Abstract

The <80 ka basalts–basanites of the Potrillo Volcanic Field (PVF) form scattered scoria cones, lava flows and maars adjacent to the New Mexico–Mexico border. MgO ranges up to 12·5%; lavas with MgO < 10·7% have fractionated both olivine and clinopyroxene. Cumulate fragments are common in the lavas, as are subhedral megacrysts of aluminous clinopyroxene (with pleonaste inclusions) and kaersutitic amphibole. REE modelling indicates that these megacrysts could be in equilibrium with the PVF melts at ∼1·6–1·7 GPa pressure. The lavas fall into two geochemical groups: the Main Series (85% of lavas) have major- and trace-element abundances and ratios closely resembling those of worldwide ocean-island alkali basalts and basanites (OIB); the Low-K Series (15%) differ principally by having relatively low K2O and Rb contents. Otherwise, they are chemically indistinguishable from the Main Series lavas. Sr- and Nd-isotopic ratios in the two series are identical and vary by scarcely more than analytical error, averaging 87Sr/86Sr = 0·70308 (SD = 0·00004) and 143Nd/144Nd = 0·512952 (SD=0·000025). Such compositions would be expected if both series originated from the same mantle source, with Low-K melts generated when amphibole remained in the residuum. Three PVF lavas have very low Os contents (<14 ppt) and appear to have become contaminated by crustal Os. One Main Series picrite has 209 ppt Os and has a γOs value of +13·6, typical for OIB. This contrasts with published 187Os/188Os ratios for Kilbourne Hole peridotite mantle xenoliths, which give mostly negative γOs values and show that Proterozoic lithospheric mantle forms a thick Mechanical Boundary Layer (MBL) that extends to ∼70 km depth beneath the PVF area. The calculated mean primary magma, in equilibrium with Fo89, has Na2O and FeO contents that give a lherzolite decompression melting trajectory from 2·8 GPa (95 km depth) to 2·2 GPa (70 km depth). Inverse modelling of REE abundances in Main Series Mg-rich lavas is successful for a model invoking decompression melting of convecting sub-lithospheric lherzolite mantle (ϵNd = 6·4; Tp ∼ 1400°C) between 90 and 70 km. Nevertheless, such a one-stage model cannot account for the genesis of the Low-K Series because amphibole would not be stable within convecting mantle at Tf ∼ 1400°C. These magmas can only be accommodated by a three-stage model that envisages a Thermal Boundary Layer (TBL) freezing conductively onto the ∼70 km base of the Proterozoic MBL during the ∼20 Myr tectonomagmatic quiescence before PVF eruptions. As it grew, this was veined by hydrous small-fraction melts from below. The geologically recent arrival of hotter-than-ambient (Tp ∼ 1400°C) convecting mantle beneath the Potrillo area re-melted the TBL and caused the magmatism.

INTRODUCTION

Small-volume basic alkalic magmatism characterizes many current and geologically recent areas of continental extension. The geochemistry of the most primitive magmatic rocks can be used to model mantle partial-melting conditions and, in geologically young examples, these results may be compared with such features as: (1) geophysical data bearing on lithospheric thickness and the state of the subjacent convecting mantle; and (2) the petrology and geochemistry of the local lithospheric mantle, in cases where this is sampled by xenolith suites in the magmatic rocks. Hence, the relative contributions of melts from lithospheric and convecting mantle can be assessed.

Wang et al. (2002) focused on the major-element compositions of <10 Ma basic lavas, many of them alkaline, from throughout the Basin and Range province and adjacent regions of the southwest USA. They used techniques originally developed for understanding MORB petrogenesis to deduce that the sources of these continental magmas were essentially entirely within the convecting mantle, beneath a lithospheric lid of varying thickness. Wang et al. (2002) showed that their modelled depths to the base of the lithosphere in western USA agreed well with geophysical data for the region.

In contrast, other authors have noted geochemical features of worldwide alkaline basic lavas that have seemed to them to be incompatible with an origin simply by partial melting of convecting mantle. They have developed alternative magma genesis models in which the lithospheric mantle also plays a role. For instance, invocation of a geologically young metasomatic event affecting older lithospheric mantle, which then contributes to the melts, is the most widely favoured method of explaining the characteristic elemental and isotopic geochemistry of such magmatism occurring on both continental and old oceanic lithosphere (e.g. Wilson & Downes, 1991; McKenzie & O'Nions, 1995; Wilson et al., 1995; Class & Goldstein, 1997; Späth et al., 2001; Yang et al., 2003).

This study concerns the <80 ka Potrillo Volcanic Field (PVF), southern New Mexico, which occurs in the region where the Rio Grande rift ceases to be a distinct physiographic feature and joins the regional Basin and Range extension zone of the southwestern USA and northern Mexico. The chemical and Sr–Nd–Os isotopic compositions of the lavas are used to attempt to define the depth and conditions of their genesis, and of the subsequent processes that affected them. Such an approach is assisted in this area by the availability of extensive published data on both the regional geophysics and the petrology and geochemistry of suites of lithospheric mantle xenoliths sampled by some of the lavas.

POTRILLO VOLCANIC FIELD

The PVF lies in the desert that extends for ∼60 km west of the segment of the Rio Grande rift between El Paso and Las Cruces. It encroaches slightly into Mexico but falls mostly within New Mexico. The magmatism is entirely basaltic (s.l.), forming scattered scoria cones (Fig. 1). Up to several short lava flows originate from each cone and there are a few locations (e.g. Aden Crater) that appear to be centres of more persistent magmatism. The region is well known for the spectacular maar and suite of mantle xenoliths at Kilbourne Hole (Padovani & Reid, 1989). Most of the scoria cones lie along the broad approximately N–S trending ridge of the West Potrillo Mountains (Fig. 1). Another small line of cones forms the Black Hills, between Kilbourne Hole and the Rio Grande. The ages of the lavas, measured by 3He surface exposure dating, range from 80 to 17 ka (Anthony & Poths, 1992).

Fig. 1.

Sketch map showing the distribution of eruptive centres and sample locations in the Potrillo area, New Mexico, southwest USA. The West Potrillo Mountains are a broad ridge oriented approximately N–S.

Fig. 1.

Sketch map showing the distribution of eruptive centres and sample locations in the Potrillo area, New Mexico, southwest USA. The West Potrillo Mountains are a broad ridge oriented approximately N–S.

The varied basement rock types beneath the Potrillo volcanics, and locally protruding through them, range from Proterozoic granites through a Phanerozoic sedimentary succession to basalt–andesite volcanics of the southern fringes of the Sierra de las Uvas volcanic field—an ∼30 Ma igneous succession that lies mostly to the north of Fig. 1 (Seager et al., 1984; unpublished work by Thompson and Leat). The Potrillo area is generally classified as part of the southern Rio Grande rift and shows the Tertiary tectonic evolution of that structure. Extension took place in an intense 30–20 Ma phase, involving low-angle normal faults, and a less intense <10 Ma phase, involving high-angle normal faults (Olsen et al., 1987; Keller et al., 1990). Mack & Seager (1995) argued that the Quaternary magmatism reached the surface via a transfer zone linking two adjacent N–S-trending, long-lived, extensional structures—the West Robledo and Camel Mountain faults.

In order to decide whether the PVF magmas originated within the subcontinental lithospheric mantle, the underlying convecting mantle or both, it is first necessary to summarize what is known about the lithospheric mantle in this area.

Previous geophysical estimates of lithosphere thickness

Previous attempts to estimate lithospheric thickness beneath the southern Rio Grande rift have been made by modelling seismic, heat flow and gravity data, and based on thermo-barometric studies of mantle xenoliths in the PVF lavas, notably at Kilbourne Hole. Sinno et al. (1986) established seismic refraction profiles crossing the Potrillo area, both E–W (reversed) and N–S (unreversed). They interpreted these data as showing that the Rio Grande rift at 32–33°N is a distinct structure, separating the Great Plains and Basin and Range tectonic provinces. The low value of Pn (P-wave velocity for the ray path reflected from immediately below the Moho) beneath the rift (7·7 km/s) led them to suggest that ‘the asthenosphere could be in direct contact with the crust’ (Sinno et al., 1986, p. 6154), i.e. that no lithospheric mantle exists beneath the Potrillo area.

Morgan et al. (1986) suggested that the very high heat flow values across the axis of the southern Rio Grande rift ‘indicate a convective heat transfer regime in the upper mantle’ (Morgan et al., 1986, p. 6269). Keller et al. (1990) refined this view by suggesting that below 35 km depth in the southern Rio Grande rift, ‘the geotherm appears to become non-conductive and is probably controlled by convection of basaltic magmas from depths of 50 to 70 km up into the lower crust’ (Keller et al., 1990, p. 31). It is not clear whether they envisaged this process as melt generation within upwelling convecting mantle (asthenosphere) or melt migration advecting heat from below 50 km into overlying lithosphere. Cordell et al. (1991) derived a map of total lithospheric thickness beneath the southern Rio Grande rift by inversion of gravity anomaly data and concluded that beneath the Potrillo area ‘the lithosphere may be as thin as 40 km’ (Cordell et al., 1991, p. 6566).

Achauer & Masson (2002) summarized previous extensive teleseismic travel-time delay tomographic studies of the rift north of the Potrillo area. They reached two main conclusions: (1) Below 185 km depth, there is an extensive zone of seismically relatively slow mantle (interpreted as relatively hot). This is centred on the physiographic rift but extends for 100–200 km on either side and widens downwards. (2) Between 35 km (∼Moho) and 185 km there are several areas along the physiographic rift axis and its flanks where seismically slow (hot) mantle extends up to the base of the crust. They interpreted the laterally extensive, deep, hot mantle as sourced by a convective plume and suggested that asthenospheric diapirism may have carried plume mantle up to near-Moho depths locally. They emphasized that in these seismic features, the Rio Grande rift resembles the East African rift and contrasts with the Rhinegraben and Baikal rifts, which both lack widespread deep seismically slow mantle. The zone of upwelling convecting mantle appears to be slightly oblique to the trend of the rift axis, such that the Potrillo area overlies the eastern flank rather than the crest of the present-day upwarp of low-velocity mantle. The geophysically defined lithospheric thickness beneath Kilbourne Hole is ∼80 km.

Petrological features of the lithospheric mantle beneath Kilbourne Hole

The extensive and well-studied suite of mantle xenoliths that occur at Kilbourne Hole (Fig. 1) gives an important perspective on the nature of the lithospheric mantle beneath the PVF (Padovani & Reid, 1989). Most of the mafic–ultramafic xenoliths are medium- to coarse-grained so-called Group I peridotites, with protogranular textures and green Cr diopside. Thin bands of pyroxenite cutting this peridotite are rare. About 10% of the xenoliths are more Fe-rich, finer-grained, tabular-textured so-called Group II peridotites, frequently veined by pyroxenite and wehrlite (Irving, 1980; Padovani & Reid, 1989; Bussod & Williams, 1991). All of these authors stress that hydrous minerals, such as amphibole and mica, are rare in the Kilbourne Hole xenolith suite.

Bussod & Williams (1991) provided crucial evidence by showing that calculated PT equilibration conditions for an extensive suite of Kilbourne Hole peridotite xenoliths defined a continuous array from 850°C at 35 km depth to 1050°C at 67 km depth. This is far below the temperature of ∼1320°C in this depth range attained by mantle convecting at a potential temperature (Tp) of 1300°C (McKenzie & Bickle, 1988). Potential temperatures significantly lower than this value are not usually suggested for convecting mantle. Thus, the Kilbourne Hole peridotite xenoliths appear to provide tangible evidence for stable (Mechanical Boundary Layer) lithospheric mantle extending to at least ∼70 km beneath the area. Furthermore, Bussod & Williams (1991, fig. 1) noted that within the set of 18 xenoliths that they studied, the depth distribution of Group I and II peridotites was anything but random; the fine-grained tabular-textured Group II peridotites and their pyroxenite-veined composite associates are confined to the shallower (35–55 km) part of the lithospheric mantle, whilst Group I coarse peridotites are confined to the 50–70 km depth range and are the only type sampled below 55 km.

PETROGRAPHY AND MINERALOGY OF THE PVF LAVAS

Sampling was designed to give regional coverage of the volcanics; sample localities are given in Electronic Appendix A (which may be downloaded from http://www.petrology.oupjournals.org). Either lava flows or large spindle bombs provided suitable material. In addition, small sets of the megacrysts and coarsely crystalline ultramafic inclusions that occur at many places in the field were collected; Kilbourne Hole is the best-known location for these suites (Padovani & Reid, 1989). A few of the localities at the northern end of the West Potrillo Mountains (Fig. 1) provided samples of the older Uvas volcanics (Seager et al., 1984; Olsen et al., 1987). The locality and geochemical data for West Potrillo Mountains Uvas lavas sampled during this study are included in Electronic Appendices A and B.

All the lavas are porphyritic, with up to ∼20% (mostly <10%) of euhedral to subhedral phenocrysts, rarely exceeding 2 mm, in fine-grained holocrystalline groundmasses of olivine, clinopyroxene, plagioclase and opaques, together with sparse patches of brown glass. The phenocrysts are of olivine alone, olivine plus clinopyroxene, or olivine plus clinopyroxene and plagioclase (occasionally). One lava (8104) contains sparse opaque-granule-rich pseudomorphs, probably after amphibole, plus trace amounts of <1 mm equant opaque and apatite microphenocrysts.

Apart from sample 8104, no hydrous minerals were observed in thin sections as either phenocrysts or groundmass phases. Nevertheless, large amphibole megacrysts (<5 cm) occur quite commonly throughout the West Potrillo Mountains area, mostly in scoria but also as rare fragments within lavas. Other abundant megacrysts are clinopyroxene and feldspar; Ortiz (1980) mentioned rare spinel. Crystals of each mineral, found loose in scoria, are euhedral but characterized by polished and pitted surfaces.

Coarse-grained ultramafic inclusions, up to ∼10 cm in size, are locally common throughout the West Potrillo Mountains area. Spinel lherzolites are bright green, with a coarse equigranular texture, curved grain boundaries, olivine with deformation lamellae and brown spinel. They predominate at Kilbourne Hole (Roden et al., 1988; Padovani & Reid, 1989; Bussod & Williams, 1991; Burton et al., 1999) but are a minor part of the West Potrillo Mountains suites. One West Potrillo Mountains green xenolith is a granular olivine websterite, with some brown spinel and coarse orthopyroxene exsolution lamellae in its clinopyroxenes. All the remaining ultramafic inclusions form a series between clinopyroxenite (± olivine) and hornblendite. Lherzolite-like granular textures characterize the hornblendites but members of the clinopyroxenite to amphibole clinopyroxenite suite have igneous cumulate textures, with euhedral to subhedral minerals surrounded by a network of brown interstitial glass. Some clinopyroxenites lack olivine (except as tiny euhedra accompanying plagioclase in the glass) and amphibole (except as flecks along clinopyroxene cleavage planes). As amphibole increases in abundance, it forms poikilocrysts and eventually equant crystals. One olivine clinopyroxenite with a few percent of brown amphibole shows small euhedra of this phase within pools of interstitial glass. Spinel forms rounded <0·5 mm grains in members of the clinopyroxenite–hornblendite suite. In fresh samples, the spinel is green in the clinopyroxenites, including some with poikilitic amphibole, and opaque in the hornblendites. Many inclusions are variably oxidized and these have opaque spinels. Plagioclase is common in the interstitial glass patches and also occurs as a minor part of the main mineral assemblage in two of the hornblende clinopyroxenites.

The olivine–clinopyroxene-phyric lavas commonly contain rounded-fritted olivine, pyroxene and feldspar macrocrysts, in addition to their euhedral phenocrysts. Electron microprobe study of the olivine and pyroxene macrocrysts (described below) shows that neither is from the same spinel lherzolite source as the mantle xenoliths. The 11 lavas used for electron microprobe studies have MgO contents between 9·1 and 12·5%. Analyses were made using the Geoscan (mostly) and Cameca SX100 instruments at Manchester University (Table 1). The analyses concentrated on phenocrysts, macrocrysts, megacrysts and other minerals included in them, in order to shed light on the post-genesis history of the lavas. The Geoscan settings were as follows: accelerating potential 15 kV; incident beam current 6 nA; take-off angle 75°. Standards used were: Si, Ca, wollastonite; Ti, rutile (synthetic); Al, corundum (synthetic); Fe, fayalite (synthetic); Mn, tephroite (synthetic); Mg, periclase (synthetic); Na, jadeite; K, orthoclase; Cr, Cr2O3 (synthetic); Ni, NiO (synthetic). The analysis system was a PGT (Princeton Gamma Tech) energy-dispersive Excalibur system. All data were ZAF corrected.

Table 1:

Representative electron microprobe mineral analyses

Sample
 
Number
 
Location of analysis
 
SiO2
 
TiO2
 
Al2O3
 
Cr2O3
 
FeO
 
MnO
 
MgO
 
NiO
 
CaO
 
Na2O
 
K2O
 
Total
 
Olivines               
898 Euhedral phenocryst core 40·04 nm nm 0·04 11·60 0·26 46·56 0·25 0·30 nm nm 99·04 
898 Groundmass 38·82 nm nm 0·01 19·51 0·28 39·99 0·30 0·55 nm nm 99·45 
8101 Euhedral phenocryst core 40·23 nm nm 0·09 12·50 0·22 46·27 0·48 0·25 nm nm 100·04 
8101 Rounded macrocryst core 38·16 nm nm 0·00 23·27 0·25 37·85 0·22 0·14 nm nm 99·89 
8138 Rounded macrocryst core 40·33 nm nm 0·00 12·36 0·23 46·53 0·38 0·27 nm nm 100·08 
8138 Rounded macrocryst core 38·07 nm nm 0·02 24·39 0·34 37·23 0·15 0·19 nm nm 100·40 
8138 Rim of 6 39·70 nm nm 0·12 14·41 0·14 44·50 0·26 0·25 nm nm 99·37 
8140 Rounded macrocryst core 37·01 nm nm 0·01 30·42 0·38 32·19 0·14 0·08 nm nm 100·22 
8140 Rounded macrocryst core 38·23 nm nm 0·00 22·87 0·39 38·16 0·01 0·26 nm nm 99·92 
8159 10 Euhedral phenocryst core 40·02 nm nm 0·07 13·73 0·25 45·25 0·22 0·24 nm nm 99·78 
8164 11 Euhedral phenocryst core 40·10 nm nm 0·05 12·88 0·22 45·78 0·37 0·25 nm nm 99·63 
8164 12 Groundmass 38·79 nm nm 0·03 19·48 0·36 40·73 0·23 0·37 nm nm 99·98 
Clinopyroxenes               
676 20 Groundmass 48·93 2·03 4·13 0·10 8·79 0·10 12·98 0·00 22·02 0·57 nm 99·64 
676 21 Groundmass 41·34 5·95 9·22 0·10 10·68 0·16 9·11 0·15 22·20 0·84 nm 99·76 
683a 22 Kilbourne Hole megacryst (13 pts) 47·82 1·53 8·76 0·09 7·19 0·14 13·32 0·07 19·24 1·37 nm 99·52 
683a 23 Kilbourne Hole megacryst rim 44·18 3·26 10·56 0·11 6·66 0·12 11·99 0·12 22·13 0·58 nm 99·70 
879 24 Groundmass 47·33 2·31 6·20 0·27 7·39 0·08 13·58 0·30 21·99 0·78 nm 100·22 
879 25 Resorbed macrocryst core 48·05 1·22 8·26 0·12 7·72 0·21 11·64 0·14 20·96 1·28 nm 99·60 
879 26 Euhedral phenocryst rim 41·48 4·71 11·10 0·10 8·46 0·06 10·42 0·12 22·57 0·80 nm 99·82 
898 27 Resorbed macrocryst core 50·17 0·95 5·42 0·39 4·55 0·11 15·33 0·01 22·15 0·83 nm 99·90 
898 28 Resorbed macrocryst core 48·04 1·45 7·33 0·10 8·48 0·23 13·37 0·03 19·10 1·14 nm 99·27 
8101 29 Resorbed macrocryst core 46·63 1·70 8·66 0·07 8·63 0·11 11·67 0·00 20·70 1·07 nm 99·25 
8101 30 Rim of 29 47·38 2·30 5·78 0·12 7·81 0·13 12·66 0·06 22·64 0·61 nm 99·48 
8134 31 Euhedral phenocryst core 49·21 1·45 5·42 0·46 5·65 0·13 15·09 0·07 21·00 0·81 nm 99·29 
8138 32 Euhedral phenocryst core 48·25 1·11 7·04 0·00 9·48 0·36 11·70 0·10 20·11 0·94 nm 99·08 
8138 33 Rim of 32 47·34 2·70 6·54 0·21 6·75 0·05 13·05 0·11 23·12 0·31 nm 100·18 
Spinels               
683 13 Inside cpx megacryst 0·18 0·76 57·54 0·00 24·19 0·20 17·20 0·13 nm nm nm 100·20 
683 14 Inside cpx megacryst 0·35 0·93 57·62 0·00 23·20 0·12 17·17 0·45 nm nm nm 99·83 
8101 15 Inside olivine macrocryst 0·08 1·65 54·14 0·07 27·28 0·18 15·56 0·06 nm nm nm 99·02 
8134 16 Inside resorbed cpx macrocryst 0·46 2·72 48·98 0·12 26·93 0·11 18·09 0·25 nm nm nm 97·66 
8134 17 Inside resorbed cpx macrocryst 0·50 0·89 56·11 0·00 22·09 0·15 17·98 0·11 nm nm nm 97·82 
8140 18 Inside olivine macrocryst 0·47 0·86 39·01 21·29 19·06 0·23 16·52 0·30 nm nm nm 97·74 
8140 19 Groundmass 0·34 24·50 4·43 0·01 63·40 0·76 2·65 0·00 nm nm nm 96·08 
Plagioclases               
8134 34 Rounded macrocryst 55·66 0·19 27·66 0·02 0·11 0·00 0·00 0·09 9·79 5·83 0·39 99·74 
8138 35 Rounded macrocryst 60·72 0·02 24·22 0·09 0·00 0·01 0·04 0·00 6·22 8·00 0·90 100·22 
Sample
 
