmatite dikelet main pegmatite albitite 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 9R-2 9R-2 9R-1 14R-3 9R-2 9R-2 9R-2 9R-2 9R-2 9R-1 8R-4 22 cm 64 cm 45 cm 85 cm 64 cm SiO2 0·05 0·02 0·02 0·06 TiO2 49·01 47·82 48·75 53·69 Al2O3 0·09 0·11 0·14 0·04 Cr2O3 0·05 0·05 0·06 0·14 MgO 2·30 3·64 3·54 6·88 22 cm 22 cm 64 cm 22 cm rim core rim core 75 cm 30 cm calcic sodic 58·52 56·59 57·53 56·04 55·25 58·60 66·87 26·75 27·95 27·48 28·33 28·50 25·90 21·42 0·03 0·04 0·02 0·02 0·02 0·04 0·03 8·61 10·14 9·46 10·44 11·06 7·87 2·00 0·26 0·32 0·18 0·27 0·24 0·13 0·19 Na2O 6·57 5·75 6·21 5·63 5·22 6·95 9·81 K2O 0·31 0·29 0·27 0·23 0·19 0·13 0·26 CaO MnO 0·67 0·63 0·59 0·76 FeO 39·35 35·95 36·95 35·32 8·54 10·92 10·64 3·01 Fe2O3 F Total 100·06 99·14 100·70 101·04 101·07 101·13 100·95 100·48 99·62 100·58 81·9 74·8 75·8 70·9 An 41·3 48·6 45·0 49·9 53·3 38·2 10·0 Hem 8·0 10·2 9·8 2·7 Ab 57·0 49·8 53·5 48·8 45·6 61·1 88·5 Gk 8·4 13·4 12·9 24·5 Or 1·8 1·6 1·5 1·3 1·1 0·8 1·5 Ilm 99·90 Fe-parg, Fe-pargasite; meta., metamorphic. 890 BEARD et al. GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN Fig. 3. Mg# vs Ti (+) and Al (Ε) in clinopyroxene. Plotted points are average analyses for individual samples. The fields outline the range of spot analyses. It should be noted that the main pegmatite is substantially less magnesian than the clast and dike samples. Fig. 4. Mg# vs Ti (+), Na (stars), and AlIV (Ε) in igneous amphibole (kaersutite). Plotted points are average analyses for individual samples. The fields outline the range of spot analyses. As with pyroxene, the clasts and dike are more magnesian than the main pegmatite. It should be noted that even the most magnesian amphiboles are still kaersutite or titanian pargasite. ilmenite is magnesian (geikelite component up to 14% in the main pegmatite and 25% in one small dikelet) and is usually zoned in Mg. There are two sets of very fine (width usually <5 m) exsolution lamellae. One set is hematite-rich (approximately hm25) ilmenite. The other, much more sparsely developed, is baddeleyite (ZrO2). 891 JOURNAL OF PETROLOGY VOLUME 43 NUMBER 5 MAY 2002 Fig. 5. Mg#, Ti, AlIV and A-site occupancy variation with distance across the rim of a zoned amphibole (kaersutite core, hastingsite and ferropargasite rim) from the main pegmatite. This type of zoning, although optically and compositionally spectacular, is uncommon and has been seen only in the main pegmatite. Other minerals Zircon occurs as rather large (up to 1 mm) grains in the albite rocks. A strongly chloritized phyllosilicate that may have once been biotite occurs in these rocks and in a single sample from the main pegmatite. Sphene occurs sporadically as epitaxial growths on ilmenite and is a late magmatic or subsolidus phase. Trace element chemistry of mineral separates Plagioclase contains 200–300 ppm Ba and 800–1300 ppm Sr, with Sr/Rb averaging >250. Despite acid leaching, K in two out of three bulk analyses is higher than the averaged microprobe data (Tables 2 and 3), whereas Rb in a small (10 mg) split analyzed by isotope dilution is five times (31 vs 6 ppm) that of a larger (100 mg) sample analyzed for Rb by ICP. This suggests that, despite leaching, a K- and Rb-rich component, probably K-feldspar, was present in at least one and possibly two of the separates. Concentrations of most other trace elements are low, especially compared with amphibole. Plagioclase is light REE (LREE) enriched with a large positive Eu anomaly and very low abundances of heavy REE (HREE) (Fig. 6). Amphibole contains substantial Nb, transition metals and REE, and is probably the major reservoir for most of these elements in the pegmatite. Both LREE and (to a lesser extent) HREE are depleted with respect to middle REE (MREE). There is a small negative Eu anomaly (Fig. 6). The three amphibole separates yield exceptionally consistent results for all trace elements for both ICP and isotope dilution, although Rb from isotope dilution (2–3 ppm) is slightly lower than the value (4 ppm) from ICP. Both K and Ti from ICP are consistent with microprobe results (Tables 2 and 3). The average of the three analyses is taken as a realistic estimate of the average kaersutite composition in the main pegmatite. The single augite separate has element distribution patterns (e.g. chondrite-normalized REE) similar to the amphibole, but much lower abundances for most elements. An exception to this is compatible transition metals, which are as abundant or (in the case of Cr) more abundant in augite than in amphibole. A problem with the augite separate is its likely contamination by very fine included septa and stringers of replacement kaersutite. Both K and Ti measured in the separate are higher than the averaged microprobe analyses, enough so to suggest 10–15% amphibole contamination. High-K and high-Ti augite analyses were screened and eliminated during compilation of the microprobe average as a check on amphibole contamination. However, the high Ti in the analysis of the mineral separate may reflect, in part, ilmenite exsolution rather than amphibole contamination. The single ilmenite separate has low abundances of most trace elements with several notable exceptions. The high concentrations of large ion lithophile elements 892 BEARD et al. GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN Table 3: Trace element analyses of mineral separates (values in ppm) Sample: 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 9R-3 9R-2 9R-1 9R-2 9R-2 9R-1 9R-2 9R-1 20 cm 28 cm 75 cm 64 cm 28 cm 75 cm 64 cm 45 cm, Mineral: plag plag plag amph amph amph cpx ilm Sr 1286 928 802 166 148 142 35 12 K 8050 7139 4068 3984 4233 4233 664 913 Rb 6 4 3 4 4 4 2 3 Ba 212 287 230 193 161 170 116 92 Cs 0·2 0·1 0·1 0·1 Th 0·1 0·0 0·1 0·39 0·35 0·34 0·3 U 0·03 0·02 0·04 0·14 0·09 0·11 0·18 Nb 27·8 Ta P Zr 26·9 1·94 1·97 27·3 2·01 2·2 0·1 130 130 130 218 262 699 130 9 10 10 90 109 108 57 Hf 0·2 Ti 700 840 Y 2 2 Ge 2·8 2·2 Ga 0·1 16 17 4·1 846 1·1 18 4·1 3·9 2·2 0·34 0·11 54 3·48 197 2·3 27671 29981 28665 6127 297843 83 90 86 47 5 4·4 4·5 16 17 47·6 3·9 16 53·2 5·3 2·6 7 4 Co 53 Ni 160 146 117 114 V 910 996 977 557 2085 Cr 142 154 106 233 230 109 108 108 140 Sc 0·7 1·1 0·7 La 1·96 2·08 2·70 Ce 3·61 3·83 4·54 Pr 0·40 0·41 0·47 Nd 1·65 1·73 1·75 34·5 35·4 34·4 Sm 0·25 0·26 0·25 12·1 12·3 11·9 Eu 0·97 1·15 1·32 Gd 0·23 0·23 0·22 8·79 7·67 28·6 28·5 5·45 5·53 4·01 3·97 15·7 15·7 Tb 0·03 0·03 0·03 Dy 0·13 0·13 0·11 Ho 0·03 0·02 0·02 3·34 3·39 Er 0·07 0·06 0·04 8·9 9·01 1·15 Tm 2·88 2·88 16·5 7·61 27·4 5·31 3·78 14·9 4·99 108 56·3 4·14 12·7 4·6 2·3 0·45 14·5 5·36 1·49 0·3 1·74 0·13 7 0·34 1·32 0·06 7·88 0·37 3·17 1·59 0·08 8·45 4·45 0·25 1·19 1·12 0·60 0·05 16·7 2·73 45·7 15·8 Yb 0·06 0·05 0·03 6·75 7·06 6·58 3·67 0·46 Lu 0·008 0·007 0·004 0·893 0·923 0·862 0·508 0·101 (LILE), especially K and Ba, undoubtedly reflect inclusions of a K-rich phase, probably orthoclase replacing plagioclase. Such inclusions are exceptionally difficult to detect in an opaque mineral. A large (1 g) sample processed for isotopic analysis yielded (by isotope dilution) much lower Rb (0·05 vs 3 ppm) and Sr (3·5 vs 12ppm) than the smaller (100 mg) separate used for ICP, consistent with feldspar contamination of the latter split. The high Zr content of the ilmenite is consistent with observed baddeleyite exsolution lamellae. High concentrations of Nb and transition metals, especially V, are not unexpected. REE concentrations are low, and the 893 JOURNAL OF PETROLOGY VOLUME 43 NUMBER 5 MAY 2002 Fig. 6. Chondrite-normalized (Sun & McDonough, 1989) rare earth elements in mineral separates from the main pegmatite. chondrite-normalized pattern probably reflects mixing of an ilmenite signal (possibly HREE-rich) with feldspar and/or other contaminants (Fig. 6). MODELS AND ESTIMATES OF MELT COMPOSITION We have estimated the whole-rock chemistry of the 1070 gabbro pegmatites by two methods: modal mass balance and partition coefficient inversion. Modal models are end-member calculations that tacitly assume that the minerals in the rock are not cumulate, but reflect a bulk liquid composition. Partition coefficient modeling provides another end-member by assuming that the minerals are entirely cumulate. Taken together, the two end-member models provide some constraints on likely liquid compositions. Isotope geochemistry Sr, Nd, and Pb isotopic compositions were determined on splits of the mineral separates that were analyzed for trace elements. Nd and Sr were measured in amphibole and pyroxene; Sr and Pb in plagioclase. Results are given in Table 4. Nd values for pyroxene