Number
 
Location of analysis
 
SiO2
 
TiO2
 
Al2O3
 
Cr2O3
 
FeO
 
MnO
 
MgO
 
NiO
 
CaO
 
Na2O
 
K2O
 
Total
 
Olivines               
898 Euhedral phenocryst core 40·04 nm nm 0·04 11·60 0·26 46·56 0·25 0·30 nm nm 99·04 
898 Groundmass 38·82 nm nm 0·01 19·51 0·28 39·99 0·30 0·55 nm nm 99·45 
8101 Euhedral phenocryst core 40·23 nm nm 0·09 12·50 0·22 46·27 0·48 0·25 nm nm 100·04 
8101 Rounded macrocryst core 38·16 nm nm 0·00 23·27 0·25 37·85 0·22 0·14 nm nm 99·89 
8138 Rounded macrocryst core 40·33 nm nm 0·00 12·36 0·23 46·53 0·38 0·27 nm nm 100·08 
8138 Rounded macrocryst core 38·07 nm nm 0·02 24·39 0·34 37·23 0·15 0·19 nm nm 100·40 
8138 Rim of 6 39·70 nm nm 0·12 14·41 0·14 44·50 0·26 0·25 nm nm 99·37 
8140 Rounded macrocryst core 37·01 nm nm 0·01 30·42 0·38 32·19 0·14 0·08 nm nm 100·22 
8140 Rounded macrocryst core 38·23 nm nm 0·00 22·87 0·39 38·16 0·01 0·26 nm nm 99·92 
8159 10 Euhedral phenocryst core 40·02 nm nm 0·07 13·73 0·25 45·25 0·22 0·24 nm nm 99·78 
8164 11 Euhedral phenocryst core 40·10 nm nm 0·05 12·88 0·22 45·78 0·37 0·25 nm nm 99·63 
8164 12 Groundmass 38·79 nm nm 0·03 19·48 0·36 40·73 0·23 0·37 nm nm 99·98 
Clinopyroxenes               
676 20 Groundmass 48·93 2·03 4·13 0·10 8·79 0·10 12·98 0·00 22·02 0·57 nm 99·64 
676 21 Groundmass 41·34 5·95 9·22 0·10 10·68 0·16 9·11 0·15 22·20 0·84 nm 99·76 
683a 22 Kilbourne Hole megacryst (13 pts) 47·82 1·53 8·76 0·09 7·19 0·14 13·32 0·07 19·24 1·37 nm 99·52 
683a 23 Kilbourne Hole megacryst rim 44·18 3·26 10·56 0·11 6·66 0·12 11·99 0·12 22·13 0·58 nm 99·70 
879 24 Groundmass 47·33 2·31 6·20 0·27 7·39 0·08 13·58 0·30 21·99 0·78 nm 100·22 
879 25 Resorbed macrocryst core 48·05 1·22 8·26 0·12 7·72 0·21 11·64 0·14 20·96 1·28 nm 99·60 
879 26 Euhedral phenocryst rim 41·48 4·71 11·10 0·10 8·46 0·06 10·42 0·12 22·57 0·80 nm 99·82 
898 27 Resorbed macrocryst core 50·17 0·95 5·42 0·39 4·55 0·11 15·33 0·01 22·15 0·83 nm 99·90 
898 28 Resorbed macrocryst core 48·04 1·45 7·33 0·10 8·48 0·23 13·37 0·03 19·10 1·14 nm 99·27 
8101 29 Resorbed macrocryst core 46·63 1·70 8·66 0·07 8·63 0·11 11·67 0·00 20·70 1·07 nm 99·25 
8101 30 Rim of 29 47·38 2·30 5·78 0·12 7·81 0·13 12·66 0·06 22·64 0·61 nm 99·48 
8134 31 Euhedral phenocryst core 49·21 1·45 5·42 0·46 5·65 0·13 15·09 0·07 21·00 0·81 nm 99·29 
8138 32 Euhedral phenocryst core 48·25 1·11 7·04 0·00 9·48 0·36 11·70 0·10 20·11 0·94 nm 99·08 
8138 33 Rim of 32 47·34 2·70 6·54 0·21 6·75 0·05 13·05 0·11 23·12 0·31 nm 100·18 
Spinels               
683 13 Inside cpx megacryst 0·18 0·76 57·54 0·00 24·19 0·20 17·20 0·13 nm nm nm 100·20 
683 14 Inside cpx megacryst 0·35 0·93 57·62 0·00 23·20 0·12 17·17 0·45 nm nm nm 99·83 
8101 15 Inside olivine macrocryst 0·08 1·65 54·14 0·07 27·28 0·18 15·56 0·06 nm nm nm 99·02 
8134 16 Inside resorbed cpx macrocryst 0·46 2·72 48·98 0·12 26·93 0·11 18·09 0·25 nm nm nm 97·66 
8134 17 Inside resorbed cpx macrocryst 0·50 0·89 56·11 0·00 22·09 0·15 17·98 0·11 nm nm nm 97·82 
8140 18 Inside olivine macrocryst 0·47 0·86 39·01 21·29 19·06 0·23 16·52 0·30 nm nm nm 97·74 
8140 19 Groundmass 0·34 24·50 4·43 0·01 63·40 0·76 2·65 0·00 nm nm nm 96·08 
Plagioclases               
8134 34 Rounded macrocryst 55·66 0·19 27·66 0·02 0·11 0·00 0·00 0·09 9·79 5·83 0·39 99·74 
8138 35 Rounded macrocryst 60·72 0·02 24·22 0·09 0·00 0·01 0·04 0·00 6·22 8·00 0·90 100·22 

All Fe measured as FeO; nm, not measured. See text for analytical methods.

Olivine

Representative olivine analyses are given in Table 1. A plot of olivine forsterite content versus whole-rock Mg-number (Fig. 2a) summarizes the core compositions and marginal zoning of euhedral phenocrysts, plus groundmass compositions. If a value of Kd for distribution of Fe2+ and Mg between olivine and melt of 0·30 is used (Ulmer, 1989), it appears that all the phenocryst cores in bulk rocks with Mg-number <68 (Table 2) are approximately in equilibrium with their bulk-rock compositions. The two most Mg-rich samples—8159 and 8164 (Table 1)—plot below this curve. This relationship can be caused by olivine phenocryst accumulation in a melt and such an explanation may be appropriate because several percent of >2 mm olivine phenocrysts occur in each lava. Nevertheless, this conclusion is substantially model-dependent because the bulk-rock Mg-number values used in Fig. 2 were calculated with 10% of the Fe oxide in each analysis allocated to Fe2O3. If all the Fe is taken as FeO, the phenocryst cores in 8159 and 8164 lie on the Kd = 0·30 line and those of the lower Mg-number samples lie above it. Thus, Fig. 2a can in no sense be taken to ‘prove’ that samples 8159 and 8164 either are, or are not, olivine cumulates.

Table 2:

Whole rock analyses of representative lavas from the Portillo Volcanic Field

Sample: 8159 8164 6110 8129 6108 6158 8144 8112 8101 676 875 870 869 8140 6140 894 8136 672 898 
Series:
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Low-K
 
Main
 
Main
 
Main
 
Main
 
Low-K
 
Low-K
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
SiO2 45·41 44·94 43·18 44·67 44·88 42·72 44·35 45·38 45·84 45·09 45·12 46·39 46·65 44·87 44·46 44·29 45·05 45·49 43·97 
TiO2 2·08 2·05 2·20 2·27 2·31 2·25 2·37 2·21 2·23 2·02 2·36 2·19 2·28 2·27 2·38 2·32 2·34 2·09 2·15 
Al2O3 13·25 13·63 13·55 14·11 13·55 14·15 13·95 14·93 14·30 14·17 15·10 13·92 14·78 14·17 14·70 14·13 14·74 14·94 14·44 
Fe2O3* 12·27 12·64 12·77 12·47 12·63 12·38 12·77 11·24 11·98 11·99 11·41 13·18 11·47 11·58 11·65 11·94 11·48 11·80 11·15 
MnO 0·17 0·18 0·20 0·20 0·18 0·20 0·19 0·18 0·18 0·19 0·17 0·19 0·19 0·18 0·18 0·17 0·17 0·18 0·17 
MgO 12·52 12·32 12·26 11·75 11·74 11·70 11·51 11·47 11·24 11·18 11·11 11·09 10·98 10·96 10·91 10·81 10·76 10·65 10·62 
CaO 10·56 9·66 9·40 11·41 9·84 11·32 10·91 10·38 10·82 10·34 10·71 10·46 10·70 10·54 10·73 10·91 11·08 10·87 11·35 
Na22·69 2·98 3·30 2·96 3·16 3·06 2·88 3·28 3·14 3·05 3·01 2·51 3·76 3·27 3·20 3·24 3·26 2·91 2·82 
K21·35 1·47 1·55 1·59 1·59 1·60 1·21 1·58 1·58 1·48 1·53 0·75 0·94 1·60 1·67 1·56 1·60 1·49 1·55 
P2O5 0·47 0·49 0·48 0·47 0·51 0·36 0·51 0·51 0·49 0·42 0·47 0·44 0·44 0·48 0·47 0·51 0·43 0·48 0·50 
Total 100·77 100·37 98·88 101·89 100·39 99·75 100·66 101·16 101·79 99·91 100·99 101·13 102·19 99·92 100·35 99·88 100·91 100·90 99·90 
LOI 1·09 0·08 gain 0·34 gain 0·01 0·55 0·13 0·26 gain 0·28 0·85 0·11 0·26 0·01 0·62 0·33 gain 1·18 
Mg-no. 66·90 65·88 65·54 65·12 64·81 65·18 64·10 66·90 65·02 64·88 65·86 62·50 65·47 65·22 64·98 64·20 65·00 64·13 65·37 
Ba 380 336 511 470 479 711 406 466 398 455 349 345 455 384 493 413 483 488 505 
Cr 375 397 409 378 326 348 324 291 398 383 312 352 325 408 335 314 369 326 367 
Hf 3·68 3·73 4·67 4·43 4·81 3·98 3·94 4·81 3·94 4·00 4·44 3·83 4·25 4·18 4·36 4·10 4·39 4·31 4·01 
Nb 43 44 52 58 53 54 53 55 50 43 53 40 49 50 53 56 49 48 53 
Ni 304 320 332 219 285 223 234 217 245 220 198 222 192 235 199 217 178 184 216 
Pb 0·81 1·89 2·40 2·30 1·56 2·19 1·04 1·49 1·24 1·70 1·23 0·76 2·26 2·22 2·18 1·15 1·26 1·53 1·26 
Rb 25 28 31 35 32 36 21 42 31 26 34 14 21 33 33 28 32 26 32 
Sc 28·7 27·8 29·5 27·7 29·3 34·1 32·5 32·2 33·9 33·0 32·0 30·6 31·7 32·1 32·3 33·3 34·3 32·6 31·5 
Sr 526 506 574 778 592 571 903 589 550 545 583 624 571 572 590 588 524 559 620 
Ta 2·93 2·83 3·42 3·95 3·61 3·62 3·52 3·91 3·35 2·96 3·57 2·69 3·44 3·34 3·55 3·85 3·49 3·29 3·60 
Th 2·86 3·13 3·95 4·41 3·81 4·25 3·40 4·40 3·51 3·18 3·90 2·53 4·23 3·95 4·16 4·02 3·88 3·64 3·71 
0·84 0·89 1·10 1·29 1·02 1·13 0·90 1·21 1·00 0·56 1·06 0·58 1·11 1·05 1·10 1·08 1·06 1·00 1·02 
211 197 228 221 223 237 240 241 236 226 238 213 215 257 231 238 270 227 228 
22 24 30 25 28 28 27 26 25 29 29 25 27 25 29 27 27 30 24 
Zn 78 91 88 84 78 82 84 73 79 78 73 84 65 74 72 78 71 79 74 
Zr 160 148 199 179 200 164 160 200 158 147 181 146 169 167 179 161 164 167 161 
La 27·7 27·1 33·9 36·7 34·9 32·0 31·5 36·2 30·4 29·3 33·4 26·3 33·1 31·1 32·7 34·0 32·7 31·6 33·4 
Ce 53·7 52·8 65·2 68·8 68·1 60·8 61·2 68·8 58·4 57·3 64·0 52·9 62·1 58·9 62·2 64·6 62·3 61·6 63·7 
Pr 7·05 6·69 8·18 8·51 8·87 7·61 7·96 8·73 7·51 7·64 8·21 7·08 7·64 7·40 7·70 8·30 8·04 8·07 8·12 
Nd 29·0 27·9 33·6 33·5 35·8 31·3 32·7 34·5 30·3 31·0 32·9 29·4 31·2 30·2 31·5 33·4 32·3 33·0 32·6 
Sm 5·92 5·70 6·78 6·52 7·20 6·37 6·59 6·81 6·20 6·41 6·63 6·17 6·30 6·13 6·39 6·77 6·50 6·70 6·42 
Eu 1·93 1·89 2·18 2·15 2·26 2·03 2·11 2·11 1·98 1·99 2·11 1·99 2·03 1·98 2·09 2·16 2·08 2·10 2·06 
Gd 5·66 5·73 6·66 6·24 6·62 6·22 6·23 6·21 5·88 6·04 6·25 5·96 6·29 6·18 6·50 6·26 6·13 6·24 5·92 
Tb 0·86 0·86 0·99 0·91 1·01 0·92 0·94 0·93 0·89 0·92 0·95 0·92 0·93 0·89 0·94 0·96 0·94 0·96 0·90 
Dy 4·71 4·63 5·35 4·90 5·55 5·05 5·07 5·12 4·85 5·07 5·24 5·09 5·05 4·88 5·13 5·23 5·27 5·36 4·87 
Ho 0·90 0·91 1·04 0·94 1·07 0·97 0·97 0·98 0·93 0·99 1·00 0·98 0·99 0·94 1·00 1·00 1·03 1·03 0·93 
Er 2·24 2·26 2·64 2·36 2·74 2·49 2·43 2·52 2·34 2·49 2·55 2·47 2·51 2·40 2·52 2·55 2·63 2·65 2·36 
Tm 0·35 0·35 0·41 0·37 0·44 0·39 0·38 0·39 0·36 0·39 0·40 0·38 0·39 0·38 0·39 0·39 0·42 0·42 0·36 
Yb 2·00 2·04 2·44 2·15 2·50 2·30 2·14 2·28 2·06 2·20 2·31 2·19 2·29 2·15 2·29 2·27 2·39 2·41 2·10 
Lu 0·31 0·31 0·39 0·33 0·39 0·36 0·34 0·36 0·33 0·35 0·37 0·34 0·35 0·33 0·35 0·36 0·38 0·38 0·33 
Sample: 8159 8164 6110 8129 6108 6158 8144 8112 8101 676 875 870 869 8140 6140 894 8136 672 898 
Series:
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Low-K
 
Main
 
Main
 
Main
 
Main
 
Low-K
 
Low-K
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
SiO2 45·41 44·94 43·18 44·67 44·88 42·72 44·35 45·38 45·84 45·09 45·12 46·39 46·65 44·87 44·46 44·29 45·05 45·49 43·97 
TiO2 2·08 2·05 2·20 2·27 2·31 2·25 2·37 2·21 2·23 2·02 2·36 2·19 2·28 2·27 2·38 2·32 2·34 2·09 2·15 
Al2O3 13·25 13·63 13·55 14·11 13·55 14·15 13·95 14·93 14·30 14·17 15·10 13·92 14·78 14·17 14·70 14·13 14·74 14·94 14·44 
Fe2O3* 12·27 12·64 12·77 12·47 12·63 12·38 12·77 11·24 11·98 11·99 11·41 13·18 11·47 11·58 11·65 11·94 11·48 11·80 11·15 
MnO 0·17 0·18 0·20 0·20 0·18 0·20 0·19 0·18 0·18 0·19 0·17 0·19 0·19 0·18 0·18 0·17 0·17 0·18 0·17 
MgO 12·52 12·32 12·26 11·75 11·74 11·70 11·51 11·47 11·24 11·18 11·11 11·09 10·98 10·96 10·91 10·81 10·76 10·65 10·62 
CaO 10·56 9·66 9·40 11·41 9·84 11·32 10·91 10·38 10·82 10·34 10·71 10·46 10·70 10·54 10·73 10·91 11·08 10·87 11·35 
Na22·69 2·98 3·30 2·96 3·16 3·06 2·88 3·28 3·14 3·05 3·01 2·51 3·76 3·27 3·20 3·24 3·26 2·91 2·82 
K21·35 1·47 1·55 1·59 1·59 1·60 1·21 1·58 1·58 1·48 1·53 0·75 0·94 1·60 1·67 1·56 1·60 1·49 1·55 
P2O5 0·47 0·49 0·48 0·47 0·51 0·36 0·51 0·51 0·49 0·42 0·47 0·44 0·44 0·48 0·47 0·51 0·43 0·48 0·50 
Total 100·77 100·37 98·88 101·89 100·39 99·75 100·66 101·16 101·79 99·91 100·99 101·13 102·19 99·92 100·35 99·88 100·91 100·90 99·90 
LOI 1·09 0·08 gain 0·34 gain 0·01 0·55 0·13 0·26 gain 0·28 0·85 0·11 0·26 0·01 0·62 0·33 gain 1·18 
Mg-no. 66·90 65·88 65·54 65·12 64·81 65·18 64·10 66·90 65·02 64·88 65·86 62·50 65·47 65·22 64·98 64·20 65·00 64·13 65·37 
Ba 380 336 511 470 479 711 406 466 398 455 349 345 455 384 493 413 483 488 505 
Cr 375 397 409 378 326 348 324 291 398 383 312 352 325 408 335 314 369 326 367 
Hf 3·68 3·73 4·67 4·43 4·81 3·98 3·94 4·81 3·94 4·00 4·44 3·83 4·25 4·18 4·36 4·10 4·39 4·31 4·01 
Nb 43 44 52 58 53 54 53 55 50 43 53 40 49 50 53 56 49 48 53 
Ni 304 320 332 219 285 223 234 217 245 220 198 222 192 235 199 217 178 184 216 
Pb 0·81 1·89 2·40 2·30 1·56 2·19 1·04 1·49 1·24 1·70 1·23 0·76 2·26 2·22 2·18 1·15 1·26 1·53 1·26 
Rb 25 28 31 35 32 36 21 42 31 26 34 14 21 33 33 28 32 26 32 
Sc 28·7 27·8 29·5 27·7 29·3 34·1 32·5 32·2 33·9 33·0 32·0 30·6 31·7 32·1 32·3 33·3 34·3 32·6 31·5 
Sr 526 506 574 778 592 571 903 589 550 545 583 624 571 572 590 588 524 559 620 
Ta 2·93 2·83 3·42 3·95 3·61 3·62 3·52 3·91 3·35 2·96 3·57 2·69 3·44 3·34 3·55 3·85 3·49 3·29 3·60 
Th 2·86 3·13 3·95 4·41 3·81 4·25 3·40 4·40 3·51 3·18 3·90 2·53 4·23 3·95 4·16 4·02 3·88 3·64 3·71 
0·84 0·89 1·10 1·29 1·02 1·13 0·90 1·21 1·00 0·56 1·06 0·58 1·11 1·05 1·10 1·08 1·06 1·00 1·02 
211 197 228 221 223 237 240 241 236 226 238 213 215 257 231 238 270 227 228 
22 24 30 25 28 28 27 26 25 29 29 25 27 25 29 27 27 30 24 
Zn 78 91 88 84 78 82 84 73 79 78 73 84 65 74 72 78 71 79 74 
Zr 160 148 199 179 200 164 160 200 158 147 181 146 169 167 179 161 164 167 161 
La 27·7 27·1 33·9 36·7 34·9 32·0 31·5 36·2 30·4 29·3 33·4 26·3 33·1 31·1 32·7 34·0 32·7 31·6 33·4 
Ce 53·7 52·8 65·2 68·8 68·1 60·8 61·2 68·8 58·4 57·3 64·0 52·9 62·1 58·9 62·2 64·6 62·3 61·6 63·7 
Pr 7·05 6·69 8·18 8·51 8·87 7·61 7·96 8·73 7·51 7·64 8·21 7·08 7·64 7·40 7·70 8·30 8·04 8·07 8·12 
Nd 29·0 27·9 33·6 33·5 35·8 31·3 32·7 34·5 30·3 31·0 32·9 29·4 31·2 30·2 31·5 33·4 32·3 33·0 32·6 
Sm 5·92 5·70 6·78 6·52 7·20 6·37 6·59 6·81 6·20 6·41 6·63 6·17 6·30 6·13 6·39 6·77 6·50 6·70 6·42 
Eu 1·93 1·89 2·18 2·15 2·26 2·03 2·11 2·11 1·98 1·99 2·11 1·99 2·03 1·98 2·09 2·16 2·08 2·10 2·06 
Gd 5·66 5·73 6·66 6·24 6·62 6·22 6·23 6·21 5·88 6·04 6·25 5·96 6·29 6·18 6·50 6·26 6·13 6·24 5·92 
Tb 0·86 0·86 0·99 0·91 1·01 0·92 0·94 0·93 0·89 0·92 0·95 0·92 0·93 0·89 0·94 0·96 0·94 0·96 0·90 
Dy 4·71 4·63 5·35 4·90 5·55 5·05 5·07 5·12 4·85 5·07 5·24 5·09 5·05 4·88 5·13 5·23 5·27 5·36 4·87 
Ho 0·90 0·91 1·04 0·94 1·07 0·97 0·97 0·98 0·93 0·99 1·00 0·98 0·99 0·94 1·00 1·00 1·03 1·03 0·93 
Er 2·24 2·26 2·64 2·36 2·74 2·49 2·43 2·52 2·34 2·49 2·55 2·47 2·51 2·40 2·52 2·55 2·63 2·65 2·36 
Tm 0·35 0·35 0·41 0·37 0·44 0·39 0·38 0·39 0·36 0·39 0·40 0·38 0·39 0·38 0·39 0·39 0·42 0·42 0·36 
Yb 2·00 2·04 2·44 2·15 2·50 2·30 2·14 2·28 2·06 2·20 2·31 2·19 2·29 2·15 2·29 2·27 2·39 2·41 2·10 
Lu 0·31 0·31 0·39 0·33 0·39 0·36 0·34 0·36 0·33 0·35 0·37 0·34 0·35 0·33 0·35 0·36 0·38 0·38 0·33 
Sample: 8160 675 6151 8111 6103 671 8161 674 8139 6104 8138 6155 6143 879 678 895 6187 6157 888 860 
Series:
 
Main
 
Main
 
Main
 
Low-K
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
SiO2 43·88 46·15 45·29 45·32 44·95 45·24 45·12 47·33 46·54 45·35 43·84 45·48 44·65 44·74 47·71 45·52 45·59 46·52 44·25 43·89 
TiO2 2·34 1·98 2·34 2·13 2·47 2·14 2·43 2·13 2·52 2·44 2·46 2·42 2·54 2·41 2·19 2·58 2·52 2·29 2·32 2·40 
Al2O3 14·30 14·33 14·18 14·74 15·28 14·99 14·81 14·59 15·31 15·03 15·07 15·15 14·97 14·97 14·35 15·34 14·99 15·44 15·33 15·29 
Fe2O3* 11·77 11·98 11·35 11·39 11·08 11·88 12·27 12·17 11·01 11·67 11·68 11·45 12·05 11·59 11·49 10·62 11·31 11·14 11·43 11·42 
MnO 0·18 0·19 0·16 0·19 0·17 0·19 0·17 0·19 0·18 0·18 0·17 0·18 0·19 0·18 0·19 0·15 0·16 0·18 0·18 0·18 
MgO 10·61 10·54 10·49 10·46 10·36 10·31 10·07 10·06 10·04 9·94 9·83 9·66 9·48 9·36 9·32 9·32 9·18 9·16 9·15 9·14 
CaO 10·38 9·96 11·06 10·74 10·05 10·96 10·00 10·54 10·54 10·06 10·79 10·02 10·45 9·88 9·02 10·08 10·34 10·03 10·36 11·49 
Na23·33 3·05 3·20 3·30 3·86 2·70 3·52 2·88 3·66 3·62 3·81 3·42 3·42 4·05 3·53 3·93 3·47 3·59 3·50 3·99 
K21·90 1·40 1·74 1·06 1·97 1·55 1·63 1·35 1·78 1·78 1·01 1·74 1·68 0·88 1·88 2·07 1·83 1·81 1·99 0·94 
P2O5 0·55 0·43 0·54 0·44 0·62 0·48 0·59 0·46 0·56 0·70 0·51 0·50 0·54 0·54 0·54 0·63 0·56 0·51 0·55 0·60 
Total 99·23 100·02 100·35 99·77 100·81 100·44 100·61 101·70 102·14 100·77 99·70 99·67 99·70 98·92 100·25 100·24 99·95 100·34 99·54 100·06 
LOI 0·93 gain gain 0·40 gain gain gain 0·32 0·47 0·61 0·55 gain gain 0·32 gain 0·17 0·81 gain 0·48 0·72 
Mg-no. 64·10 63·54 64·68 64·53 64·94 63·22 61·92 62·09 64·37 62·79 62·51 62·57 60·90 61·53 61·64 63·48 61·66 61·97 61·31 61·34 
Ba 414 413 553 473 473 488 402 457 494 471 610 452 528 454 561 512 505 425 497 552 
Cr 379 347 349 426 215 314 266 308 330 252 248 266 247 269 279 270 223 255 238 207 
Hf 4·37 4·06 4·52 4·46 4·98 4·35 4·18 4·25 4·81 4·89 4·75 4·28 4·49 4·62 5·02 5·14 4·93 4·60 4·79 4·61 
Nb 56 41 56 54 61 49 54 44 63 61 57 52 57 56 57 68 59 53 64 65 
Ni 225 207 219 229 197 174 198 183 178 171 169 163 154 150 230 168 155 150 163 149 
Pb 1·18 1·31 1·31 1·56 1·48 1·66 2·02 1·60 1·46 1·67 1·38 1·53 1·23 0·91 5·01 1·88 1·49 1·75 1·74 1·50 
Rb 31 26 35 39 39 27 31 25 40 41 12 32 28 25 44 43 37 36 46 20 
Sc 31·6 30·7 33·4 33·1 30·0 32·0 28·5 30·7 31·8 27·7 32·8 28·9 31·2 29·8 24·1 29·7 28·6 28·9 30·5 32·6 
Sr 632 533 613 596 667 580 658 555 671 701 666 585 643 659 626 728 663 605 665 716 
Ta 3·83 2·83 3·83 3·73 4·16 3·40 3·58 3·04 4·28 4·12 3·88 3·46 3·80 3·76 2·11 4·68 3·97 3·70 4·55 4·38 
Th 3·71 3·18 3·86 4·27 4·54 3·72 3·83 3·41 4·73 4·51 4·12 3·77 3·73 4·47 5·20 5·12 4·48 4·10 5·24 4·64 
1·15 0·88 1·05 1·16 1·34 0·97 1·11 0·91 1·26 1·25 1·12 1·07 0·98 1·08 1·50 1·46 1·22 1·08 1·49 1·13 
230 211 221 236 236 227 217 216 221 195 298 211 218 237 197 209 199 211 223 263 
26 30 25 26 26 32 26 30 27 28 26 30 31 29 34 30 29 25 27 27 
Zn 78 81 72 72 69 81 82 86 72 71 74 68 78 76 81 70 69 72 73 76 
Zr 173 163 183 185 203 172 169 171 199 201 193 175 180 191 221 210 198 182 197 185 
La 34·8 28·0 34·6 35·2 38·9 32·3 33·2 30·2 40·1 39·2 36·5 31·9 34·3 36·2 35·7 43·0 38·3 34·3 40·6 40·3 
Ce 66·9 55·4 66·5 66·1 73·6 62·8 63·4 59·1 75·9 75·3 70·4 62·1 66·0 68·2 68·3 81·0 73·2 66·4 75·4 76·7 
Pr 8·57 7·37 8·55 8·32 9·40 8·29 7·85 7·81 9·57 9·64 9·01 8·10 8·58 8·37 8·63 10·23 9·34 8·58 9·46 9·78 
Nd 34·4 30·1 34·4 32·8 37·2 33·5 32·3 32·0 37·4 38·6 35·7 32·9 34·9 33·6 34·39 39·9 37·0 34·5 37·0 38·7 
Sm 6·88 6·20 6·90 6·49 7·28 6·84 6·46 6·67 7·24 7·50 7·12 6·64 7·06 6·81 6·87 7·70 7·35 6·92 7·15 7·47 
Eu 2·20 1·94 2·20 2·07 2·30 2·13 2·13 2·07 2·28 2·34 2·26 2·10 2·26 2·19 2·09 2·43 2·32 2·18 2·24 2·36 
Gd 6·34 5·88 6·38 6·02 6·59 6·44 6·44 6·29 6·60 6·84 6·58 6·31 6·65 6·68 6·34 6·94 6·72 6·51 6·45 6·76 
Tb 0·95 0·90 0·95 0·92 0·99 0·98 0·95 0·96 0·98 1·02 0·99 0·95 0·99 0·98 0·97 1·03 1·02 0·98 0·96 1·01 
Dy 5·12 5·03 5·23 5·09 5·44 5·48 5·07 5·35 5·37 5·56 5·46 5·29 5·41 5·38 5·43 5·73 5·66 5·42 5·21 5·49 
Ho 0·98 0·97 0·99 0·98 1·04 1·07 0·97 1·02 1·03 1·06 1·04 1·01 1·03 1·05 1·08 1·09 1·09 1·04 0·99 1·04 
Er 2·43 2·52 2·52 2·52 2·61 2·70 2·42 2·65 2·62 2·68 2·66 2·54 2·59 2·63 2·83 2·78 2·76 2·65 2·53 2·63 
Tm 0·38 0·39 0·39 0·39 0·42 0·42 0·36 0·41 0·41 0·42 0·42 0·40 0·40 0·40 0·43 0·43 0·43 0·41 0·39 0·41 
Yb 2·15 2·25 2·25 2·29 2·36 2·47 2·14 2·42 2·38 2·41 2·38 2·30 2·32 2·37 2·70 2·50 2·49 2·39 2·28 2·35 
Lu 0·33 0·36 0·35 0·37 0·37 0·39 0·33 0·38 0·37 0·38 0·38 0·37 0·36 0·38 0·44 0·40 0·39 0·38 0·37 0·37 
Sample: 8160 675 6151 8111 6103 671 8161 674 8139 6104 8138 6155 6143 879 678 895 6187 6157 888 860 
Series:
 
Main
 
Main
 
Main
 
Low-K
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
SiO2 43·88 46·15 45·29 45·32 44·95 45·24 45·12 47·33 46·54 45·35 43·84 45·48 44·65 44·74 47·71 45·52 45·59 46·52 44·25 43·89 
TiO2 2·34 1·98 2·34 2·13 2·47 2·14 2·43 2·13 2·52 2·44 2·46 2·42 2·54 2·41 2·19 2·58 2·52 2·29 2·32 2·40 
Al2O3 14·30 14·33 14·18 14·74 15·28 14·99 14·81 14·59 15·31 15·03 15·07 15·15 14·97 14·97 14·35 15·34 14·99 15·44 15·33 15·29 
Fe2O3* 11·77 11·98 11·35 11·39 11·08 11·88 12·27 12·17 11·01 11·67 11·68 11·45 12·05 11·59 11·49 10·62 11·31 11·14 11·43 11·42 
MnO 0·18 0·19 0·16 0·19 0·17 0·19 0·17 0·19 0·18 0·18 0·17 0·18 0·19 0·18 0·19 0·15 0·16 0·18 0·18 0·18 
MgO 10·61 10·54 10·49 10·46 10·36 10·31 10·07 10·06 10·04 9·94 9·83 9·66 9·48 9·36 9·32 9·32 9·18 9·16 9·15 9·14 
CaO 10·38 9·96 11·06 10·74 10·05 10·96 10·00 10·54 10·54 10·06 10·79 10·02 10·45 9·88 9·02 10·08 10·34 10·03 10·36 11·49 
Na23·33 3·05 3·20 3·30 3·86 2·70 3·52 2·88 3·66 3·62 3·81 3·42 3·42 4·05 3·53 3·93 3·47 3·59 3·50 3·99 
K21·90 1·40 1·74 1·06 1·97 1·55 1·63 1·35 1·78 1·78 1·01 1·74 1·68 0·88 1·88 2·07 1·83 1·81 1·99 0·94 
P2O5 0·55 0·43 0·54 0·44 0·62 0·48 0·59 0·46 0·56 0·70 0·51 0·50 0·54 0·54 0·54 0·63 0·56 0·51 0·55 0·60 
Total 99·23 100·02 100·35 99·77 100·81 100·44 100·61 101·70 102·14 100·77 99·70 99·67 99·70 98·92 100·25 100·24 99·95 100·34 99·54 100·06 
LOI 0·93 gain gain 0·40 gain gain gain 0·32 0·47 0·61 0·55 gain gain 0·32 gain 0·17 0·81 gain 0·48 0·72 
Mg-no. 64·10 63·54 64·68 64·53 64·94 63·22 61·92 62·09 64·37 62·79 62·51 62·57 60·90 61·53 61·64 63·48 61·66 61·97 61·31 61·34 
Ba 414 413 553 473 473 488 402 457 494 471 610 452 528 454 561 512 505 425 497 552 
Cr 379 347 349 426 215 314 266 308 330 252 248 266 247 269 279 270 223 255 238 207 
Hf 4·37 4·06 4·52 4·46 4·98 4·35 4·18 4·25 4·81 4·89 4·75 4·28 4·49 4·62 5·02 5·14 4·93 4·60 4·79 4·61 
Nb 56 41 56 54 61 49 54 44 63 61 57 52 57 56 57 68 59 53 64 65 
Ni 225 207 219 229 197 174 198 183 178 171 169 163 154 150 230 168 155 150 163 149 
Pb 1·18 1·31 1·31 1·56 1·48 1·66 2·02 1·60 1·46 1·67 1·38 1·53 1·23 0·91 5·01 1·88 1·49 1·75 1·74 1·50 
Rb 31 26 35 39 39 27 31 25 40 41 12 32 28 25 44 43 37 36 46 20 
Sc 31·6 30·7 33·4 33·1 30·0 32·0 28·5 30·7 31·8 27·7 32·8 28·9 31·2 29·8 24·1 29·7 28·6 28·9 30·5 32·6 
Sr 632 533 613 596 667 580 658 555 671 701 666 585 643 659 626 728 663 605 665 716 
Ta 3·83 2·83 3·83 3·73 4·16 3·40 3·58 3·04 4·28 4·12 3·88 3·46 3·80 3·76 2·11 4·68 3·97 3·70 4·55 4·38 
Th 3·71 3·18 3·86 4·27 4·54 3·72 3·83 3·41 4·73 4·51 4·12 3·77 3·73 4·47 5·20 5·12 4·48 4·10 5·24 4·64 
1·15 0·88 1·05 1·16 1·34 0·97 1·11 0·91 1·26 1·25 1·12 1·07 0·98 1·08 1·50 1·46 1·22 1·08 1·49 1·13 
230 211 221 236 236 227 217 216 221 195 298 211 218 237 197 209 199 211 223 263 
26 30 25 26 26 32 26 30 27 28 26 30 31 29 34 30 29 25 27 27 
Zn 78 81 72 72 69 81 82 86 72 71 74 68 78 76 81 70 69 72 73 76 
Zr 173 163 183 185 203 172 169 171 199 201 193 175 180 191 221 210 198 182 197 185 
La 34·8 28·0 34·6 35·2 38·9 32·3 33·2 30·2 40·1 39·2 36·5 31·9 34·3 36·2 35·7 43·0 38·3 34·3 40·6 40·3 
Ce 66·9 55·4 66·5 66·1 73·6 62·8 63·4 59·1 75·9 75·3 70·4 62·1 66·0 68·2 68·3 81·0 73·2 66·4 75·4 76·7 
Pr 8·57 7·37 8·55 8·32 9·40 8·29 7·85 7·81 9·57 9·64 9·01 8·10 8·58 8·37 8·63 10·23 9·34 8·58 9·46 9·78 
Nd 34·4 30·1 34·4 32·8 37·2 33·5 32·3 32·0 37·4 38·6 35·7 32·9 34·9 33·6 34·39 39·9 37·0 34·5 37·0 38·7 
Sm 6·88 6·20 6·90 6·49 7·28 6·84 6·46 6·67 7·24 7·50 7·12 6·64 7·06 6·81 6·87 7·70 7·35 6·92 7·15 7·47 
Eu 2·20 1·94 2·20 2·07 2·30 2·13 2·13 2·07 2·28 2·34 2·26 2·10 2·26 2·19 2·09 2·43 2·32 2·18 2·24 2·36 
Gd 6·34 5·88 6·38 6·02 6·59 6·44 6·44 6·29 6·60 6·84 6·58 6·31 6·65 6·68 6·34 6·94 6·72 6·51 6·45 6·76 
Tb 0·95 0·90 0·95 0·92 0·99 0·98 0·95 0·96 0·98 1·02 0·99 0·95 0·99 0·98 0·97 1·03 1·02 0·98 0·96 1·01 
Dy 5·12 5·03 5·23 5·09 5·44 5·48 5·07 5·35 5·37 5·56 5·46 5·29 5·41 5·38 5·43 5·73 5·66 5·42 5·21 5·49 
Ho 0·98 0·97 0·99 0·98 1·04 1·07 0·97 1·02 1·03 1·06 1·04 1·01 1·03 1·05 1·08 1·09 1·09 1·04 0·99 1·04 
Er 2·43 2·52 2·52 2·52 2·61 2·70 2·42 2·65 2·62 2·68 2·66 2·54 2·59 2·63 2·83 2·78 2·76 2·65 2·53 2·63 
Tm 0·38 0·39 0·39 0·39 0·42 0·42 0·36 0·41 0·41 0·42 0·42 0·40 0·40 0·40 0·43 0·43 0·43 0·41 0·39 0·41 
Yb 2·15 2·25 2·25 2·29 2·36 2·47 2·14 2·42 2·38 2·41 2·38 2·30 2·32 2·37 2·70 2·50 2·49 2·39 2·28 2·35 
Lu 0·33 0·36 0·35 0·37 0·37 0·39 0·33 0·38 0·37 0·38 0·38 0·37 0·36 0·38 0·44 0·40 0·39 0·38 0·37 0·37 
Sample: 6185 8134 695 8103 866 8107 6152 6100 699 6102 864 883 882 8157 865 8109 6130 8110 8128 6127 Calculated parental magma; see text 
Series:
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Low-K
 