and amphibole cluster in a tight group with initial (125 Ma) Nd of 4·7–4·9 (modern values 5·0–5·3). For amphibole, pyroxene and plagioclase, initial 87Sr/86Sr ranges from 0·70362 to 0·70463. Two of the amphiboles have initial 87Sr/86Sr <0·70366; all other analyses yielded initial 87Sr/86Sr >0·704. The feldspar has average 206Pb/204Pb, 207Pb/ 204 Pb, and 208Pb/204Pb, respectively, of 18·65; 15·48, and 38·3 (Table 4). Nd and Sr were also analyzed in the ilmenite separate, which has the lowest concentrations of Sr, Rb, Nd and Sm, and the highest initial 87Sr/86Sr (0·70623) and lowest Nd (4·4) of any sample analyzed for isotopes. Given the low concentrations of REE in the ilmenite (Table 3), Nd probably lies within measurement error of the other separates. The anomalously radiogenic ilmenite Sr presumably reflects contamination, probably with feldspar alteration products (Fig. 7). Modal estimates of major and trace element composition The coarse-grained nature of the pegmatite and the necessarily small size of the samples precludes meaningful determination of a thin-section-based mode. The modal estimates used here to calculate bulk-rock chemistry rely heavily upon shipboard modal data. These modes are based on visual inspection of the whole core and on a semiquantitative scanned image analysis (Whitmarsh et al., 1998). The volume modes are then converted to mass (wt %) modes for the whole-rock calculation. The general relationship used for calculating wholerock chemistry from modal data in these rocks is 894 BEARD et al. GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN Table 4: Isotopic compositions of mineral separates Sample: Mineral: 87 Sr/86Sr 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 1070a- 9R-2 9R-2 9R-1 9R-2 9R-3 9R-1 9R-2 9R-1 28 cm 64 cm 75 cm 28 cm 20 cm 75 cm 64 cm 45 cm amph amph amph plag plag plag cpx ilm 0·703656 0·704273 0·703618 0·512909 0·512911 0·512893 0·704322 0·704625 0·704005 143 Nd/144Nd 206 Pb/204Pb 18·63 18·73 18·57 207 Pb/204Pb 15·50 15·51 15·45 208 Pb/204Pb 38·30 38·39 38·19 0·704461 0·706227 0·512912 0·512867 Sr and Nd are initial ratios calculated at 127 Ma. Pb ratios are uncorrected. Fig. 7. (a) Ciwr = Xpl × Cipl + Xcpx × Cicpx + Xamp × Ciamp + Xilm × Ciilm mode. The third mode is also mafic, but more amphibole rich. where X is the mass fraction of the phase (plagioclase, clinopyroxene, amphibole, ilmenite) and Ci is the concentration of a given element i. The accuracy of the calculation is limited by uncertainties in the mode, by analytical error, and, in the case of trace elements, by the purity of the separates. Three calculated whole-rock compositions based on modal data are presented (Table 5). One is based on a plagioclase-rich mode, the second, on a more mafic Trace element estimates based on partition coefficients 895 Any attempt to reconstruct melt compositions by inverting partition coefficients will involve large errors. We have attempted to manage the error in our calculations in several ways. First, we chose a wide range of possible partition coefficients, based on experimental studies of JOURNAL OF PETROLOGY VOLUME 43 NUMBER 5 MAY 2002 Fig. 7. (a) Covariation of Sr and Nd isotopes. MAR, Mid-Atlantic Ridge; OFZ, Oceanographer Fracture Zone. (b) Covariation of Pb isotopes. MORB fields from Staudigel et al. (1984) and Wilson (1989); OFZ data from Shirey et al. (1987); Azores data from Turner et al. (1997). Results of modeling basalts or ion probe phenocryst–matrix studies for all four major mineral phases (Nielsen, 2000, GERM database). We complemented this with data from one widely accepted compilation (Rollinson, 1993). Second, we used the partition coefficients to frame a range of potential melt compositions, rather than try to select a single ‘best’ partition coefficient value for any given element. Third, we interpret this range largely on the basis of relative, rather than absolute, element abundances. Model melt compositions are given in Table 6. The partition coefficients and primary data sources are given in the Appendix. 