Low-K
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 

 
SiO2 43·98 42·90 45·11 44·97 45·40 45·83 44·68 45·08 45·11 45·41 46·04 46·94 45·06 45·82 44·50 43·70 47·43 46·81 46·59 47·93 44·40 
TiO2 2·42 2·42 2·54 2·18 2·41 2·59 2·53 2·42 2·45 2·45 2·16 2·18 2·81 2·27 2·31 2·36 2·22 2·15 2·13 2·35 2·11 
Al2O3 15·08 14·85 15·78 15·03 15·81 15·68 15·57 15·44 15·52 15·65 14·97 15·33 15·75 15·88 15·41 14·90 15·72 15·51 15·66 17·62 13·31 
Fe2O3* 11·62 11·16 10·71 12·28 10·67 10·76 11·26 11·78 11·88 12·00 11·30 11·76 12·16 11·33 11·30 11·14 10·87 11·37 12·45 10·36 12·00 
MnO 0·19 0·18 0·18 0·17 0·17 0·16 0·18 0·19 0·19 0·19 0·18 0·20 0·19 0·18 0·18 0·18 0·16 0·17 0·20 0·17 0·17 
MgO 9·11 9·09 9·08 9·01 9·00 8·99 8·76 8·72 8·71 8·67 8·67 8·63 8·60 8·42 8·40 8·35 8·09 7·94 7·54 5·91 13·39 
CaO 10·69 9·76 10·25 10·58 9·88 10·10 9·51 9·99 9·53 9·45 10·72 10·67 9·64 10·41 10·09 11·21 9·53 9·58 9·19 9·62 9·88 
Na23·93 3·89 3·75 3·62 4·03 4·09 3·99 3·97 4·12 4·25 4·16 4·20 4·37 3·53 3·54 2·51 3·70 3·89 4·45 3·91 2·93 
K20·81 1·88 1·98 1·14 2·11 2·08 2·17 2·00 2·09 2·20 1·02 0·71 1·94 2·19 1·99 1·14 1·77 2·29 1·95 2·25 1·37 
P2O5 0·56 0·57 0·66 0·47 0·65 0·68 0·61 0·62 0·62 0·64 0·58 0·57 0·80 0·63 0·55 0·68 0·53 0·55 0·67 0·67 0·44 
Total 98·95 96·70 100·15 99·27 100·06 100·96 99·16 100·21 100·22 100·91 99·80 101·18 101·32 100·67 99·26 99·38 100·02 100·26 100·83 100·79 100·00 
LOI 0·57 0·10 0·13 gain gain gain –0·10 0·57 gain gain 0·81 0·79 0·07 1·03 1·00 3·22 gain 1·04 gain 0·70  
Mg-no. 60·82 61·74 62·68 59·23 62·56 62·34 60·64 59·46 59·22 58·87 60·32 59·25 58·35 59·55 59·58 59·76 59·59 58·04 54·54 53·05 68·85 
Ba 546 508 698 351 604 213 565 585 621 620 545 502 596 559 547 574 520 529 555 585  
Cr 226 209 182 399 209 213 201 136 136 134 355 262 147 190 227 275 182 259 184 55  
Hf 4·96 5·16 5·37 3·86 5·49 5·26 5·24 5·81 5·94 5·96 5·20 5·04 5·64 4·56 5·38 5·06 5·32 5·66 6·49 5·65  
Nb 62 63 73 42 71 70 68 71 72 74 65 59 74 62 59 72 51 62 70 65  
Ni 169 159 152 147 151 143 128 113 113 111 220 170 105 119 135 148 125 124 107 47  
Pb 1·50 2·62 1·74 0·83 1·92 1·63 1·52 3·14 2·18 3·52 2·14 1·78 1·93 1·53 2·19 1·55 2·06 2·43 2·45 3·49  
Rb 26 41 38 18 45 43 44 44 46 47 12 28 50 41 43 47 34 47 58 42  
Sc 31·8 28·9 28·8 31·7 28·4 28·6 28·0 27·0 27·5 26·9 29·0 29·4 25·7 26·3 28·5 26·7 26·5 25·5 23·3 23·8  
Sr 690 703 953 558 761 721 727 858 784 789 686 699 820 679 690 787 617 684 755 728  
Ta 4·21 4·42 4·88 2·84 4·98 4·83 4·55 4·67 4·85 4·86 4·77 4·39 5·05 4·08 4·38 5·28 3·54 4·59 5·13 4·34  
Th 4·64 5·35 5·46 2·72 5·53 5·38 5·16 5·90 5·60 6·18 5·74 4·94 5·46 4·58 5·07 6·04 4·24 5·62 5·91 5·40  
1·17 1·50 1·47 0·82 1·54 1·49 1·47 1·68 1·48 1·58 1·48 1·24 1·51 1·34 1·42 1·66 1·11 1·61 1·64 1·59  
246 229 206 236 192 228 227 207 220 212 225 226 212 217 202 154 206 195 189 191  
29 28 31 25 31 29 26 30 29 29 27 26 33 29 28 27 30 27 31 31  
Zn 75 77 73 82 71 69 76 79 78 74 75 75 78 70 75 65 69 80 89 69  
Zr 205 208 227 153 221 205 222 246 252 258 207 203 237 208 216 206 216 237 276 237  
La 39·2 41·6 45·3 26·5 45·8 45·0 42·2 44·9 45·9 47·5 43·8 41·3 46·4 38·5 42·5 46·9 36·8 44·3 50·2 42·3  
Ce 74·4 77·5 84·7 52·3 85·0 83·8 78·7 83·3 85·9 87·6 80·6 77·2 87·1 72·0 79·9 86·1 71·8 81·5 94·0 79·9  
Pr 9·46 9·49 10·57 6·92 10·60 10·56 9·94 10·04 10·83 10·58 10·06 9·68 11·01 8·94 10·02 10·64 9·32 10·02 11·72 9·94  
Nd 37·6 37·6 41·3 28·7 41·2 41·1 38·8 39·8 42·1 41·5 38·9 37·9 43·2 34·7 39·3 41·2 37·5 38·2 45·2 40·0  
Sm 7·32 7·26 7·78 6·02 7·77 7·79 7·47 7·54 7·99 7·85 7·36 7·22 8·29 6·61 7·51 7·67 7·48 7·16 8·42 7·84  
Eu 2·32 2·32 2·49 1·99 2·46 2·47 2·38 2·42 2·55 2·50 2·31 2·29 2·60 2·11 2·37 2·40 2·33 2·27 2·58 2·49  
Gd 6·71 6·94 7·01 5·79 7·10 7·02 6·74 7·15 7·11 7·47 6·53 6·45 7·36 5·94 6·70 6·82 6·98 6·33 7·34 7·61  
Tb 1·01 1·01 1·05 0·89 1·05 1·04 1·00 1·04 1·08 1·09 0·99 0·98 1·10 0·89 1·01 1·01 1·06 0·97 1·10 1·11  
Dy 5·49 5·49 5·70 4·87 5·69 5·73 5·52 5·64 5·89 5·86 5·36 5·31 5·90 4·83 5·54 5·49 5·93 5·30 5·97 6·18  
Ho 1·06 1·06 1·09 0·93 1·08 1·09 1·04 1·08 1·14 1·11 1·04 1·02 1·11 0·92 1·07 1·05 1·15 1·02 1·14 1·19  
Er 2·68 2·69 2·79 2·30 2·75 2·77 2·61 2·77 2·91 2·87 2·63 2·59 2·82 2·33 2·69 2·64 2·98 2·61 2·96 3·09  
Tm 0·42 0·42 0·43 0·36 0·43 0·43 0·42 0·44 0·45 0·45 0·41 0·41 0·44 0·36 0·42 0·42 0·47 0·41 0·46 0·49  
Yb 2·41 2·45 2·51 2·03 2·53 2·51 2·40 2·59 2·67 2·67 2·43 2·39 2·53 2·12 2·46 2·43 2·78 2·43 2·75 2·88  
Lu 0·39 0·39 0·40 0·31 0·40 0·40 0·38 0·41 0·44 0·42 0·39 0·38 0·40 0·33 0·39 0·38 0·44 0·39 0·44 0·45  
Sample: 6185 8134 695 8103 866 8107 6152 6100 699 6102 864 883 882 8157 865 8109 6130 8110 8128 6127 Calculated parental magma; see text 
Series:
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Low-K
 
Low-K
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 
Main
 

 
SiO2 43·98 42·90 45·11 44·97 45·40 45·83 44·68 45·08 45·11 45·41 46·04 46·94 45·06 45·82 44·50 43·70 47·43 46·81 46·59 47·93 44·40 
TiO2 2·42 2·42 2·54 2·18 2·41 2·59 2·53 2·42 2·45 2·45 2·16 2·18 2·81 2·27 2·31 2·36 2·22 2·15 2·13 2·35 2·11 
Al2O3 15·08 14·85 15·78 15·03 15·81 15·68 15·57 15·44 15·52 15·65 14·97 15·33 15·75 15·88 15·41 14·90 15·72 15·51 15·66 17·62 13·31 
Fe2O3* 11·62 11·16 10·71 12·28 10·67 10·76 11·26 11·78 11·88 12·00 11·30 11·76 12·16 11·33 11·30 11·14 10·87 11·37 12·45 10·36 12·00 
MnO 0·19 0·18 0·18 0·17 0·17 0·16 0·18 0·19 0·19 0·19 0·18 0·20 0·19 0·18 0·18 0·18 0·16 0·17 0·20 0·17 0·17 
MgO 9·11 9·09 9·08 9·01 9·00 8·99 8·76 8·72 8·71 8·67 8·67 8·63 8·60 8·42 8·40 8·35 8·09 7·94 7·54 5·91 13·39 
CaO 10·69 9·76 10·25 10·58 9·88 10·10 9·51 9·99 9·53 9·45 10·72 10·67 9·64 10·41 10·09 11·21 9·53 9·58 9·19 9·62 9·88 
Na23·93 3·89 3·75 3·62 4·03 4·09 3·99 3·97 4·12 4·25 4·16 4·20 4·37 3·53 3·54 2·51 3·70 3·89 4·45 3·91 2·93 
K20·81 1·88 1·98 1·14 2·11 2·08 2·17 2·00 2·09 2·20 1·02 0·71 1·94 2·19 1·99 1·14 1·77 2·29 1·95 2·25 1·37 
P2O5 0·56 0·57 0·66 0·47 0·65 0·68 0·61 0·62 0·62 0·64 0·58 0·57 0·80 0·63 0·55 0·68 0·53 0·55 0·67 0·67 0·44 
Total 98·95 96·70 100·15 99·27 100·06 100·96 99·16 100·21 100·22 100·91 99·80 101·18 101·32 100·67 99·26 99·38 100·02 100·26 100·83 100·79 100·00 
LOI 0·57 0·10 0·13 gain gain gain –0·10 0·57 gain gain 0·81 0·79 0·07 1·03 1·00 3·22 gain 1·04 gain 0·70  
Mg-no. 60·82 61·74 62·68 59·23 62·56 62·34 60·64 59·46 59·22 58·87 60·32 59·25 58·35 59·55 59·58 59·76 59·59 58·04 54·54 53·05 68·85 
Ba 546 508 698 351 604 213 565 585 621 620 545 502 596 559 547 574 520 529 555 585  
Cr 226 209 182 399 209 213 201 136 136 134 355 262 147 190 227 275 182 259 184 55  
Hf 4·96 5·16 5·37 3·86 5·49 5·26 5·24 5·81 5·94 5·96 5·20 5·04 5·64 4·56 5·38 5·06 5·32 5·66 6·49 5·65  
Nb 62 63 73 42 71 70 68 71 72 74 65 59 74 62 59 72 51 62 70 65  
Ni 169 159 152 147 151 143 128 113 113 111 220 170 105 119 135 148 125 124 107 47  
Pb 1·50 2·62 1·74 0·83 1·92 1·63 1·52 3·14 2·18 3·52 2·14 1·78 1·93 1·53 2·19 1·55 2·06 2·43 2·45 3·49  
Rb 26 41 38 18 45 43 44 44 46 47 12 28 50 41 43 47 34 47 58 42  
Sc 31·8 28·9 28·8 31·7 28·4 28·6 28·0 27·0 27·5 26·9 29·0 29·4 25·7 26·3 28·5 26·7 26·5 25·5 23·3 23·8  
Sr 690 703 953 558 761 721 727 858 784 789 686 699 820 679 690 787 617 684 755 728  
Ta 4·21 4·42 4·88 2·84 4·98 4·83 4·55 4·67 4·85 4·86 4·77 4·39 5·05 4·08 4·38 5·28 3·54 4·59 5·13 4·34  
Th 4·64 5·35 5·46 2·72 5·53 5·38 5·16 5·90 5·60 6·18 5·74 4·94 5·46 4·58 5·07 6·04 4·24 5·62 5·91 5·40  
1·17 1·50 1·47 0·82 1·54 1·49 1·47 1·68 1·48 1·58 1·48 1·24 1·51 1·34 1·42 1·66 1·11 1·61 1·64 1·59  
246 229 206 236 192 228 227 207 220 212 225 226 212 217 202 154 206 195 189 191  
29 28 31 25 31 29 26 30 29 29 27 26 33 29 28 27 30 27 31 31  
Zn 75 77 73 82 71 69 76 79 78 74 75 75 78 70 75 65 69 80 89 69  
Zr 205 208 227 153 221 205 222 246 252 258 207 203 237 208 216 206 216 237 276 237  
La 39·2 41·6 45·3 26·5 45·8 45·0 42·2 44·9 45·9 47·5 43·8 41·3 46·4 38·5 42·5 46·9 36·8 44·3 50·2 42·3  
Ce 74·4 77·5 84·7 52·3 85·0 83·8 78·7 83·3 85·9 87·6 80·6 77·2 87·1 72·0 79·9 86·1 71·8 81·5 94·0 79·9  
Pr 9·46 9·49 10·57 6·92 10·60 10·56 9·94 10·04 10·83 10·58 10·06 9·68 11·01 8·94 10·02 10·64 9·32 10·02 11·72 9·94  
Nd 37·6 37·6 41·3 28·7 41·2 41·1 38·8 39·8 42·1 41·5 38·9 37·9 43·2 34·7 39·3 41·2 37·5 38·2 45·2 40·0  
Sm 7·32 7·26 7·78 6·02 7·77 7·79 7·47 7·54 7·99 7·85 7·36 7·22 8·29 6·61 7·51 7·67 7·48 7·16 8·42 7·84  
Eu 2·32 2·32 2·49 1·99 2·46 2·47 2·38 2·42 2·55 2·50 2·31 2·29 2·60 2·11 2·37 2·40 2·33 2·27 2·58 2·49  
Gd 6·71 6·94 7·01 5·79 7·10 7·02 6·74 7·15 7·11 7·47 6·53 6·45 7·36 5·94 6·70 6·82 6·98 6·33 7·34 7·61  
Tb 1·01 1·01 1·05 0·89 1·05 1·04 1·00 1·04 1·08 1·09 0·99 0·98 1·10 0·89 1·01 1·01 1·06 0·97 1·10 1·11  
Dy 5·49 5·49 5·70 4·87 5·69 5·73 5·52 5·64 5·89 5·86 5·36 5·31 5·90 4·83 5·54 5·49 5·93 5·30 5·97 6·18  
Ho 1·06 1·06 1·09 0·93 1·08 1·09 1·04 1·08 1·14 1·11 1·04 1·02 1·11 0·92 1·07 1·05 1·15 1·02 1·14 1·19  
Er 2·68 2·69 2·79 2·30 2·75 2·77 2·61 2·77 2·91 2·87 2·63 2·59 2·82 2·33 2·69 2·64 2·98 2·61 2·96 3·09  
Tm 0·42 0·42 0·43 0·36 0·43 0·43 0·42 0·44 0·45 0·45 0·41 0·41 0·44 0·36 0·42 0·42 0·47 0·41 0·46 0·49  
Yb 2·41 2·45 2·51 2·03 2·53 2·51 2·40 2·59 2·67 2·67 2·43 2·39 2·53 2·12 2·46 2·43 2·78 2·43 2·75 2·88  
Lu 0·39 0·39 0·40 0·31 0·40 0·40 0·38 0·41 0·44 0·42 0·39 0·38 0·40 0·33 0·39 0·38 0·44 0·39 0·44 0·45  
*

Total Fe recorded as Fe2O3.

Analyses are in order of descending MgO content: those with MgO > 10·7% were used in the LKP and INVMEL models. Major elements analysed by XRF; trace elements Ba, Cr, Nb, Ni, Rb, Sr, V, Y, Zn and Zr analysed by XRF; other trace elements analysed by ICP-MS; see text for methods. See Electronic Appendix A for sample localities.

Fig. 2.

Mg-number of PVF olivine vs Mg-number of the rocks containing them. (a) Euhedral phenocrysts and groundmass crystals. Individual olivine analyses are short horizontal bars. Crosses for sample 8134 are olivines in a cluster also containing Al-rich clinopyroxenes and Mg–Al-rich spinels (see text). (b) Subhedral, rounded, irregular olivine macrocrysts. Individual olivine analyses are crosses. See text for details.

Fig. 2.

Mg-number of PVF olivine vs Mg-number of the rocks containing them. (a) Euhedral phenocrysts and groundmass crystals. Individual olivine analyses are short horizontal bars. Crosses for sample 8134 are olivines in a cluster also containing Al-rich clinopyroxenes and Mg–Al-rich spinels (see text). (b) Subhedral, rounded, irregular olivine macrocrysts. Individual olivine analyses are crosses. See text for details.

An unambiguous feature of many of these lavas is evident in Fig. 2b, which illustrates the compositions of the cores of the rounded macrocrysts. Although some of these are approximately in equilibrium with the bulk-rock compositions, most are much too Fe-rich for this to be so. Ulmer (1989) showed that Kd might rise to 0·34 if the olivine formed at 15 kbar, but Fig. 2b emphasizes that this cannot explain the most abundant macrocryst composition, around Fo75, let alone the Fo65 in sample 8140. These widespread, partly dissolved macrocrysts are relatively Fe-rich olivine xenocrysts within these lavas. The significance of this point will be discussed below.

Clinopyroxene

Euhedral clinopyroxene phenocrysts in the PVF lavas have a substantial range of compositions, generated by sector, radial and oscillatory zoning. Compositions range from: SiO2 = 41·3–50·1 wt %; Al2O3 = 4·1–11·4 wt %; TiO2 = 0·8–6·0 wt %; Na2O = 0·3–1·6 wt % (Table 1). The highest Ti and Al values characterize pink titanaugite phenocryst rims and groundmass grains. A section through a large clinopyroxene megacryst (683a) from the Kilbourne Hole breccia was also analysed. Its mean composition (Table 1) has 8·8 wt % Al2O3 but only 1·5 wt % TiO2 and 1·4 wt % Na2O. When calculated into pyroxene components, using the scheme proposed by Kushiro (1962), this megacryst has small amounts of acmite, Ti-Tschermak's (Ti-Ts) and Fe3+-Tschermak's (Fe-Ts; Fe3+ estimated by charge balance) components. No jadeite is calculated but the megacryst is rich in Ca-Tschermak's component (Ca-Ts). In contrast, the euhedral clinopyroxene phenocrysts in the typical Potrillo lava (683) pierced by the Kilbourne Hole maar have core compositions with little Ca-Ts and almost twice the Ti-Ts component of the megacryst. Ti-Ts rises to >15% in groundmass titanaugites within this lava. In alkalic lavas with similar bulk compositions, especially Si-saturation, Ca-Ts is sensitive to crystallization pressure (Thompson, 1974; Soesoo, 1997).

The Kilbourne Hole clinopyroxene megacryst has a lower CaO content than those of the Potrillo lava euhedral phenocryst or groundmass clinopyroxenes (Table 1). This is clear on a pyroxene Di–Hed–En–Fs quadrilateral plot (Fig. 3). If they both formed from similar basic alkalic melts, the megacryst crystallized at higher temperatures than the euhedral phenocrysts because it falls closer to the two-pyroxene solvus in the pyroxene system. In some lavas, the phenocrysts have rounded-irregular cores of clinopyroxene, with a greenish tint, that show signs of having been partly redissolved by the melt, producing a ‘fingerprint’ texture. Occasional <5 mm irregular to rounded fragments of clinopyroxene macrocrysts also occur. As apparent in Fig. 3, analyses of these disequilibrium clinopyroxene remnants overlap both the megacryst and phenocryst data. We therefore infer that these are remnants of partly dissolved sub-calcic clinopyroxene high-pressure megacrysts or other clinopyroxenes that also formed at relatively high pressures (e.g. precipitated onto the walls of magma conduits). Figure 3 also shows that the Kilbourne Hole megacryst and Ca-Ts-rich sub-calcic remnants are appreciably richer in Fe than the cores of euhedral Potrillo clinopyroxenes.

Fig. 3.

Compositional variations in PVF clinopyroxenes within samples 676, 683, 870, 879, 898, 8101, 8129, 8134 and 8138. The key to the symbols is given on the diagram.

Fig. 3.

Compositional variations in PVF clinopyroxenes within samples 676, 683, 870, 879, 898, 8101, 8129, 8134 and 8138. The key to the symbols is given on the diagram.

Amphibole

The amphibole megacrysts in PVF scoria were not subjected to microprobe analysis because Irving & Frey (1984) have already studied this mineral in detail and identified it as kaersutite. Their mean PVF amphibole megacryst major-element analysis is as follows: SiO2, 40·0; TiO2, 5·1; Al2O3, 14·5; FeO, 9·6; MnO, 0·11; MgO, 13·9; CaO, 11·1; Na2O, 2·4; K2O, 1·7 (all values in wt %).

Spinels

Three types of spinel occur in these samples (Table 1). Dark-green Mg–Fe–Al spinel (pleonaste) plots close to the spinel–hercynite join in Fig. 4b. It forms inclusions up to 5 mm in diameter within the Kilbourne Hole clinopyroxene megacryst. It also occurs as very small equant inclusions within a subhedral clinopyroxene phenocryst in sample 8134 and within a Fo74·5 olivine macrocryst in 8101. A loose cluster of subhedral olivine and clinopyroxene phenocrysts in 8134 includes olivines that appear to be a liquidus phase in this composition (marked as crosses in Fig. 2a). The greenish core of a coexisting clinopyroxene in the cluster contains a pleonaste grain. Brown chrome spinel occurs within the euhedral olivine phenocrysts of the lavas and titanomagnetite in their groundmasses (Table 1).

Fig. 4.