896 All modal models yield broadly basaltic major element compositions. Absolute and relative quantities of most elements vary largely as a function of chosen mineral mode. TiO2 ranges from 2·2 to 3·4% and exceeds 1·5% in most models even when ilmenite is (unjustifiably) excluded from the mode. Na2O is nearly 3·2% even in the most plagioclase-poor model. In the modal models, most trace element abundances are controlled by amphibole. In model 3, for example, 70–85% of Nb, Ta, Y and REE (except La and Eu) reside in amphibole. Partly in consequence of this, the BEARD et al. GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN Table 5: Modal estimates of bulk composition Table 6: Models of trace-element composition for bulk rock %Plag:∗ 70 40 45 Low Kd High Kd Modal %Cpx: 4 30 18 model model model 3 %Amph: 25 28 35 %Ilm: 1 2 2 Sr 1875 111 406 K 12694 5101 2713 SiO2 51·85 49·71 49·05 Rb 80 TiO2 2·17 3·11 3·42 Ba 1519 4·3 162 2·74 188 Al2O3 21·50 13·89 15·84 Cs MgO 4·32 8.l5 7·69 Th 18 0·35 0·23 CaO 10·42 13·48 12·12 Ta 25 1·29 0·86 MnO 0·09 0·18 0·16 Nb 342 23 FeO 4·76 7·93 7·82 Ce 317 34 4·7 Na2O 4·80 3·17 3·61 Zr 814 66 59·2 K2O 0·34 0·27 0·32 Hf 18 2·6 F 0·09 0·09 0·12 Sm 33 6·7 Total 100·32 100·33 100·16 Mg# 61·77 65·66 63·66 Modes are given in volume percent, but composition calculated based on weight percent. calculated whole-rock REE have a concave-down pattern that parallels amphibole (Figs 6 and 8). In this model plagioclase accounts for 86% of the Sr, 40–60% of most other LILE, and >20% of the La and Eu. The model based on the highest partition coefficients (i.e. those yielding the lowest abundances of a given element in the melt) resembles the modal model, except that it predicts slight LREE enrichment (Fig. 8) and yields lower values for compatible elements. This similarity is not surprising, as both models would be expected to yield near minimum values. The model based on the lowest partition coefficients predicts very high concentrations, especially for incompatible elements. DISCUSSION Comparisons with Iberian margin and other oceanic rocks The isotopes and modeled bulk compositions of the Site 1070 gabbro pegmatites are similar to EMORB in most respects. The pegmatites are generally less enriched than Azores basalts and, even allowing for fractionation, more enriched than most normal MORB (NMORB). The variation in Sr isotopes in the Site 1070 pegmatite mineral separates is consistent with contamination with Cretaceous seawater ( 87Sr/86Sr = 0·70725; Veizer et al., 1999; Fig. 7a). There is little variation in Nd isotopic 1·90 1·90 0·11 12·3 2·1 5·6 Ti 33072 9281 21629 Y 216 52 42 Yb 14 4·14 Sc 173 26 69 Cr 137 7 102 Ni 76 8 74 La 166 15 Pr Nd 3·3 4·98 2·65 121 26 16·47 Eu 12·7 2·5 Gd 25 7·7 7·19 2·02 1·32 8·1 7·66 1·1 1·55 Tb Dy Ho Er 3·14 26 7·5 16 5·1 Tm Lu 2·26 4·16 0·55 2·8 0·57 0·44 Kd models based on partition coefficients given in Table A1 (Appendix). Modal model is model 3 from Table 5. composition. In an Sr–Nd isotopic diagram, the samples with the least radiogenic Sr plot at the juncture of overlap amongst (non-plume?) EMORB from the Oceanographer Fracture Zone (Shirey et al., 1987), basalts from the Azores (Turner et al., 1997), and the least depleted Atlantic NMORB (Staudigel et al., 1984) (Fig. 7a). Feldspar Pb isotopes are similar both to NMORB and to EMORB from the Oceanographer Fracture Zone. They are less radiogenic than basalts from the Azores (Fig. 7b). Analyses of metagabbro from the Iberia Abyssal Plain yield Nd ranging from +6 to +10 (Seifert et al., 1997; Cornen et al., 1999). However, these rocks are now 897 JOURNAL OF PETROLOGY VOLUME 43 NUMBER 5 MAY 2002 Fig. 8. Chondrite-normalized (Sun & McDonough, 1989) REE for the model melt compositions. recognized as Paleozoic continental basement (Rubenach, 1999). On the Galicia Bank, basaltic and gabbroic rocks have Nd of +2·2 to +8·8, with older rocks (e.g. 122 Ma) having less depleted values than younger (e.g. 100 Ma) rocks (Charpentier et al., 1998). This trend was interpreted by Charpentier et al. (1998) as reflecting a decreasing lithospheric and increasing asthenospheric component with time. Young (60–66 Ma) alkaline basalts from the Gorringe Bank have initial 87Sr/86Sr of 0·7031 and Nd of +6·6 (Bernard-Griffith et al., 1997). A dolerite from the west end of Gorringe Bank has Nd of +5·1 and initial 87Sr/86Sr of 0·70484 (Cornen et al., 1999). Scharer et al. (2000) reported strongly depleted Hf of +19·5 to +20·5 for 135 Ma gabbroic rocks from Gorringe Bank and slightly less depleted