Compositional variations in PVF spinels within samples 683, 870, 879, 898, 8101, 8129, 8134, 8138 and 8140. The Al-rich spinels (pleonaste) form inclusions within the Kilbourne Hole clinopyroxene megacryst (see Fig. 3) and rarely within euhedral clinopyroxene and olivine phenocrysts in lavas. Cr-spinel is the most common opaque inclusion in olivine phenocrysts. Titanomagnetite occurs only in lava groundmasses.

Fig. 4.

Compositional variations in PVF spinels within samples 683, 870, 879, 898, 8101, 8129, 8134, 8138 and 8140. The Al-rich spinels (pleonaste) form inclusions within the Kilbourne Hole clinopyroxene megacryst (see Fig. 3) and rarely within euhedral clinopyroxene and olivine phenocrysts in lavas. Cr-spinel is the most common opaque inclusion in olivine phenocrysts. Titanomagnetite occurs only in lava groundmasses.

Plagioclase

Rounded megacrysts of plagioclase, up to 4 mm in size, occur in three of the analysed samples. Their compositions are in the range An27 to An48 (Table 1).

LAVA GEOCHEMISTRY

All lava samples with low loss-on-ignition (below 1·2 wt % in all but one instance and positive in about 30% of samples) were analysed by XRF at Birmingham University for a range of trace elements. From this set of 145 samples, subsets were analysed for major elements (XRF), at Birmingham University, and additional trace elements (INAA), at Durham University, to describe the full range of elemental compositions in greater detail. Subsequently, a subset of 60 samples was analysed by ICP-MS at Durham, using the methodology of Ottley et al. (2003), so that modelling of the rare earth elements (REE) could use high-quality data, with the shapes of the chondrite-normalized patterns fully delineated. These samples were mostly those with the highest abundances of MgO and Ni. One clinopyroxene and two amphibole megacrysts were also analysed by ICP-MS (Table 3).

Table 3:

ICP-MS analyses of PVS megacrysts

 Megacrysts with rounded polished margins (from pyroclastics; crystals in the 4–5 cm size range)
 
       
 Amphibole Amphibole  Three splits of one cpx megacryst from Kilbourne Hole
 
    

 
Pot1
 
878k
 
Average
 
683a-1
 
683a-2
 
683a-3
 
Average
 

 
K21·67 1·59 1·63 0·01 0·01 0·01 0·01  
TiO2 5·13 7·10 6·12 1·45 1·46 1·46 1·46 0·018 
MnO 0·11 0·10 0·11 0·13 0·13 0·13 0·13 0·004 
Ba 378 375 377 1·15 1·15 1·15 1·15 0·986 
Co 61 68 64 39 40 39 40 1·449 
Cr 29 28 28 5·9 6·5 6·7 6·4 0·770 
Cs 0·005 0·007 0·006 0·000 0·002 0·002 0·001 0·002 
Cu 23 48 35 25 27 25 26 3·106 
Ga 14 13 13 11 11 11 11 0·651 
Hf 2·35 1·90 2·13 2·14 2·15 2·17 2·15 0·035 
Nb 32·3 23·1 27·7 0·62 0·60 0·66 0·63 0·067 
Ni 62 113 88 139 148 143 144 9·064 
Pb 2·02 2·84 2·43 0·13 0·45 4·29 1·62 4·633* 
Rb 6·71 5·28 6·00 0·06 0·06 0·08 0·07 0·031 
Sc 36 49 42 66 67 68 67 1·703 
Sr 584 707 646 81 84 83 83 3·178 
Ta 1·93 1·40 1·66 0·15 0·15 0·14 0·15 0·010 
Th 0·04 0·04 0·04 0·06 0·05 0·06 0·06 0·011 
0·01 0·02 0·02 0·01 0·01 0·01 0·01 0·001 
339 485 412 372 377 378 375 6·091 
24·7 24·2 24·4 22·0 22·5 22·5 22·3 0·600 
Zn 85·9 106·6 96·3 29·5 31·4 30·3 30·4 1·924 
Zr 62·4 47·1 54·7 48·4 49·6 49·5 49·2 1·298 
La 5·86 4·39 5·13 2·71 2·78 2·79 2·76 0·085 
Ce 19·89 15·04 17·46 9·72 9·91 9·95 9·86 0·254 
Pr 3·30 2·59 2·95 1·64 1·68 1·66 1·66 0·032 
Nd 21·53 17·49 19·51 11·07 11·32 11·26 11·21 0·258 
Sm 6·02 5·36 5·69 3·60 3·64 3·65 3·63 0·050 
Eu 2·08 1·93 2·01 1·25 1·28 1·28 1·27 0·028 
Gd 6·72 6·32 6·52 4·54 4·64 4·61 4·59 0·103 
Tb 0·94 0·91 0·92 0·71 0·74 0·74 0·73 0·030 
Dy 5·02 4·89 4·96 4·15 4·24 4·24 4·21 0·104 
Ho 0·91 0·89 0·90 0·80 0·82 0·83 0·82 0·033 
Er 2·12 2·14 2·13 2·05 2·05 2·04 2·05 0·010 
Tm 0·30 0·30 0·30 0·32 0·32 0·32 0·32 0·005 
Yb 1·59 1·55 1·57 1·68 1·71 1·71 1·70 0·041 
Lu 0·23 0·23 0·23 0·25 0·26 0·26 0·26 0·010 
 Megacrysts with rounded polished margins (from pyroclastics; crystals in the 4–5 cm size range)
 
       
 Amphibole Amphibole  Three splits of one cpx megacryst from Kilbourne Hole
 
    

 
Pot1
 
878k
 
Average
 
683a-1
 
683a-2
 
683a-3
 
Average
 

 
K21·67 1·59 1·63 0·01 0·01 0·01 0·01  
TiO2 5·13 7·10 6·12 1·45 1·46 1·46 1·46 0·018 
MnO 0·11 0·10 0·11 0·13 0·13 0·13 0·13 0·004 
Ba 378 375 377 1·15 1·15 1·15 1·15 0·986 
Co 61 68 64 39 40 39 40 1·449 
Cr 29 28 28 5·9 6·5 6·7 6·4 0·770 
Cs 0·005 0·007 0·006 0·000 0·002 0·002 0·001 0·002 
Cu 23 48 35 25 27 25 26 3·106 
Ga 14 13 13 11 11 11 11 0·651 
Hf 2·35 1·90 2·13 2·14 2·15 2·17 2·15 0·035 
Nb 32·3 23·1 27·7 0·62 0·60 0·66 0·63 0·067 
Ni 62 113 88 139 148 143 144 9·064 
Pb 2·02 2·84 2·43 0·13 0·45 4·29 1·62 4·633* 
Rb 6·71 5·28 6·00 0·06 0·06 0·08 0·07 0·031 
Sc 36 49 42 66 67 68 67 1·703 
Sr 584 707 646 81 84 83 83 3·178 
Ta 1·93 1·40 1·66 0·15 0·15 0·14 0·15 0·010 
Th 0·04 0·04 0·04 0·06 0·05 0·06 0·06 0·011 
0·01 0·02 0·02 0·01 0·01 0·01 0·01 0·001 
339 485 412 372 377 378 375 6·091 
24·7 24·2 24·4 22·0 22·5 22·5 22·3 0·600 
Zn 85·9 106·6 96·3 29·5 31·4 30·3 30·4 1·924 
Zr 62·4 47·1 54·7 48·4 49·6 49·5 49·2 1·298 
La 5·86 4·39 5·13 2·71 2·78 2·79 2·76 0·085 
Ce 19·89 15·04 17·46 9·72 9·91 9·95 9·86 0·254 
Pr 3·30 2·59 2·95 1·64 1·68 1·66 1·66 0·032 
Nd 21·53 17·49 19·51 11·07 11·32 11·26 11·21 0·258 
Sm 6·02 5·36 5·69 3·60 3·64 3·65 3·63 0·050 
Eu 2·08 1·93 2·01 1·25 1·28 1·28 1·27 0·028 
Gd 6·72 6·32 6·52 4·54 4·64 4·61 4·59 0·103 
Tb 0·94 0·91 0·92 0·71 0·74 0·74 0·73 0·030 
Dy 5·02 4·89 4·96 4·15 4·24 4·24 4·21 0·104 
Ho 0·91 0·89 0·90 0·80 0·82 0·83 0·82 0·033 
Er 2·12 2·14 2·13 2·05 2·05 2·04 2·05 0·010 
Tm 0·30 0·30 0·30 0·32 0·32 0·32 0·32 0·005 
Yb 1·59 1·55 1·57 1·68 1·71 1·71 1·70 0·041 
Lu 0·23 0·23 0·23 0·25 0·26 0·26 0·26 0·010 
*

Megacryst contains sporadic sulphide inclusions. See Table 1 for major-element analysis of clinopyroxene 683a. Other major elements in a PVS amphibole megacryst are (in wt %): SiO2, 40·0; Al2O3, 14·5; FeO, 9·6; MnO, 0·11; MgO, 13·9; CaO, 11·1; Na2O, 2·4 (Irving & Frey, 1984).

Analytical methods

The trace elements analysed by XRF (Birmingham) were determined on a Philips PW 1400 spectrometer, using a Rh anode. Appropriate corrections were made for overlap of Rb Kβ on Y Kα, of Sr Kβ on Zr Kα and target contamination on Ni. The Rh Compton-scatter peak was used for absorption corrections. All samples were crushed in an agate ‘swing mill’ to ∼50 µm, and pelleted with a small amount of 7% aqueous PVA. Major-element analyses (Birmingham) were analysed using glass discs—a minimum of two for each sample. Synthetic standards were used (Leake & Hendry, 1969).

Rare earth elements Cr, Hf, Sc, Ta, Th and U were determined by instrumental neutron-activation analysis (INAA) at Durham. Samples were irradiated for five consecutive days at the University of London Reactor Centre at Silwood Park, in batches of nine samples interspaced with three standards to achieve satisfactory correction for variations in neutron flux. Standards were made by drying multi-element chemical solutions onto filter papers. For counting, samples were loaded onto a horizontal autochanger that accurately reproduced a sample-to-detector distance of 1·5–2·0 cm. Gamma-ray photons were collected by an EG&G® Ortec planar high-purity Ge low-energy photon detector (25 mm in diameter by 10 mm in thickness). Resolution of this instrument was excellent [nominally 550 eV full width at half maximum (FWHM) at 122 keV]. Gamma-rays were collected in the range 50–900 keV and processed as 4096-channel spectra. The efficiency of the detector was good up to 300–400 keV but declined dramatically in the energy range 400–920 keV. Nevertheless, good results were obtained using the following peaks with energies >300 keV: La (328·8 and 487·1 keV); Th (311·8 keV); Sc (889·4 keV). Leat et al. (1990) give results for international standards.

The current Durham technique for preparing and analysing samples by ICP-MS (Ottley et al., 2003) is not yet published in an international journal and therefore an abbreviated version is given here. The powdered samples, together with several blanks and a suitable standard, were prepared using a HF–HNO3 mixture (Romil SPA-grade) in a Savillex 22 ml PFA vial followed by evaporation to near dryness and further addition of HNO3. After driving off the initial HF–HNO3 mixture, samples were taken up in 1 ml of concentrated HNO3, ‘dried’ again and then a further 2·5 ml of concentrated HNO3 was added together with 10–15 ml of high-purity water, and then samples were left overnight on a hot plate. Following cooling, samples were further diluted to exactly 50 ml, to yield a solution in approximately 3·5% HNO

Sample and standard weight was always 0·1000 ± 0·001 g, in order to attempt to equalize the dissolved solids load and hence potential matrix effects. Any sample weight deviations were corrected for later. The final sample dilution was 1:5000. Re and Rh internal ‘spikes’ were added during the final dilution, to yield 20 ppb in solution. The use of Re and Rh as internal standards compensated for: (1) calibration drift because of changes in instrument sensitivity; (2) matrix suppression; (3) dilution errors.

Calibration of the Perkin Elmer Sciex Elan 6000 ICP-MS was achieved via the use of in-house standards and international reference materials (e.g. W-2, BHVO-1 and AGV1), together with procedural blanks. A blank and standard were routinely run every 10 samples to ascertain the magnitude of calibration drift. On the very few occasions that the RSD of the drift monitor exceeded 5% of the measured value, the samples were either re-analysed or a drift correction algorithm was incorporated into the data.

The isotopes selected for analysis were chosen to be free from isobaric interference wherever practical. Nevertheless, some of the REE in particular suffer from molecular oxide interferences. The magnitude of the interference correction is directly related to the plasma conditions and sample introduction system. All correction factors have been calculated relative to a CeO/Ce ratio of 2·6%. The Elan 6000 ICP-MS was optimized each day to obtain a 2·6% CeO/Ce ratio by changing the nebulizer gas flow.

Using these procedures, the elements determined showed good to excellent detection limits with all REE except Ce, being better than 5 ppb in the rock (Ottley et al., 2003). Routine detection limits (in the rock) of better than 10 ppb were obtained for most other trace elements of geological interest (LILE, HFSE). The reproducibility of elemental concentrations assessed by replicate runs of the same solution during an analysis period was usually better than 5% and frequently 3%. The external reproducibility of samples prepared from multiple separate dissolutions and run in separate analytical sessions were mostly better than 10% (RSD) and frequently better than 5% (2 SD) in typical basaltic rocks (Ottley et al., 2003).

Zr/Hf values for two international standards (NBS688 and BE-N) are between 37 and 38, and within error of the less precisely determined ‘accepted’ values. More importantly, the reproducibility (for separate dissolutions) of these values is excellent in each standard (1·2–1·7% RSD). Nb/Ta values given in the literature for Primitive Mantle and chondrites show more agreement than Zr/Hf, and are typically 17·3–17·6. External reproducibility for the two international standards (NBS688 and BE-N) was excellent (4·4 and 2·7% RSD, respectively) and shows good agreement with ‘accepted’ values. Both our values and the ‘accepted’ values are different from chondritic, but the good reproducibility indicates that this is probably because of a mineralogical factor rather than analytical problems (Ottley et al., 2003).

Sr and Nd isotopes were determined at McMaster University. Samples were leached with 6 M HCl overnight, then rinsed twice with high-purity de-ionized water before dissolution. The residues were dissolved in concentrated HF and HNO3 in Savillex ‘bomb’ vessels at 125°C for 3 days, before conversion to the chloride form for ion exchange. Standard cation and reverse-phase column separation methods were used. Analytical blanks were less than 0·02 ng for Nd (Sr not measured). Sr and Nd isotope analyses were performed on a VG isomass 354 mass-spectrometer. Sr was analysed by dynamic multi-collection, using a three-collector algorithm, and was normalized to a ratio of 0·1194 for 86Sr/88Sr. Nd was analysed using double filaments and a four-collector peak switching algorithm, and was normalized to a 146Nd/144Nd ratio of 0·7219. During this work, an average value of 0·710287 ± 35 (2σ population) was determined for the NBS 987 standard (n = 10), and an average value of 0·511860 ± 15 (2σ population) was determined for the La Jolla standard (n = 9). Sr isotope data were normalized to a value of 0·71024 for the NBS 987 standard.

Os isotopes were determined at the Arthur Holmes Isotope Geology Laboratory, Durham University. Whole-rock Re–Os chemical procedures followed the methods of Pearson & Woodland (2000). Digestions were performed in Carius tubes spiked for Re and Os. Total procedural Os blanks were 0·5 pg; all samples were blank corrected. Os samples were analysed by N-TIMS on a Thermo Electron Triton mass spectrometer. All analyses were carried out on the secondary electron multiplier, via ion-counting, in peak-hopping mode using a Ba(OH)2 activator solution. Using this procedure, our long-term mean 187Os/188Os value for 161 runs of the University of Maryland College Park standard, at signal sizes equivalent to those of the samples, was 0·11383 ± 32 (2 SD; 2·8 per mil) and is within error of the value of 0·113791 ± 3 produced from static Faraday runs of large loads by Walker et al. (1997). The mean 189Os/188Os over this period is 1·21976 ± 192 (2 SD; equates to 1·6 per mil). Re was analysed on a Thermo Electron Neptune mass spectrometer on the SEM, in peak-hoping mode, using Ir as a mass bias correction.

Element and oxide variation

Analyses of the lavas with ICP-MS data are given in Table 2; the remainder are in Electronic Appendix B. On a plot of total alkalis versus silica (TAS; Fig. 5), the Potrillo lavas classify as basanites, basalts and hawaiites (Le Maitre, 2002), in agreement with earlier analyses (Anthony et al., 1992). MgO in the Potrillo suite varies from 5·6 to 12·5 wt %, with three samples classifying as picrites (IUGS; MgO > 12 wt %; Le Bas, 2000). Our analyses of PVF lavas with MgO < 6·5 wt % are both from the Santo Tomas–Black Mountain area, east of the main PVF (Electronic Appendix A) and we have not studied them in detail (Hoffer, 1971; Anthony et al., 1992).

Fig. 5.

Variation of total alkalis (Na2O + K2O) with SiO2 in the PVF lavas. TAS subdivisions are from Le Maitre (2002). Filled squares are the Main Series; open diamonds are the Low-K Series (see text).

Fig. 5.

Variation of total alkalis (Na2O + K2O) with SiO2 in the PVF lavas. TAS subdivisions are from Le Maitre (2002). Filled squares are the Main Series; open diamonds are the Low-K Series (see text).

Clues to the nature of the subdivisions and fractionating assemblages in the PVF magmas are given by a plot of MgO versus various elements and oxides (Fig. 6). Strong positive correlations between MgO and both Ni and Cr (Fig. 6a) show that one or more ferromagnesian phases fractionated from this suite (the marked trend-lines are linear least-squares fits through non-inflected data arrays). In contrast, a plot of MgO versus K2O (Fig. 6b) shows that the overall PVF suite contains two series. About 85% of the lavas form the Main Series, in which K2O behaves incompatibly throughout the sampled MgO range (Fig. 6b). The remainder all have variable but distinctly lower K2O abundances (<1·3%), with a mean of ∼1·0% that does not vary as a function of MgO. This Low-K Series is distinguished from the Main Series in Fig. 5 and subsequent diagrams. Figure 6 is discussed in more detail below.

Fig. 6.

Variations of selected oxides and trace elements in the PVF lavas as functions of MgO. Filled squares are the Main Series; open diamonds are the Low-K Series (see text). Trend lines in (a), (b) and (h) are least-squares fits; in (c), (d), (e) and (f), the trends are inflected. The vectors for removal of 10% clinopyroxene and amphibole megacrysts use the major-element analyses of these phases in Table 1 and Irving & Frey (1984).

Fig. 6.

Variations of selected oxides and trace elements in the PVF lavas as functions of MgO. Filled squares are the Main Series; open diamonds are the Low-K Series (see text). Trend lines in (a), (b) and (h) are least-squares fits; in (c), (d), (e) and (f), the trends are inflected. The vectors for removal of 10% clinopyroxene and amphibole megacrysts use the major-element analyses of these phases in Table 1 and Irving & Frey (1984).

Isotopes

There are no published isotopic studies of PVF lavas where the elemental compositions of the samples are also recorded and MgO exceeds 9 wt %. As the Mg-rich lavas are the main focus of this study, nine samples were analysed for Sr and Nd isotopes, and four of these were analysed for Os isotopes. The results are given in Tables 4 and 5.

Table 4:

Strontium and neodymium isotopes in PVF lavas

Sample
 
Series
 
87Sr/86Sr
 
± (2SE)
 
143Nd/144Nd
 
± (2SE)
 
MgO (wt %)
 
8159 Main 0·70304 0·00002 0·512952 0·000011 12·52 
8164 Main 0·70310 0·00003 0·512928 0·000018 12·32 
6110 Main 0·70313 0·00002 0·512909 0·000012 12·26 
8129 Main 0·70300 0·00004 0·512983 0·000011 11·75 
8144 Low-K 0·70306 0·00007 0·512965 0·000020 11·51 
8112 Main 0·70308 0·00003 0·512969 0·000014 11·47 
870 Low-K 0·70310 0·00002 0·512940 0·000018 11·09 
869 Low-K 0·70311 0·00007 0·512920 0·000024 10·98 
8111 Low-K 0·70308 0·00003 0·512964 0·000020 10·46 
Sample
 
Series
 
87Sr/86Sr
 
± (2SE)
 
143Nd/144Nd
 
± (2SE)
 
MgO (wt %)
 
8159 Main 0·70304 0·00002 0·512952 0·000011 12·52 
8164 Main 0·70310 0·00003 0·512928 0·000018 12·32 
6110 Main 0·70313 0·00002 0·512909 0·000012 12·26 
8129 Main 0·70300 0·00004 0·512983 0·000011 11·75 
8144 Low-K 0·70306 0·00007 0·512965 0·000020 11·51 
8112 Main 0·70308 0·00003 0·512969 0·000014 11·47 
870 Low-K 0·70310 0·00002 0·512940 0·000018 11·09 
869 Low-K 0·70311 0·00007 0·512920 0·000024 10·98 
8111 Low-K 0·70308 0·00003 0·512964 0·000020 10·46 

See Table 2 for Rb, Sr, Sm and Nd abundances.

Table 5:

Osmium isotopes in PVF lavas

Sample
 
Series
 
Re, ppt
 
Os, ppt
 
187Re/188Os
 
187Os/188Os
 
± (2SE)
 
Gamma Os initial*
 
MgO (wt %)
 
6110 Main 322 208·8 7·44 0·14497 0·00016 13·6 12·26 
8159 Main 117 14·1 40·5 0·2372 0·0004 85·9 12·52 
870 Low-K 202 5·5 181 0·2972 0·0006 133 11·09 
8164 Main 140 4·1 169 0·3413 0·0008 167 12·32 
Sample
 
Series
 
Re, ppt
 
Os, ppt
 
187Re/188Os
 
187Os/188Os
 
± (2SE)
 
Gamma Os initial*
 
MgO (wt %)
 
6110 Main 322 208·8 7·44 0·14497 0·00016 13·6 12·26 
8159 Main 117 14·1 40·5 0·2372 0·0004 85·9 12·52 
870 Low-K 202 5·5 181 0·2972 0·0006 133 11·09 
8164 Main 140 4·1 169 0·3413 0·0008 167 12·32 
*

Calculated to 48·5 ka using parameters listed by Pearson et al. (2004). Bulk Earth 187Os/188Os taken as 0·1271.

Sr and Nd isotopes

Five Main Series and four Low-K Series Mg-rich lavas were analysed for Sr and Nd isotopes (Table 4). Their 87Sr/86Sr and 143Nd/144Nd ratios define a range that is scarcely larger than analytical precision but, nevertheless, shows a weak negative correlation between 87Sr/86Sr and 143Nd/144Nd (Fig. 7a, inset). The elemental differences between the Main and Low-K series are not reflected in their Sr–Nd isotopes, which occupy the same small field (Fig. 7a).

Fig. 7.

New isotopic data for PVF lavas and previously published results for mantle xenoliths erupted at Kilbourne Hole. (a) 87Sr/86Sr vs 143Nd/144Nd. The key to the symbols is given on the diagram. Inset shows the new PVF lava data at enlarged scale. MORB and OIB fields are from Gibson et al. (1995). Kilbourne Hole mantle xenoliths and their pyroxenite veins are from Roden et al. (1988). (b) Distribution of 187Os/188Os expressed as γOs. See Table 4 for PVF lava data. Kilbourne Hole mantle xenolith data are from Burton et al. (1999) and Meisel et al. (2001). (c) Osmium abundance versus 187Os/188Os for the PVF lavas. Sources of comparative data: oceanic OIB field, Hauri (2002); continental crust, Saal et al. (1998); Kilbourne Hole mantle xenoliths, Burton et al. (1999) and Meisel et al. (2001). Mixing curve is between PVF lava 6110 and average Kilbourne Hole mantle xenolith, taken as 3500 ppt (see text).

Fig. 7.

New isotopic data for PVF lavas and previously published results for mantle xenoliths erupted at Kilbourne Hole. (a) 87Sr/86Sr vs 143Nd/144Nd. The key to the symbols is given on the diagram. Inset shows the new PVF lava data at enlarged scale. MORB and OIB fields are from Gibson et al. (1995). Kilbourne Hole mantle xenoliths and their pyroxenite veins are from Roden et al. (1988). (b) Distribution of 187Os/188Os expressed as γOs. See Table 4 for PVF lava data. Kilbourne Hole mantle xenolith data are from Burton et al. (1999) and Meisel et al. (2001). (c) Osmium abundance versus 187Os/188Os for the PVF lavas. Sources of comparative data: oceanic OIB field, Hauri (2002); continental crust, Saal et al. (1998); Kilbourne Hole mantle xenoliths, Burton et al. (1999) and Meisel et al. (2001). Mixing curve is between PVF lava 6110 and average Kilbourne Hole mantle xenolith, taken as 3500 ppt (see text).

Os isotopes

The most striking feature of the results (Table 5) is the wide range of Os contents within lavas having a relatively narrow range of MgO contents (11·1–12·5 wt %). Thus, one lava (6110; Main Series) has an Os content (209 ppt) that might be expected for a picrite elsewhere (e.g. Hauri, 2002), whereas the Os contents of the other lavas are much lower (4–14 ppt). Figure 7b and c show that the low-Os lavas have a much higher 187Os/188Os than sample 6110. As with Sr–Nd isotopes, the Low-K Series sample (870) has a similar value of 187Os/188Os to the two Main Series lavas with low Os contents.

DISCUSSION OF PVF LAVA GEOCHEMISTRY

What minerals fractionated from the magmas: the phenocrysts, megacrysts or both?

At the outset of this discussion, it is necessary to establish what process was predominant in causing the large range of chemical variation in the PVF lavas. Because the Low-K lavas are only a small proportion of the sample set, their genesis will be considered separately below. It is first necessary to identify what process can cause MgO to vary continuously from 12·5 to 5·9 wt % (Table 2) with equally wide ranges of Ni (332–47 ppm) and Cr (409–55 ppm). Although both mantle melting and fractional crystallization processes could cause such variation in Mg-rich lavas, the Mg–Ni–Cr-poor PVF lavas clearly require fractional crystallization from more magnesian parent magmas to be the main factor in their genesis. Furthermore, all the PVF lavas are in equilibrium with olivine far less forsteritic than any plausible mantle composition (Fig. 2). The distinctive patterns of variation in Fig. 6, where some trends are straight and others inflected, gives clues as to which phases separated from the magmas.

Phenocrysts. Olivine is a euhedral phenocryst in all the porphyritic PVF lavas; alone in samples with >10·7 wt % MgO and accompanied by euhedral, sector-zoned clinopyroxene in those with 7·0–10·7 wt % MgO. If olivine alone fractionated from PVF lavas with >10·7 wt % MgO, and was joined by clinopyroxene subsequently, the types of plot forming Fig. 6 should show inflections in trends of CaO and Sc vs MgO, as they do (Fig. 6c and d). Separation of plagioclase in the low-Mg lavas is probably marked by the slight inflection in the Al2O3 trend at 7–8 wt % MgO but there are insufficient data for this to be clear. Euhedral plagioclase phenocrysts occur only in lavas with MgO < 7 wt %. Rounded, relatively sodic plagioclase macrocrysts occur in more magnesian lavas (Table 1), but these are unlikely to have precipitated from such basic magmas (Irving & Frey, 1984). Their significance will be discussed below.

Megacrysts. The reduction in both CaO and Sc contents of the magmas, with falling MgO, requires that a major fractionating phase was either clinopyroxene or amphibole (or both). The ∼11% MgO Main Series lavas contain 10·0–11·4% CaO and 34 ppm Sc (Table 2). These values cannot be compared quantitatively with those of the small euhedral PVF clinopyroxene phenocrysts because the latter are sector-zoned, with extremely internally variable compositions. Apart from olivine, the other obvious candidates for fractionating phases are the clinopyroxene and amphibole megacrysts (Tables 1 and 3).The megacryst clinopyroxene contains several percent of green Ti-poor Mg–Al–Fe spinel (Table 1) but none has been found within megacryst amphibole. The coarse, cumulate-textured ultramafic xenoliths (see above) show textural relationships, suggesting that small amounts of both olivine and Mg–Al–Fe spinel co-precipitated with clinopyroxene, whereas amphibole crystallized at lower temperatures and alone. Extensive removal of amphibole alone from ∼11 wt % MgO PVF magmas could reduce CaO and Sc (Table 3) in the residua but would not permit a concomitant increase in Al2O3 because the Al2O3 content of the amphibole (14·5%; Irving & Frey, 1984) and the ∼11 wt % MgO melts are approximately the same (Tables 2 and 3). Amphibole separation (mean TiO2 content by ICP-MS 6·1 wt %, Table 3) would also strongly deplete TiO2 in the residual melt. Vectors for removal of 10 wt % megacryst clinopyroxene and amphibole from a typical PVF Mg-rich basalt are marked on the Al2O3 and TiO2 diagrams of Fig. 6.

Both the lava mineralogy and chemistry, and the ultramafic inclusions described above closely resemble those of the Geronimo Volcanic Field, southern Arizona, about 200 km west of the PVF (Kempton et al., 1987; Kempton & Dungan, 1989). Kempton et al. (1987) modelled the fractionation of the Geronimo magmas in detail, using a major-element mass-balance approach. Their preferred model for hawaiite derivation from alkali basalt required fractionation of a 5:1 assemblage of megacryst aluminous clinopyroxene and Mg–Al–Fe spinel, plus minor olivine. Two features of the PVF lavas rule out a similar quantified fractional crystallization model for this suite: (1) the trends on such plots as Fig. 6 are all scattered and this is a characteristic of melts where repetitive episodes of magma mixing have taken place; (2) petrographic evidence for such mixing is clear from the abundant rounded/fritted macrocryst fragments in many thin sections. In each sample studied in detail (e.g. Fig. 2b), the rounded olivine macrocrysts are more Fe-rich than coexisting euhedral phenocrysts of this phase. This is consistent with magmatic plumbing beneath the PVF in which uprising magma batches frequently intercepted and mixed with the fractionated residua of previous batches. Thus, the composition of any specific sample is a function of three post-genesis processes: (1) fractional crystallization; (2) mixing between such liquids and others encountered during their uprise; (3) re-fusion of unknown amounts of previous crystal cumulates from earlier magma batches.

Kempton (1987; Kempton & Dungan, 1989) considered that the Geronimo aluminous clinopyroxene and Mg–Al–Fe spinel megacrysts precipitated at ‘moderate pressures’. Two thermobarometers have subsequently been proposed to calculate the P and T of Ca-rich clinopyroxene formation from its elemental composition (Putirka et al., 1996; Soesoo, 1997). The thermodynamic approach of Putirka et al. (1996) requires the equilibrium compositions of both clinopyroxene and coexisting melt. Therefore, it cannot be applied to the PVF because these euhedral–subhedral phenocrysts are sector-zoned and therefore not equilibrium phases (Thompson, 1972).

The empirical approach of Soesoo (1997) uses principal component analysis of the major-element compositions of clinopyroxenes crystallized from melting experiments to construct eigenvector grids for P and T. The eigenvectors for the mean composition of the Kilbourne Hole clinopyroxene megacryst 683a (Table 1) are XPT = 32·35 and YPT = −27·55. On Soesoo's plots, these correspond to P = 1·1–1·5 GPa and T ∼ 1210°C. Greenwood (2001) compared Soesoo's results with the experimental data of Thompson (1974) for alkalic melts and suggested that the Soesoo thermobarometer might underestimate P by ∼0·4 GPa and T by ∼40°C. Therefore, the Ca-Ts-rich clinopyroxene megacryst 683a may have formed at ∼1·7 GPa pressure. This corresponds to a depth of ∼58 km, if a 30 km sialic crustal thickness is used (Sinno et al., 1986; Bussod & Williams, 1991). Perhaps 50–60 km is a sufficiently cautious estimate; i.e. 10–20 km above the base of the lithosphere, as defined by Bussod & Williams (1991).

The green Mg–Al–Fe spinel forming <5 mm rounded inclusions within the ∼5 cm clinopyroxene megacryst 683a appears, on textural grounds, to have crystallized either before or together with the latter. The cumulus-textured ultramafic inclusions within the PVF volcanics likewise give textural evidence of a clinopyroxene–green spinel–olivine assemblage. Thus, we conclude that PVF magmas with MgO < 10·7 wt % have undergone fractional crystallization dominated by removal of aluminous sub-calcic augite, accompanied by Mg–Al–Fe spinel and olivine at a pressure of ∼1·7 GPa. This pressure agrees well with the value of ∼1·6 GPa calculated by Huckenholz et al. (1992) for equilibrium between clinopyroxene and amphibole, based on the PVF megacrysts analysed by Irving & Frey (1984).

The REE data for clinopyroxene and amphibole megacrysts (Table 3) may be used to test the hypothesis that these phases crystallized from melts of the same composition as the PVF magmas. On a chondrite-normalized REE plot (Fig. 8), the patterns of all the PVF lavas analysed using ICP-MS (Table 2) plot in a compact group. The REE patterns for the average Potrillo clinopyroxene and amphibole megacrysts (Table 3) are also shown. Bottazzi et al. (1999) investigated trace-element partitioning between a kaersutitic amphibole (TiO2 = 4·06 wt %) and an alkali basalt melt—a situation very similar to the Potrillo example. The melt calculated to be in equilibrium with the Potrillo amphibole megacryst (Fig. 8), using the kaersutite/melt REE partition coefficients of Bottazzi et al. (1999), has similar REE abundances to the PVF lavas.

Fig. 8.

Rare earth elements in PVF lavas (Table 2) and clinopyroxene and amphibole megacrysts (Table 3). The shaded field is the PVF lavas (both series). Filled stars and crosses show the REE contents of the melts calculated to be in equilibrium with the clinopyroxene and amphibole megacrysts, using the partition coefficients of Wood & Blundy (1997) and Bottazzi et al. (1999). In the Wood & Blundy model, dotted lines are values for 1·7 GPa; 1250°C (upper line) and 1·7 GPa; 1200°C (lower line). See text for details.

Fig. 8.

Rare earth elements in PVF lavas (Table 2) and clinopyroxene and amphibole megacrysts (Table 3). The shaded field is the PVF lavas (both series). Filled stars and crosses show the REE contents of the melts calculated to be in equilibrium with the clinopyroxene and amphibole megacrysts, using the partition coefficients of Wood & Blundy (1997) and Bottazzi et al. (1999). In the Wood & Blundy model, dotted lines are values for 1·7 GPa; 1250°C (upper line) and 1·7 GPa; 1200°C (lower line). See text for details.

The modelling of melts in equilibrium with a clinopyroxene megacryst is more complicated. Recent experimental studies of clinopyroxene/melt partition coefficients, using ion microprobe data, have produced very variable values, especially for the heavy REE (e.g. Hart & Dunn, 1993; Blundy et al., 1998). Therefore, we used the predictive thermodynamic model of Wood & Blundy (1997) to calculate the values of REE partition coefficients appropriate to the crystallization of the Kilbourne Hole clinopyroxene megacryst (Table 3) from a PVF basaltic melt. Figure 8 shows the calculated equilibrium melt REE pattern appropriate to the P and T estimated above using the method of Soesoo (1997). Additional patterns are given for equilibria at 1·7 GPa—our preferred pressure for the clinopyroxene megacryst formation—and both 1200 and 1250°C. All three patterns are very close to each other and also to the result obtained from the Potrillo amphibole megacrysts. It is clear from Fig. 8 that at the depths relevant to this study, the effects of P are much smaller that those of T in the Wood & Blundy model. It is concluded that the calculations shown in Fig. 8 support the hypothesis that the amphibole and clinopyroxene megacrysts in the PVF lavas are cognate.

Isotope geochemistry

Sr–Nd isotopes.Figure 7a shows that all the analysed PVF lavas have 87Sr/86Sr higher and 143Nd/144Nd lower than MORB, with values of these ratios falling in the same range as OIB. Recent alkali basalts and basanites in the western USA have similar Sr–Nd isotopic ratios (e.g. Wang et al., 2002, and references therein). The complete overlap between both 87Sr/86Sr and 143Nd/144Nd values in the Main and Low-K series samples (Fig. 7a, inset) removes the simplest explanation for these two suites, namely that they originated from different mantle sources. Possibly, they originated from the same mantle source and the differences between them developed subsequently. Alternatively, the mantle source might have varied mineralogically, with a K-bearing mineral (such as amphibole) remaining in the residue after partial melting in some places but not in others (see below). Furthermore, crustal contamination can also be ruled out as the cause of the Low-K Series because any such process, on a scale sufficient to make changes to their elemental compositions as large as those seen in Fig. 6, would be bound also to affect their Sr–Nd isotopic ratios. The ∼1·6 Ga acid crust sampled as xenoliths in the lava at Kilbourne Hole has values of 87Sr/86Sr up to 0·79 and 143Nd/144Nd down to 0·5119 (Padovani & Reid, 1989). The Sr–Nd isotopic compositions of the lithospheric mantle beneath the PVF are represented in Fig. 7a by analyses of peridotite xenolith samples (some with pyroxenite veins) from Kilbourne Hole (Roden et al., 1988). The majority of these have lower 87Sr/86Sr and higher 143Nd/144Nd than the PVF lavas. Nevertheless, two of these xenoliths have more radiogenic compositions than the Potrillo lavas and therefore Sr–Nd isotopic ratios alone cannot completely exclude the possibility that the sub-PVF lithospheric mantle has an isotopically variable composition that was averaged when the PVF magmas melted from it.

Os isotopes. Such a possibility can be evaluated in the light of the Os isotope results. Osmium is an ideal element for discriminating between the melting products of convecting and lithospheric mantle (e.g. Hauri, 2002). During the melting associated with lithosphere formation, Re is partitioned preferentially into the melt and Os into the residuum, thus creating large differences in Re/Os. Over geological time, as 187Re decays to 187Os, the ratio 187Os/188Os in a given volume of mantle peridotite rises at very variable rates, depending on its melting history and whether or not it has avoided the constant remixing associated with convection. Thus, geologically old lithospheric mantle characteristically has relatively low Re/Os and 187Os/188Os (e.g. Pearson et al., 2004). If this were to melt to produce basaltic magmas, they would inherit low 187Os/188Os ratios.

If the lithospheric mantle contains abundant geologically old pyroxenite veins, the situation is more complicated. In peridotite massifs such as Beni Bousera, Morocco, the pyroxenite veins have high enough Os contents and 187Os/188Os ratios that a subsequent basaltic melt from both the peridotite and its included pyroxenite veins could have significantly greater 187Os/188Os than a simple peridotite melt (Pearson & Nowell, 2004). Nevertheless, before speculating about a hypothesis of pyroxenite-veined lithospheric mantle fusion for PVF magmatism, it would seem prudent to find evidence that such radiogenic pyroxenite veins actually occur within the lithospheric mantle sampled at Kilbourne Hole. This suite of mantle xenoliths has been subjected to one of the most extensive Os-isotope studies of any non-cratonic, basalt-hosted example to date (Burton et al., 1999; Meisel et al., 2001), but, unfortunately, no pyroxenites were analysed. Nevertheless, even if such radiogenic pyroxenites from the Kilbourne Hole xenolith suite are reported in the future, it is important to remember that Bussod & Williams (1991) showed that such veins mostly came from the shallowest parts of the sub-PVF lithospheric mantle, in peridotites with pre-eruption pyroxene equilibration temperatures between 850 and 1000°C, in the 35–55 km depth range—well above the depths at which the Potrillo magmas appear to have originated (see below). Therefore, their relevance to PVF magma genesis would be minimal.

Figure 7b shows the ranges of 187Os/188Os (expressed as γOs) and the published data for Kilbourne Hole lithospheric mantle peridotites. The parameter γOs is used because this logarithmic-scale plot clarifies relationships in sample suites with 187Os/188Os close to Bulk Earth [using the value of Pearson & Nowell (2004)]. The γOs range of the PVF lavas is clearly higher than that of the Kilbourne Hole lithospheric mantle xenoliths, with no overlap between the groups. Meisel et al. (2001) calculated a model rhenium depletion age for the Kilbourne Hole lherzolite xenoliths of 1·6 Ga, comparable with the basement age of the local crust. This suggests that both crust and mantle formation in this area were results of a single 1·6 Ga event, and that no subsequent process in the lithospheric mantle disturbed the Os-isotope systematics of the latter.

The PVF samples (Table 5) include three from the Main Series and one from the Low-K Series, all with MgO contents among the highest found in our sample set. Three of the samples (two Main; one Low-K) have extremely low Os contents (4·1–14·1 ppt), far lower than could occur in an unmodified basic magma (MgO > 12 wt %) that had originated in equilibrium with mantle peridotite (∼45 ppt Os or more; Hauri, 2002). Clearly, despite their high MgO contents, these magmas have not proceeded without delay from their sources to the surface. Their low Os contents are combined with extremely high 187Os/188Os. This suggests that they paused in crustal reservoirs during their uprise and that this delay allowed two processes to take place: (1) a sulphide phase precipitated from the melts, so that most of the original Os in the melt was lost; (2) as a result, the Os-depleted melts became hypersensitive to contamination by continental crust with 187Os/188Os 0·5–2·5 (Saal et al., 1998). These three samples may be ‘failures’ in any attempt to decipher the original sources of the PVF magmas but they confirm our interpretation, detailed above, that even the MgO-rich members of this suite are affected by complex crystal–liquid differentiation and subtle crustal contamination.