values (+14·0 to +14·6) for 122 Ma gabbro on the Galicia Bank. Despite the uncertainties in the bulk compositional models, several general characteristics of the magma from which the pegmatites formed can be recognized (Figs 8 and 9): (1) the model based on the minimum partition coefficients (i.e. the model yielding the highest concentrations of the modeled elements) appears to be largely unrealistic (predicting, for example, 340 ppm Nb). However, abundance patterns in all three models are roughly parallel for most elements. (2) None of the models exhibit the strong LREE depletion typical of NMORB. As noted above, however, chondrite-normalized La/Yb is higher in the partition coefficient models (3–10) than in the modal models (about unity) (Fig. 8). (3) All models are enriched in the most incompatible elements with respect to the NMORB of Sun & McDonough (1989). For many strongly incompatible elements, this enrichment is greater than can be accommodated by any realistic (e.g. 75%) degree of fractionation (Fig. 9). (4) The modal and high-Kd models resemble the EMORB of Sun & McDonough (1989). (5) The least radiogenic (i.e. excluding lavas from São Miguel) Azores samples of Turner et al. (1997) are mostly bracketed by the models. (6) Nb and Ta are not depleted in the models. (7) Yb and Y are unique in that they are higher in all models than in NMORB, EMORB, or Azores lavas (Fig. 9). Highly altered basalts, of unknown age, but interpreted as syn-rift by Seifert et al. (1997), were recovered from ODP Site 899, >40 km east of Site 1070. Some of these basalts are Ti rich (>3% TiO2) and contain Ti-augite, kaersutite and Ti-rich biotite (Cornen et al., 1999). Most of the basalts are LREE enriched and have been interpreted as coming from a somewhat enriched mantle source (Seifert et al., 1997; Cornen et al., 1999). Gorringe Bank includes a younger series of 60–77 Ma alkaline basalts, diorites, and gabbros, and an older series (137 Ma) of cumulate gabbros and dioritic, doleritic and basaltic differentiates. These latter rocks have flat to slightly LREE-enriched REE patterns, interpreted by Cornen et al. (1999) as transitional MORB ranging to a mixed MORB–OIB (ocean-island basalt) source. 898 BEARD et al. GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN Fig. 9. Modal model number 3 (Table 5; Ε), average Azores basalt (Φ; Turner et al. (1997), excluding São Miguel) and partition coefficient models (shaded field) normalized to (a) NMORB and (b) EMORB of Sun & McDonough (1989). Many mafic igneous rocks, including gabbro, diorite, and basalt, have been recovered from the Galicia Bank. These include sheared kaersutite–plagioclase diorite dikes with coarse (greater than centimeter-sized) relict kaersutite porphyroclasts and accessory biotite and ilmenite (Beslier et al., 1990). Many Galicia Bank basaltic rocks also contain accessory brown amphibole and biotite and show mild enrichment in LREE (Charpentier et al., 1998; Cornen et al., 1999). In summary, both elemental and isotopic chemistry and mineralogy of the Site 1070 gabbros are consistent with derivation from an enriched mantle source, possibly subcontinental lithosphere, and mitigate against a depleted, NMORB-type asthenospheric source. The pegmatites have mineralogical and chemical similarities to many other syn-rift intrusive and extrusive rocks found elsewhere in the region. In the continuum of lithospheric vs asthenospheric magmas recognized at the west Iberian margin (Charpentier et al., 1998; Cornen et al., 1999), the Site 1070 rocks lie near the lithospheric end-member. Origin of high Mg in gabbroic clasts and dikelets 899 The interpretation of the origin of the high-Mg pegmatitic dikes and clasts bears on the overall interpretation of the pegmatitic rocks at Site 1070. If the high Mg (and Cr) JOURNAL OF PETROLOGY VOLUME 43 content of their mafic phases (ilmenite, augite, kaersutite) simply reflects chemical exchange with the peridotite host, they provide little information on pegmatite petrogenesis. If the high Mg is a primary feature, however, these dikelets and fragments arguably represent melts parental to the main Site 1070 gabbro. The most plausible explanations for the high Mg#s are that they reflect either original igneous values or high-temperature exchange with