The fourth sample, 6110, is the lava that erupted to form Potrillo maar (Fig. 1). This has exhumed a suite of mantle xenoliths, testifying to its uninterrupted uprise from mantle depths. Its Os content of 209 ppt is high enough to make it strongly resistant to crustal contamination and its 187Os/188Os value of 0·145 places it firmly in the OIB field (Fig. 7c) close to the maximum value of 0·150 (Hauri, 2002). A mixing curve is plotted in Fig. 7c between sample 6110 and average Kilbourne Hole lithospheric mantle peridotite, using an arithmetic mean of the Os isotope compositions published for Kilbourne Hole peridotites (Burton et al., 1999; Meisel et al., 2001). The 187Os/188Os range of these xenoliths is 0·117–0·130 and the average is 0·126. The Os concentration of this end-member was taken as 3500 ppt—typical of lithospheric mantle. This value was assumed, rather than calculated from the published analytical data, because peridotite xenoliths erupted in alkali basalts have clearly lost a substantial fraction of their original Os content because of suphide breakdown and alteration associated with eruption (Pearson et al., 2004). This is evident in the Kilbourne Hole suite, which has an average Os content of 2550 ppt, whereas the average for massif peridotites—residues from similar levels of melting—is 4000 ppt (Pearson et al., 2004). It is clear that sample 6110 could dissolve only a few percent of its peridotite xenolith load without developing both an extremely high Os abundance and a value of 187Os/188Os below the OIB range. This is 0·131–0·150, excluding several samples containing semi-dissolved mantle peridotite xenoliths from the database (e.g. Widom et al., 1999). This sensitivity would be even greater if one of the least-radiogenic peridotite 187Os/188Os values was used.

The Low-K Series

Figure 9 shows the normalized abundances in PVF lavas of a range of elements that are incompatible during anhydrous mantle fusion. The predominant Main Series lavas have coherent patterns on this plot, all closely resembling those of mildly alkalic OIB. The normalized patterns of all element abundances except K, Rb, U and Th in the Low-K Series lavas are within the same range as the Main Series. In two Low-K patterns, Sr is slightly relatively enriched (Fig. 9), as would be anticipated if these two lavas contained dissolved sodic plagioclase megacrysts (see above). Normalized U and Th abundances in the Low-K lavas differ little from those in the Main Series. K is, of course, deficient in all the Low-K lavas, with variable K/Nb. Rb relative abundances are extremely variable.

Fig. 9.

Normalized multi-element plots [see Thompson et al. (1984) for factors] for PVF lavas. Rounded plagioclase macrocrysts cause the small positive Sr anomalies in two Low-K Series samples. Analytical reproducibility is illustrated by duplicate analyses of a Namibian picrite (RNT & CJO, unpublished).

Fig. 9.

Normalized multi-element plots [see Thompson et al. (1984) for factors] for PVF lavas. Rounded plagioclase macrocrysts cause the small positive Sr anomalies in two Low-K Series samples. Analytical reproducibility is illustrated by duplicate analyses of a Namibian picrite (RNT & CJO, unpublished).

Having already shown that lavas of the two series have identical Sr–Nd isotopic ratios (Fig. 7a), the scope for hypothetical mechanisms that could generate the two contrasting PVF series is quite small. A mineral or minerals with relatively high partition coefficients for K, Rb, U and Th could be residual in the mantle sources of the Low-K but not the Main Series. Alternatively, such a phase could have separated from the Low-K melts during their fractional crystallization. But which mineral could be responsible? Discussion above of the variation of TiO2 in PVF lavas (Fig. 6) ruled out the megacryst-forming amphibole (Table 3) as a potential fractionating phase from any PVF magmas. Their K2O contents are also much too low for phlogopite saturation to occur at any pressure (e.g. Späth et al., 2001).

Therefore, it is logical to focus on hydrous mantle melting scenarios. However, even these present major difficulties as models for the origin of the PVF Low-K lavas. The behaviour of K in the Low-K Series is emphasized in Fig. 10. Although Na2O/K2O is virtually constant throughout the large MgO range of the Main Series (Fig. 10a), it is very scattered in the Low-K Series, possibly becoming progressively more variable with falling MgO. K/Nb is also essentially constant throughout the Main Series (Fig. 10b), but in the Low-K Series it falls with decreasing MgO. Although the average K2O content of the Low-K Series is constant, the amounts of this oxide vary by a factor of more than two at a given MgO content (Fig. 6b).

Fig. 10.

Na2O/K2O (a) and K/Nb (b) vs MgO for PVF lavas. These plots highlight the differences between the Main and Low-K Series.

Fig. 10.

Na2O/K2O (a) and K/Nb (b) vs MgO for PVF lavas. These plots highlight the differences between the Main and Low-K Series.

Superficially, the elemental chemistry of the PVF Low-K Series resembles that of the lavas of the Chyulu Hills, Kenya, modelled by Späth et al. (2001) as incremental partial melts of amphibole-bearing lherzolite, with amphibole present in the residuum. The forward modelling of Späth et al. (2001, fig. 12) for non-modal equilibrium batch melting of an amphibole lherzolite shows K2O abundances constant (buffered by residual amphibole) and K/Nb rising during progressive partial melting. The Low-K Series PVF samples show the same behaviour (Figs 6b and 10b) but in a suite of lavas with olivine phenocrysts (Fig. 2) having Mg-number far below values that could have been in equilibrium with any plausible peridotite mantle composition (e.g. Hirose & Kushiro, 1993; Falloon et al., 1997; Robinson et al., 1998). Thus, even though the Low-K Series magmas all appear to have undergone post-genesis fractional crystallization (see above), a ‘memory’ of geochemical parameters imposed by previous mantle melting processes is still preserved in the suite.

Figure 11 applies this reasoning to a diagram used by le Roex et al. (2001) to illustrate their concepts of genetic processes for the lavas of the Chyulu Hills and adjacent parts of the East African rift. The PVF analyses and both the East African rift data fields and the calculated mantle melting trends discussed by le Roex et al. (2001) are plotted in Fig. 11. At first sight, the PVF Low-K Series data fit the geochemical pattern of the Chyulu Hills magmatism well. Thus, the Low-K Series, such as the Chyulu Hills lavas, could be modelled by melting a hydrous mantle source with amphibole remaining in the residue. The problem with applying this concept to the Potrillo Low-K lavas, as noted earlier on the plot of K/Nb versus MgO (Fig. 10b), can be demonstrated in Fig. 11 by appending MgO content beside each Potrillo Low-K Series datapoint. The Potrillo Low-K lavas are clearly aligned along a trend of falling MgO with increasing Nb content. In contrast, a plot of K/Nb for the Chyulu Hills lavas (not shown; data from le Roex et al., 2001) shows points with very variable K/Nb in a 7–18% wt MgO range and no sign of any overall trend. If microprobe analyses of the olivine phenocrysts in the PVF lavas (Table 1; Fig. 2) were not available, it would be tempting to speculate that the progressive changes along the Potrillo Low-K Series trend in Fig. 11 occur because these are literally primary magmas, formed by melting of amphibole-bearing mantle and subsequently unmodified in any way. However, it is clear that these Low-K magmas have all undergone fractional crystallization (see above) and, as Anthony et al. (1992) showed, the PVF contains dozens of individual small volcanic centres, each evolving independently and spread over a large area (Fig. 1). The possibility that all these centres are fed from a single large underlying magma chamber can be discounted because convection within such a reservoir would eliminate the sorts of variations in incompatible elements that characterize the Low-K PVF lavas (Fig. 9). In such circumstances, it is quite reasonable that minor local events, affecting individual magma batches, caused the sorts of random minor variations in the normalized trace-element patterns of the Low-K Series lavas that occur in Fig. 9.

Fig. 11.

K/Nb vs Nb in PVF lavas. Filled squares are the Main Series; open diamonds are the Low-K Series (see text). Other fields and both mantle compositions and melting model trends (amphibole-bearing and amphibole-free lherzolites) are from le Roex et al. (2001). MgO contents are marked beside the Low-K Series PVF points. Only Main Series PVF lavas with >9% MgO are plotted.

Fig. 11.

K/Nb vs Nb in PVF lavas. Filled squares are the Main Series; open diamonds are the Low-K Series (see text). Other fields and both mantle compositions and melting model trends (amphibole-bearing and amphibole-free lherzolites) are from le Roex et al. (2001). MgO contents are marked beside the Low-K Series PVF points. Only Main Series PVF lavas with >9% MgO are plotted.

The Main Series

Because it was shown above that the lavas of the Main Series have undergone extensive fractional crystallization, only those with >9 wt % MgO are plotted in Fig. 11. Using the forward modelling of le Roex et al. (2001) for amphibole-free lherzolite, the PVF Main Series lavas appear to be small-fraction partial melts of mantle with K/Nb ∼ 260 and some variability in this ratio. Before modelling the genesis of this series in more detail, it is necessary to return to their fractional crystallization history and emphasize one further point. Figure 12 shows that La/Yb varies considerably at a given MgO content in the Main Series lavas but that, nevertheless, this ratio also increases with falling MgO.

Fig. 12.

La/Yb vs MgO for PVF lavas. Filled squares are the Main Series; open diamonds are the Low-K Series (see text). The vectors illustrate the effects of removing 10% of megacryst clinopyroxene or amphibole (Tables 2 and 3; Irving & Frey, 1984) from a Main Series lava.

Fig. 12.

La/Yb vs MgO for PVF lavas. Filled squares are the Main Series; open diamonds are the Low-K Series (see text). The vectors illustrate the effects of removing 10% of megacryst clinopyroxene or amphibole (Tables 2 and 3; Irving & Frey, 1984) from a Main Series lava.

There are two ways to explain this relationship: (1) it can be argued that clinopyroxene was prominent amongst the fractionating phases, as emphasized in the discussion of Fig. 6; (2) it could also be suggested that a suite of small-fraction lherzolite melts, such as these, will have primary variability in La/Yb because residual diopside in the source will affect the REE slopes of individual melt batches. Whatever the balance of processes causing the variation in La/Yb shown in Fig. 12, the fractional crystallization alternative cannot be ruled out. This means that even La/Yb cannot be taken as primary in Main Series lavas with MgO < 10·7 wt %. Therefore, only those lavas with higher MgO contents are used in REE modelling below. Before proceeding to such modelling, it is worth emphasizing that this study of post-genesis processes in the PVF magmas reaches a very different conclusion from those of many previous studies of basaltic suites. Because such an approach works well with MORB (Langmuir et al., 1992), a widespread view has developed that basaltic magmas with >8 wt % MgO are mostly saturated with only olivine and can therefore be restored to primary compositions by calculations that add this phase alone (e.g. Wang et al., 2002). In the PVF suites, it appears that melts with up to 10·7 wt % MgO fractionated clinopyroxene and also, judging from the rounded and fritted macrocrysts in their thin sections, they probably additionally redissolved various clinopyroxene populations during magma mixing events. It will also be shown below that the PVF magmas probably contained enough H2O to invalidate the assumption that this fractional crystallization occurred in anhydrous conditions. Asimow et al. (2004) have shown that even small magmatic H2O contents cause the fractionation paths of basic magmas to deviate substantially from anhydrous models.

MODELLING THE GENESIS OF THE MAIN SERIES

In previous sections, it was shown that all PVF Main Series magmas with <10·7 wt % MgO appear to have undergone fractional crystallization during their source-to-surface uprise and that this process involved a major fraction of a phase or phases additional to olivine; most probably, high-pressure aluminous sub-calcic augite with appreciable Na and Ti content. This brings uncertainty into any scheme for modelling magma genesis that relies on REE and other trace elements (e.g. McKenzie & O'Nions, 1991). The same problem also applies to any attempt to calculate the Na and Fe contents of the PVF parental melts from those of their fractionation residua, once one or more Na-bearing ferromagnesian phases is thought to be involved (see Wang et al., 2002). Fortunately, the PVF sample set includes 17 lavas with >10·7 % wt MgO; 14 in the Main Series and three in the Low-K Series (Table 2). The remainder of this discussion will concentrate on these because they appear to have fractionated only olivine (± traces of chromite) and the effects of this can be calculated. In particular, the Mg-rich basalts and picrites of the Main Series are suitable for petrogenetic modelling. The modelling below aims to use several approaches and to see whether they give similar results. First, the CIPW-normative compositions of the lavas will be compared with those of the melts (glasses) produced by isobaric phase-equilibria experiments on natural and synthetic peridotites. Next, the major and trace elements of the lavas will be modelled as products of decompression melting, using the approaches proposed by McKenzie & O'Nions (1991) and Langmuir et al. (1992).

Simple phase-equilibria model

The process of melt extraction from upwelling convecting mantle is now sufficiently well understood to show that any direct comparison of basalt (s.l.) compositions with those of experimental mantle melts is an oversimplification. Nevertheless, a plot of CIPW normative Di–Ol–Hy–(Ne + Lc)–Qz is illuminating (Fig. 13a). Experimental anhydrous partial melts of ‘fertile’ lherzolite mantle (KLB-1 and similar compositions) from recent studies (Hirose & Kushiro, 1993; Takahashi et al., 1993; Kushiro, 1996; Falloon et al., 1997, 1999; Hirschmann et al., 1998; Robinson et al., 1998; Walter, 1998) are plotted on this diagram. These authors have tended to emphasize small differences between their results, whereas Fig. 13a shows their overall coherence. The anhydrous isobaric melts at 1·0, 1·5 and 3 GPa define three separate trends of dry partial fusion between ∼3 and ∼40–45%. The data at 1 and 1·5 GPa are from more sources than those at 3 GPa and this probably explains why the 3 GPa trend is so much clearer than the others. Apart from overlap of the 1·0 and 1·5 GPa trends near the Di–Ol join (plane of critical undersaturation; Yoder & Tilley, 1962), the three trends are separate and show progressive migration of the melt compositions to higher olivine contents with increasing pressure.

Fig. 13.

CIPW normative diopside, olivine, hypersthene, nepheline and quartz in PVF and comparable samples. All data are calculated with 10% of total iron as Fe2O3 (Thompson & Gibson, 2000). Cotectics at 1 atm for the equilibrium [ol + plag + cpx + basaltic liquid] are from Thompson (1982) and the data of Sack et al. (1987); arrows mark directions of falling temperature. The continuous-line 1 atm cotectic is the best fit to the experimental data; all of the latter fall between the dashed lines. See key and text for explanations of symbols. Filled stars indicate mantle compositions used in the experimental studies. The outlined field encloses analyses of Potrillo lavas. (a) Compositions of the PVF lavas compared with those of glasses formed during experimental studies of fertile (KLB-1 and similar) lherzolite fusion under anhydrous conditions and with added H2O and CO2. Sources of the experimental data are: Hirose & Kushiro (1993); Takahashi et al. (1993); Hirose & Kawamoto (1995); Hirose (1996); Kushiro (1996); Falloon et al. (1997, 1999); Hirschmann et al. (1998); Robinson et al. (1998); Walter (1998). (b) Dredged glasses from Loihi Seamount and from the North Arch Volcanic Field, 200–300 km north of Oahu, Hawaii (Dixon et al., 1997; Dixon & Clague, 2001). Volatile contents of the glasses are summarized on the diagram. Experimental lherzolite melting data are simplified into generalized trends, showing the liquids produced by variable degrees of melting of anhydrous lherzolite, plus the contrasting effects of adding H2O or CO2 to the system. Some data points from (a) are reproduced in (b) as pale grey ‘ghosts’ to assist orientation. Note that the term ‘fertile’—as used by experimental petrologists—is not synonymous with the use of the same word, or of terms such as ‘Bulk Earth’, by isotope geochemists.

Fig. 13.

CIPW normative diopside, olivine, hypersthene, nepheline and quartz in PVF and comparable samples. All data are calculated with 10% of total iron as Fe2O3 (Thompson & Gibson, 2000). Cotectics at 1 atm for the equilibrium [ol + plag + cpx + basaltic liquid] are from Thompson (1982) and the data of Sack et al. (1987); arrows mark directions of falling temperature. The continuous-line 1 atm cotectic is the best fit to the experimental data; all of the latter fall between the dashed lines. See key and text for explanations of symbols. Filled stars indicate mantle compositions used in the experimental studies. The outlined field encloses analyses of Potrillo lavas. (a) Compositions of the PVF lavas compared with those of glasses formed during experimental studies of fertile (KLB-1 and similar) lherzolite fusion under anhydrous conditions and with added H2O and CO2. Sources of the experimental data are: Hirose & Kushiro (1993); Takahashi et al. (1993); Hirose & Kawamoto (1995); Hirose (1996); Kushiro (1996); Falloon et al. (1997, 1999); Hirschmann et al. (1998); Robinson et al. (1998); Walter (1998). (b) Dredged glasses from Loihi Seamount and from the North Arch Volcanic Field, 200–300 km north of Oahu, Hawaii (Dixon et al., 1997; Dixon & Clague, 2001). Volatile contents of the glasses are summarized on the diagram. Experimental lherzolite melting data are simplified into generalized trends, showing the liquids produced by variable degrees of melting of anhydrous lherzolite, plus the contrasting effects of adding H2O or CO2 to the system. Some data points from (a) are reproduced in (b) as pale grey ‘ghosts’ to assist orientation. Note that the term ‘fertile’—as used by experimental petrologists—is not synonymous with the use of the same word, or of terms such as ‘Bulk Earth’, by isotope geochemists.

The shape of the compositional trend of anhydrous isobaric melts produced by progressive mantle fusion is distinctive in Fig. 13a (Thompson, 1984). Initial (∼3%) anhydrous melts are nepheline (± leucite) normative. With rising T, the isobaric melts become increasingly less silica-undersaturated; the 1·0 GPa liquids approach silica saturation. The inflection as the melt compositions cross the Di–Ol join is an artefact of the projection. When all the diopside (± garnet at higher pressures) in the source is consumed, further melting of harzburgite produces an inflection of the melting trend away from Di towards the Ol–Hy join. Finally, the trend inflects sharply again, when all source enstatite is consumed, and becomes radial to Ol. For illustration, this trend is drawn through the 3 GPa data in Fig. 13a (Hirose & Kushiro, 1993; Walter, 1998). The contrasting effects of added CO2 or H2O on mantle melt compositions are also shown in Fig. 13a. Glasses formed by partial fusion of carbonated lherzolite KLB-1 at 3 GPa (Hirose, 1997) and hydrous KLB-1 at 1 GPa (Hirose & Kawamoto, 1995) are plotted in the diagram.

The field occupied by all the PVF samples is shown in Fig. 13a, together with the individual points for all >10·7 wt % MgO lavas. Some of the latter plot along the fertile mantle experimental partial melting trend at 1·0 GPa; others are slightly displaced towards the curved line drawn through published experimental data for the equilibria at 1 atm between olivine, plagioclase, Ca-rich clinopyroxene and natural basaltic (s.l.) liquids (Thompson et al., 2001).

Figure 2 showed that the olivines in all PVF lavas (up to Fo88) are appreciably less forsteritic than would be in equilibrium with fertile lherzolite mantle (Fo89; Pearson & Nowell, 2002; Wang et al., 2002). Therefore, a parental magma for the PVF was derived by first calculating the average of all analyses with MgO > 10·7 wt % (for reasons explained above) and then adding olivine in increments of 1 wt %. At each step, the equilibrium olivine was recalculated using a value of KdFe–Mg between olivine and liquid of 0·30 (Ulmer, 1989), thus employing the same approach as Larsen & Pedersen (2000), Wang et al. (2002) and others. The calculated parental melt (Table 2) has 13·4 wt % MgO and plots amongst the data points for anhydrous small-fraction (∼5%) experimental melts of ‘fertile’ lherzolite mantle at 1–1·5 GPa pressure in Fig. 13a. Obviously, this pressure is predicated on the choice of Fo89 for primary olivine composition and therefore somewhat arbitrary. The interesting feature of Fig. 13a is the clear way it shows that the parental/primary magmas of the Potrillo alkalic lavas have normative compositions that fit well with those of anhydrous small-fraction fertile lherzolite melts.

At first sight, this is unexpected because it is generally assumed that volatiles, both H2O and CO2, play a significant role in basanite genesis (e.g. Green et al., 2001). Therefore, this point is investigated further in Fig. 11b, by plotting the analyses of dredged glasses from Loihi Seamount and from the North Arch Volcanic Field, 200–300 km north of Oahu, Hawaii (Dixon et al., 1997; Dixon & Clague, 2001). These submarine glasses have escaped much of the near-surface degassing that occurs during subaerial eruption and have therefore preserved most of their pre-eruption gas contents (Fig. 13b, inset). In particular, the North Arch basanite glasses have the highest published H2O and CO2 contents of all well-studied basanites. Dixon et al. (1997) used a similar incremental olivine addition to ours to calculate the likely North Arch parental melt compositions. Again, they plot in Fig. 13b amongst the anhydrous experimental small-fraction fertile lherzolite melts at pressures up to ∼3 GPa. Dixon et al. (1997) estimated that the North Arch parental magmas were produced by 1·6–9% batch melting of a lherzolite mantle—a figure consistent with the array of points in Fig. 13b. To explain why anhydrous isobaric phase relationships provide a satisfactory simple model for basanite genesis, it is necessary to move to more complex polybaric decompression melting models.