peridotite. Because the minerals in question are essentially unzoned, any high-T diffusive exchange would have to have gone to completion. Fe/Mg diffusion coefficients for clinopyroxene at 900–1200°C range from 10−18 to 10−14 cm2/s (Dimanov & Sautter, 2000). Complete exchange with centimeter-sized crystals hence requires times of the order of 1 my at 1200°C, increasing to >10 my at 1150°C. Evidence suggests that the pegmatites are slowly cooled: the argon closure temperature for amphibole (>600°C) was attained several million years after the intrusion age inferred from zircon (Turrin, 1999). If diffusion were aided by grain-boundary diffusion along cracks and cleavage planes, it is just possible that the pegmatites were hot enough for long enough to effect the required Fe/ Mg exchange. However, we argue that several other lines of reasoning support an igneous origin for the high-Mg minerals. First, the augite and kaersutite, although relict in rodingite, still resemble their counterparts in the main pegmatite. The amphibole in the dikelets, for example, retains its characteristic red–brown color. With the exception of elevated Mg# (and Cr content), the relict mineral chemistry of the dikelets and clasts resembles that of igneous minerals in the main pegmatite (Figs 3 and 4; Table 2). In our view, this unrealistically restricts high-T exchange to only Fe, Mg, and Cr. Second, and perhaps more telling, there is substantial variation in Mg# and Cr amongst the various dike and clast samples, although all are more magnesian than the main pegmatite. If subsolidus processes (e.g. high-T diffusive exchange with the surrounding peridotite) were controlling Mg–Fe–Cr exchange, then Mg# and Cr content should be more or less the same in all of the exchanged samples. This is clearly not the case (Figs 3 and 4). Locally incomplete exchange cannot explain this lack of uniformity, as the variation reflects the compositions of relatively homogeneous minerals and is not an artifact of strong zoning or other intra-sample inhomogeneities. Finally, the Mg#s of coexisting amphibole and augite are consistent: high-Mg augite coexists with high-Mg amphibole (Figs 3 and 4). Origin of the albitites The coarse feldspar in the albitites appears to be of igneous origin. Indications for this include zoning, low, NUMBER 5 MAY 2002 but finite calcium content, and the observation that some of the albitites have granoblastic textures (similar to those locally developed in the gabbroid rocks) suggesting highT deformation of the albitic feldspar. This last observation is consistent with the albitites (and the sodic feldspar) being present before the onset of serpentinization. The lack of albite as a plagioclase alteration product in the pegmatites and rodingites (analcime is the sodic alteration phase) is circumstantial evidence for an igneous origin of the albite. The altered amphibole in one albitite, although it is basically tremolitic hornblende, still contains >1 wt % TiO2, probably as submicroscopic exsolved plates of ilmenite (Table 2). We interpret the albitites as modified differentiation products of the Site 1070 gabbroids. Timing of gabbro pegmatite intrusion vis-àvis rift development The Site 1070 pegmatites were emplaced at 127 Ma, underwent non-pervasive shearing at granulite facies conditions, and cooled through amphibole closure (>600°C) by 119 Ma and through plagioclase closure (>250°C) at 110 Ma (Manatschal et al., 2002). This long deformation and cooling history shows that melting predates complete unroofing and serpentinization. Furthermore, pegmatite intrusion post-dates syn-rift sediments that overlie serpentinite elsewhere on the Iberia Abyssal Plain (Whitmarsh et al., 1998). Pegmatite intrusion also post-dates the initiation of faulting at Hobby High, 100 km inboard of Site 1070. At the same time, the gabbro pegmatites appear to be slightly younger than the ocean crust in the western IAP (Whitmarsh & Wallace, 2001). All of this indicates that serpentinization, unroofing and magmatism were time-transgressive during nonvolcanic (more accurately, magma-poor) rifting and that the age of the sea floor formed by the unroofed mantle grows younger outboard, just as is the case for normal, volcanic ocean crust. The Penninic nappes of eastern Switzerland: an on-land analog The mafic and ultramafic rocks of the Penninic nappes in eastern Switzerland probably constitute the best documented on-land analog of the Iberian ocean–continent transition (Froitzheim & Manatschal, 1996; Manatschal & Bernoulli, 1999; Desmurs et al., 2002). Structural and petrological features (e.g. ophicalcites) all suggest that ultramafic rocks in these nappes were exhumed during continental rifting, serpentinized, and exposed at the sea floor. The serpentinites are intruded by a variety of igneous rocks including magnesian gabbros, Fe–Ti-enriched gabbros, dioritic and gabbroic pegmatites, and 900 BEARD et al. GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN albitites (Desmurs et al., 2002). The gabbroic rocks constitute >10% of the outcrop of the serpentinites. Overlying the mantle section are sedimentary rocks that are sporadically intercalated with pillowed basalt flows. Postrift, deep-water sediments overlie the entire package. The sequence of emplacement of igneous rocks has been interpreted as reflecting the transition from a non-volcanic margin to a slow spreading ridge. Petrogenesis The elemental and isotopic chemistry of the pegmatites is consistent with a moderately enriched mantle source (Figs 7 and 9). We argue that the composition and mineralogical character of the local peridotite (i.e. peridotites recovered on the IAP) permits it to be the source for the magmas that produced the pegmatites. Although the precise nature (supra-subduction zone or subcontinental lithosphere) of the IAP peridotite is a matter of debate, there is general agreement that it is moderately enriched (in LREE and LILE, for example) with respect to abyssal peridotite (Seifert & Brunotte, 1996; Abe, 2002; Hebert et al., 2002). In addition, Nd isotopic data (Nd = 4·0; Charpentier et al., 1998) on west Iberian mantle rocks both confirm their enriched character and are consistent with the isotopic composition of the Site 1070 gabbro pegmatites. The presence of modal kaersutite in several mantle samples from the IAP and elsewhere at the West Iberian margin (Agrinier et al., 1988; Cornen et al., 1996, 1999; Abe, 2002) is consistent with its enriched character, demonstrates that the mantle is not only enriched, but hydrous, and links the mantle with hydrous, kaersutite-bearing igneous rocks at Site 1070 and elsewhere at the West Iberian margin. More difficult to constrain than the nature of the mantle source is the issue of how and where the melting of the source occurred. We postulate that the gabbroic pegmatites represent melts that were generated in the upper mantle (probably in the plagioclase stability field) then collected and fractionated close to the site where they were generated. Our reasoning is as follows: (1) plagioclase peridotites (some kaersutite bearing) are widespread at the West Iberian margin (Evans & Girardeau, 1988; Cornen et al., 1996; Whitmarsh et al., 1998; Abe, 2002). These peridotites locally preserve evidence for partial melting (Cornen et al., 1996; Charpentier et al., 1998), although the age of the melting event is unknown. (2) Minerals in some Site 1070 pegmatite dikes are nearly magnesian enough to be in equilibrium with the mantle. These dikes may be parental to a suite of more differentiated rocks, particularly the main pegmatite. If so, this suggests that the local magmatic framework involved collection and storage of melt, a necessary (albeit not sufficient) characteristic of magma source regions. (3) The pegmatites, including the very magnesian examples, were probably near or at vapor saturation at the time of emplacement and, thus, not capable of sustained upward movement because of pressure-quench constraints. (4) The enrichment of the pegmatitic rocks in Y and HREE suggests melting at pressures less than garnet stability, and the sodic and aluminous nature of the pegmatites is consistent with hydrous melting in a plagioclase-bearing, low-pressure system (e.g. Rapp, 1995). Magmatism at Site 1070 was driven by a combination of decompression that accompanied unroofing and passive upwelling of hot, asthenospheric mantle. The presence of amphibole in the lithospheric mantle at the west Iberian margin is, we argue, key to magma genesis in the region, as amphibole peridotite in the plagioclase stability field begins to melt (dehydration-melting) at temperatures below 1050°C (Niida & Green, 1999). Our model for how this melting initiated is as follows: (1) a metasomatic event of unknown age, but probably related to Paleozoic rifting (Cornen et al., 1996) or subduction (Abe, 2002) results in the enrichment and, most importantly, partial hydration of the Iberian lithosphere with concomitant formation of plagioclase and kaersutitic amphibole. (2) A combination of unroofing owing to crustal thinning (Whitmarsh & Wallace, 2001) and northward migration of an upwelling asthenospheric bulge beneath the northward-opening rift (Scharer et al., 2000) causes the mantle to approach solidus conditions. (3) Where kaersutite is present in the lithosphere, the solidus temperature is relatively low and dehydrationmelting occurs at temperatures as low as 1050°C. This is >50°C below the dry solidus of depleted asthenospheric mantle at the same pressure (Niida & Green, 1999). Dehydration-melting generates small volumes of enriched, hydrous melts. Plagioclase is likely to be at least partially consumed during the dehydration-melting reactions [Niida & Green (1999); also see Vielzeuf & Schmidt (2001) for a review of dehydration-melting reactions and P–T geometry]. (4) As rifting progresses, asthenospheric melting begins and, eventually, progresses to the point where true oceanic crust is generated. Where the lithosphere is anhydrous, it is likely that melting of the much hotter asthenosphere will dominate magmatism throughout the early and, indeed, entire history of the rift. The observation that a hydrous lithosphere, where present, is the source for early melts in an evolving oceanic rift probably holds as a general case. ACKNOWLEDGEMENTS 901 I would like to thank the scientific and ship’s crew of ODP Leg 173 for a productive and stimulating cruise. JOURNAL OF PETROLOGY VOLUME 43 Thanks go to Jay Thomas for preparing the mineral separates. The manuscript has benefited considerably from the reviews of D. Bernoulli, O. 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Only amphibole and clinopyroxene Kd values were considered for REE because (1) there were virtually no data for ilmenite and (2) Kd values and concentrations for HREE in plagioclase were very low. In every case (except Cs), Kd values for at least two mineral species were considered, although if the high and low model were based on the same mineral, this may not be evident from Table A1. Sources for Table A1 are given in the footnote to the table. Other studies that were considered (and yielded intermediate values) are: Ringwood (1970); McCallum & Charette (1978); McKay et al. (1986); Dunn (1987); Green et al. (1989); Johnson & Kinzler (1989); LaTourrette & Burnett (1992); Nielsen et al. (1992); Adam et al. (1993); Hart & Dunn (1993); Forsythe et al. (1994); Jenner et al. (1994); Johnson (1994); Lundstrom et al. (1994); McKay et al. (1994); Bindeman et al. (1998). BEARD et al. GABBROIC PEGMATITES, IBERIA ABYSSAL PLAIN Table A1: Partition coefficients and data sources for inverse compositional models High value partition coefficient Element Mineral Partition Low value partition coefficient Source Mineral coefficient Partition Source coefficient Sr cpx 0·05 8 amph 0·08 K amph 0·96 1 plag 0·17 3 1 Rb amph 0·58 3 plag 0·05 12 Ba plag 1·5 12 plag 0·16 13 Cs plag 0·07 12 plag 0·07 12 Th plag 0·19 12 amph 0·02 7 Ta ilm 2·7 2 amph 0·08 3 Nb ilm 2·3 2 amph 0·08 3 Ce amph 0·84 1 cpx 0·04 8,9 Zr amph 1·56 1 cpx 0·06 8 Hf amph 1·53 1 cpx 0·12 9 Sm amph 1·8 1 amph 0·37 4 Ti amph 3·1 5 amph 0·87 4 Y cpx 0·9 1 amph 0·4 3 Yb amph 1·64 1 cpx 0·32 9 Sc amph 4·2 1 cpx 0·81 10 Cr cpx 34 1 cpx 1·7 10 Ni cpx 14 1 cpx 1·5 1 La amph 0·54 1 cpx 0·03 8 Nd amph 1·34 1 cpx 0·12 8,9 Eu amph 1·56 1 amph 0·31 6 Gd amph 2·01 1 amph 0·63 1 Tb amph 1·4 1 cpx 0·42 11 Dy amph 2·02 1 cpx 0·30 8,9 Ho amph 3 6 amph 0·44 4 Er amph 1·74 1 amph 0·55 1 Lu amph 1·57 1 amph 0·32 4 Sources: 1, Rollinson (1993); 2, Green & Pearson (1987); 3, Green et al. (1993); 4, Adam & Green (1994); 5, Sisson (1994); 6, Green & Pearson (1985); 7, Brenan et al. (1995); 8, Skulski et al. (1994); 9, Fujimaki & Tatsumoto (1984); 10, Hauri et al. (1994); 11, Paster et al. (1974); 12, Dunn & Sen (1994); 13, Drake & Weill (1975). 905