More detailed models

Major elements

Langmuir et al. (1992) used published experimental data to design a method whereby the primary Na2O and FeO contents of a magma, after adjustment for any post-genesis processes, could be used to infer the decompression melting path of its source mantle, provided that the major-element composition of the latter was assumed. Wang et al. (2002) have updated this approach (referred to as ‘LKP’ below) to incorporate more recent experimental data and applied it specifically to post-10 Ma basaltic (s.l.) magmatism in the southwestern USA. They discussed every aspect of this modelling approach in detail and we have followed their procedures with the PVF data, in order to allow direct comparison with their results.

Figure 14 shows Na2O and FeO variations in a range of decompression melting paths calculated by Wang et al. (2002) for a ‘fertile’ mantle broadly resembling KLB-1 (see Fig. 13). Also plotted are the calculated primary magmas (in equilibrium with Fo89·0) of three volcanic suites in the southwestern USA: Big Pine Volcanic Field (BP), Pinto Peak (PP) and Western Grand Canyon (WGC), representative of those studied by Wang et al. (2002). Each of these points was considered by Wang et al. to approximate the Na2O and FeO contents of the accumulated melt at the top of the local decompression melting column beneath each named locality. This depth, in turn, was considered by Wang et al. (2002) to be the base of the rigid lithospheric lid overlying the convecting mantle in each of the three areas. To take point BP (Big Pine Volcanic Field) as an example, this plots on the decompression–melting curve that originates (P0) at 2·0 GPa. The Big Pine primary magma plots on this line at 1·3 GPa (Pf) and this is taken in the LKP model as the base of the overlying rigid lithospheric lid. Similarly, this plot estimates the pressure at the base of the lithospheric beneath Pinto Peak (PP) to be ∼2·3 GPa, whereas beneath the Western Grand Canyon (WGC), the LKP model estimates the lithospheric thickness to be substantially more.

Fig. 14.

Decompression melting model for the estimated Potrillo primary magma (see text), using the model of Langmuir et al. (1992)—widely known as LKP, as modified by Wang et al. (2002). BP (Big Pine), PP (Pinto Peak) and WGC (Western Grand Canyon) refer to examples of Basin and Range calculated primary magmas given by Wang et al. (2002). Curved lines are polybaric fertile-mantle melting paths during adiabatic decompression. Each tick mark represents 0·1 GPa of decompression, after which fractional melts are accumulated to give primary FeO and Na2O contents. Arrow shows direction of melting, from pressure of intersection of the solidus (P0) to the surface (P = 0). Pf is the final pressure recorded by a magma and equates with thickness of lithospheric lid above the convecting mantle. See text for discussion of oxidation state of the Potrillo magmas.

Fig. 14.

Decompression melting model for the estimated Potrillo primary magma (see text), using the model of Langmuir et al. (1992)—widely known as LKP, as modified by Wang et al. (2002). BP (Big Pine), PP (Pinto Peak) and WGC (Western Grand Canyon) refer to examples of Basin and Range calculated primary magmas given by Wang et al. (2002). Curved lines are polybaric fertile-mantle melting paths during adiabatic decompression. Each tick mark represents 0·1 GPa of decompression, after which fractional melts are accumulated to give primary FeO and Na2O contents. Arrow shows direction of melting, from pressure of intersection of the solidus (P0) to the surface (P = 0). Pf is the final pressure recorded by a magma and equates with thickness of lithospheric lid above the convecting mantle. See text for discussion of oxidation state of the Potrillo magmas.

Three variants of the PVF average parental magma (Table 2) are plotted in Fig. 14, allocating differing amounts of the total Fe to Fe2O3 in each. Wang et al. (2002) used no Fe2O3 in their calculations and therefore our 0% Fe2O3 PVF point is the one that is comparable with their results. This lies on a decompression melting path originating at approximately P = 3·5 GPa (≡115 km depth) and terminating below the lithospheric lid at P = 3·1 GPa (≡100 km depth). The pressure–depth conversions from Wang et al. (2002, fig. 7) have been used, in order to maintain comparability with their results. Thompson & Gibson (2000) and Thompson et al. (2001) argued that the ratio of Fe3+ to total Fe (Fe3+/FeT) in tholeiitic basalts is probably close to the value for Hawaiian basalts, namely ∼0·1. We consider that this is a more suitable Fe3+/FeT value to use than zero and therefore prefer the PVF model in Fig. 14 with 10% Fe3+/FeT, close to the FMQ buffer. The decompression–melting trajectory appropriate to this composition originates at ∼2·8 GPa (≡95 km depth; interpolating between the trajectories at 2·0 and 3·0 GPa) and terminates beneath lithosphere at ∼2·2 GPa (≡70 km depth). This is the same as the lithospheric thickness beneath Kilbourne Hole derived from mineral equilibria studies of mantle xenoliths, summarized by Bussod & Williams (1991) and discussed further below. The total final extent of mantle fusion predicted by this model would be ∼7% (Wang et al., 2002). A mantle potential temperature (Tp) value of ∼1400°C can be derived by applying equation (1) of Hirschmann (2000) to the pressure at the base of the melting column implied by the LKP model above (Fig. 14).

Rare earth elements

An alternative way to model the genesis of the PVF magmas is to use the method of McKenzie & O'Nions (1991), by deriving partial melt distributions from REE concentrations. We have followed the approach of Smith (2001), using the INVMEL program of McKenzie & O'Nions (1991), in developing forward and inverse models for different scenarios and have varied the lithospheric thickness and Tp value so as to jointly best fit the Potrillo REE data. This is an independent approach from that of Langmuir et al. (1992) because it focuses on incompatible trace elements. In the case of inverting for the REE concentrations, the lid thickness must be pre-specified and, for this, the lithosphere thickness of 70 km derived from Kilbourne Hole lherzolite PT equilibration data (Bussod & Williams, 1991) is used. The potential temperature (Tp) of the underlying convecting mantle is inferred by comparing the modelled melt distribution as a function of depth with those calculated using the parameterization of McKenzie & Bickle (1988). Two results that give satisfactory fits to the PVF data are presented below. In each of these cases, the spinel–garnet transition zone is set at 75–80 km, which corresponds to Tp∼1400°C (Klemme & O'Neill, 2000).

One-stage model. In this scenario, all the PVF melt is produced by decompression fusion of convecting mantle beneath a 70 km rigid lithospheric lid. The Nd-isotopic ratios of the PVF lavas, if they are unmodified by processes affecting the magmas since their genesis, can be used to estimate the abundances of REE, and thus other elements, in the uppermost convecting mantle beneath the area. The average value of εNd for all PVF lavas analysed to date is 6·4 (Table 4; Roden et al., 1988; Gibson et al., 1992). This implies that the convecting mantle beneath the PVF is definitely not geochemically the same as the source of MORB, but has a composition between the Depleted and Primitive mantle sources used by McKenzie & O'Nions (1995). Trace-element abundances in the model source mantle are adjusted accordingly.

The input PVF data is the average of the 14 Main Series samples with MgO > 10·7 wt % (Table 2). The model results (Fig. 15b–d) compare calculated model element abundances for decompression melts of the convecting mantle (continuous lines) with the mean sample element abundances of the PVF samples (filled circles), normalized to the Depleted Mantle source of McKenzie & O'Nions (1991). The best-fit model is very good for REE (Fig. 15b) and good for most of the other trace elements considered, except for Pb, Nb and Ta (Fig. 15c and d). The misfit for Pb is possibly a result of uncertainty in the partition coefficients used for this element, but this is not a likely explanation for the Nb and Ta misfits. These will be discussed in more detail below, in the context of a three-stage model.

Fig. 15.

Best-fit, REE-constrained, one-stage inversion model for the genesis of the PVF Main Series magmas, using the INVMEL program of McKenzie & O'Nions (1991). Lithospheric thickness—input to the inversion—is 70 km (Bussod & Williams, 1971). The magma source is assumed to be convecting asthenospheric mantle upwelling beneath this rigid lid. The convecting mantle composition is calculated by INVMEL from an input εNd value of +6·4, which is the mean PVF value (Table 4; Roden et al., 1988; Gibson et al., 1992), and falls between the Depleted and Primitive Mantle sources of McKenzie & O'Nions (1995). (a) The dashed line is the melt distribution with depth calculated for the asthenospheric melt. The dotted lines are theoretical anhydrous melt distributions, calculated using the parameterization of McKenzie & Bickle (1988) with ΔS = 350 J kg−1K−1 (Kojitani & Akaogi, 1997), for potential temperatures of 1400 and 1450°C. (b)–(d) Filled circles are mean sample elemental abundances, normalized to the Depleted Mantle source of McKenzie & O'Nions (1991). Error bars are the standard deviations of samples, plus the estimated error in source compositions. Continuous lines are calculated model abundances for asthenospheric melts.

Fig. 15.

Best-fit, REE-constrained, one-stage inversion model for the genesis of the PVF Main Series magmas, using the INVMEL program of McKenzie & O'Nions (1991). Lithospheric thickness—input to the inversion—is 70 km (Bussod & Williams, 1971). The magma source is assumed to be convecting asthenospheric mantle upwelling beneath this rigid lid. The convecting mantle composition is calculated by INVMEL from an input εNd value of +6·4, which is the mean PVF value (Table 4; Roden et al., 1988; Gibson et al., 1992), and falls between the Depleted and Primitive Mantle sources of McKenzie & O'Nions (1995). (a) The dashed line is the melt distribution with depth calculated for the asthenospheric melt. The dotted lines are theoretical anhydrous melt distributions, calculated using the parameterization of McKenzie & Bickle (1988) with ΔS = 350 J kg−1K−1 (Kojitani & Akaogi, 1997), for potential temperatures of 1400 and 1450°C. (b)–(d) Filled circles are mean sample elemental abundances, normalized to the Depleted Mantle source of McKenzie & O'Nions (1991). Error bars are the standard deviations of samples, plus the estimated error in source compositions. Continuous lines are calculated model abundances for asthenospheric melts.

Figure 15a shows the best-fit melt distribution with depth obtained by inversion of the REE concentrations. According to this model, melting clearly occurred in two spatially distinct phases within the decompressing mantle: (1) melting began at ∼90 km but the melt fraction produced during decompression remained very small (<1 vol. %) until ∼75 km; (2) most of the liquid was produced between 75 and 70 km. This melt developed with a depth distribution sub-parallel to and within 10°C of the theoretical anhydrous melt distribution for mantle with Tp = 1400°C (McKenzie & Bickle, 1988). Thus, the Tp value derived by this REE-based model is the same as was derived from the major-element-based LKP modelling above (Fig. 14). The small-melt-fraction, low-melt-productivity ‘tail’ in the decompression melting pattern shown in Fig. 15a is the way that the INVMEL program detects the ‘low-productivity, low-degree melting’ zone at depths below the main melting zone which Asimow & Langmuir (2003) and Asimow et al. (2004) modelled by other approaches in their analyses of the role of water in mantle decompression melting (see below). The total melt thickness is calculated to be ∼250 m. McKenzie & Bickle (1988) explained why it is better to express total amounts of melting as thicknesses, rather than percentages, when they are small.

Three-stage model. The first question to ask at this point is: if the one-stage REE-based model fits the observed element abundances in the PVF Main Series so well, why not use Occam's Razor and ignore more complex scenarios? There are two reasons why this approach is unsatisfactory.

(1) The Low-K Series lavas are an integral part of PVF magmatism. It would be perverse to suggest that their identical Sr–Nd isotopic ratios to the Main Series lavas (Fig. 7a) are no more than a coincidence. Therefore, the postulated mantle source of the PVF lavas must be capable of generating parental melts for both series. Although it was explained above that post-genesis processes, such as fractional crystallization and magma mixing, had modified the original element abundances and ratios of the Low-K lavas, nevertheless, their distinctive K contents are strong evidence that their parental melts originated in equilibrium with residual mantle amphibole. Class & Goldstein (1997) summarized previous experimental studies to show that amphibole is unstable within convecting mantle with a potential temperature of ∼1400°C. Therefore, as these authors and others have deduced for magmatism elsewhere (e.g. Späth et al., 2001), it appears necessary to invoke a role for lithospheric mantle in the genesis of the PVF magmas.

(2) The one-stage model outlined above used, as input, a depth of 70 km to the base of the sub-PVF lithosphere, deriving this from the calculated PT equilibration conditions of Kilbourne Hole lherzolite xenoliths (Bussod & Williams, 1991). Figure 15a shows that this lithospheric thickness allows the top few kilometres of decompression melting in underlying mantle to take place within spinel-facies lherzolite. Numerous studies have been made of alkali-basalt-basanite magmas generated beneath thicker lithospheres, so that the underlying convecting mantle melts only within the garnet stability field. In all such cases, a one-stage decompression–melting model fails to reproduce the REE abundances of the lavas (e.g. Smith, 2001). These points are discussed in more detail below.

The three stages are: (1) veining and metasomatism of the lithosphere by hydrous melts from the underlying asthenosphere, as proposed by McKenzie & O'Nions (1995); (2) perturbation of the system, by lithospheric extension and/or arrival of hotter convecting mantle, promoting melting in the convecting mantle; (3) re-melting of metasomatized lithosphere and formation of a hybrid melt with contributions from both convecting and lithospheric mantle sources.

In the PVF example, εNd is again taken as 6·4 (for reasons explained above) for both the convecting and initial (pre-veining) lithospheric mantles. Such a situation would arise in cases where the lowermost lithosphere was being formed by progressive freezing of convecting mantle onto its base. The veining is modelled as 10% addition of a metasomatizing agent, which was itself formed by 0·4% melting of underlying convecting depleted garnet lherzolite in the presence of volatiles, but could potentially be derived from another garnet-bearing mafic lithology [e.g. the MIX1G pyroxenite of Hirschmann et al. (2003)] forming veins, streaks or blobs within the convecting mantle. Figure 16b shows that the best-fit result models REE in the PVF magmas excellently; the light REE fit is slightly better than in the one-stage case (Fig. 15b). The melt distribution with depth ranges from 80 to 70 km (Fig. 16a) and is again close to (∼10°C above) and sub-parallel to the theoretical isentropic anhydrous melting curve of McKenzie & Bickle (1988) for Tp = 1400°C. The final melt fraction is ∼4 vol. % and this is sourced 80% from asthenosphere and 20% from veined and metasomatized lithosphere. The total melt thickness is ∼350 m.

Fig. 16.

Best-fit, REE-constrained, three-stage forward melt model for the genesis of the PVF magmas. See text for details, including the reasons why this scenario is necessary to explain the genesis of the Low-K Series PVF magmas. (a)–(d) As for Fig. 15, except that: dashed lines are calculated normalized elemental abundances for asthenospheric melts; dotted lines are the same for lithospheric melts; continuous lines are the same for a mixture of 80% asthenospheric and 20% lithospheric melts. The asthenospheric source is as in Fig. 15; the lithospheric source is the same, but enriched by the addition of 10% metasomatic melt. The modelled metasomatic melt is formed by melting the asthenospheric source by up to 0·4 wt % in the garnet-lherzolite zone (see text).

Fig. 16.

Best-fit, REE-constrained, three-stage forward melt model for the genesis of the PVF magmas. See text for details, including the reasons why this scenario is necessary to explain the genesis of the Low-K Series PVF magmas. (a)–(d) As for Fig. 15, except that: dashed lines are calculated normalized elemental abundances for asthenospheric melts; dotted lines are the same for lithospheric melts; continuous lines are the same for a mixture of 80% asthenospheric and 20% lithospheric melts. The asthenospheric source is as in Fig. 15; the lithospheric source is the same, but enriched by the addition of 10% metasomatic melt. The modelled metasomatic melt is formed by melting the asthenospheric source by up to 0·4 wt % in the garnet-lherzolite zone (see text).

As with the one-stage modelling, the Nb and Ta contents of the PVF parental magmas are substantially higher than those predicted by a three-stage INVMEL model (Fig. 14c). This hints that the lithospheric veins were not just the closed-system crystallization products of invading small-fraction melts. Instead, they appear to have been relatively enriched in a phase containing abundant Nb and Ta, such as ilmenite or rutile. Further discussion of such an idea encounters a circular argument.

Although this model invokes trace-element enrichment of a peridotite mantle source by means of subsequent pyroxenite veining, it does not follow Hirschmann et al. (2003) and Kogiso et al. (2003) in suggesting that the magmas of the PVF might simply be decompression melts of such Mg-rich pyroxenite veins within convecting mantle, without involving their enclosing peridotite. Kogiso et al. (2004) have argued that correlations between 187Os/188Os and other radiogenic isotopic ratios in various ocean-island alkalic lava suites is evidence that their Mg-rich pyroxenite source veins, streaks and blobs melted to give rise to magmas that migrated upwards without any significant interaction with enclosing peridotite, in contrast to proposals by Foley (1992) and Yaxley & Green (1998). A substantial number of such pyroxenite source bodies could be expected to occur within the convecting mantle underlying a region of PVF size and it would be inevitable that differences in Rb/Sr and Sm/Nd amongst them would eventually lead to a range of 87Sr/86Sr and 143Nd/144Nd in such pyroxenites. Thus, whilst some involvement of recycled pyroxenite bodies cannot be ruled out, the Sr–Nd isotopic homogeneity of the Mg-rich PVF lavas (Fig. 7a) is evidence against the idea that their sources were entirely pyroxenitic, with efficient melt segregation.

DISCUSSION OF PETROGENETIC PROCESSES

Convecting mantle source: hot, wet or both?

All the models developed above required decompression melting of convecting mantle between 90 and 70 km depth. Figure 14 and the data of Hirschmann (2000) show that such melting necessitates a mantle Tp of ∼1400°C. Within the overall anhydrous model of McKenzie & Bickle (1988), this potential temperature is between the global ambient convecting mantle Tp (∼1300°C) and their estimate of a typical plume value (∼1550°C). The next question to arise is whether the volatile content of the PVF lavas might indicate sufficient H2O in their mantle source to lower the Tp value calculated above significantly. Green et al. (2001) have compared MORB and Hawaiian tholeiitic picrites and argued that the water contents of the latter have such a large effect on their liquidus temperatures that it is unnecessary to postulate that the sub-Hawaiian convecting mantle has higher Tp than that beneath mid-ocean ridges elsewhere.

Concentrating first on water, in the absence of direct determinations on PVF lavas, it is necessary to make an indirect estimate of their H2O contents. The Hawaiian North Arch basanites (Fig. 13b) are the most hydrous, well characterized, geochemically comparable suite to the PVF and their water contents equate to an estimated source-mantle H2O concentration of 525 ± 75 ppm (Dixon et al., 1997). This value may either apply approximately to the PVF source mantle, if the magmas are simple melts of convecting mantle (Fig. 15), or must be reduced somewhat for the PVF convecting mantle partial source, if our three-stage REE-based model (Fig. 16) is the preferred one. This is because the latter postulates a stage three boosting of H2O in the melts by fusion of hydrous, veined, metasomatized lower lithosphere. Therefore, a realistic minimum H2O concentration in the convecting mantle beneath the Potrillo area would be ∼450 ppm—the value calculated for the Hawaiian mantle plume by Wallace (1998) and about three times the estimated MORB-source mantle value (Michael, 1988).

Wallace (1998) argued that even if sufficient fluorine is present to stabilize fluor–phlogopite and/or fluor–K-richterite, the amounts of these phases that could be present in KLB-1-type mantle would be insufficient to accommodate the water. Most of it would be contained as OH in nominally anhydrous silicates (Hirth & Kohlstedt, 1996). Wallace (1998, fig. 2) showed that Hawaiian plume mantle probably began to generate small-fraction hydrous melts at ∼250 km depth and that this process had probably consumed virtually all the H2O in the source mantle before anhydrous decompression melting commenced at ∼100 km depth. Thus, the water in the Hawaiian plume source contributes a small amount of deep-seated small-fraction hydrous melt to the upward migrating silicate liquids but does not affect the essentially anhydrous decompression melting zone, where most of the magma is produced. Figure 15a shows how the REE-based one-stage INVMEL modelling procedure generates a melt distribution with depth for the PVF magmas that is also most easily explained by the same scenario. It also closely resembles the melt productivity as a function of depth predicted by Asimow et al. (2004, fig. 2), using the pHMELTS algorithm that they have developed to model decompression melting of hydrous mantle beneath the Azores.

Similar reasoning can be applied to the major-element model shown in Fig. 14. Wang et al. (2002), Asimow & Langmuir (2003) and Asimow et al. (2004) have discussed this aspect in detail. Likewise, consideration of the possible CO2 content of PVF primary magmas follows similar reasoning to the same conclusion—by <100 km depth in upwelling convecting mantle, both H2O and CO2 will be almost entirely contained in a small-fraction, volatile-rich, high-pressure, low-viscosity melt that will mingle with anhydrous melts as they form but play relatively little part in the anhydrous melting process. The key factors are the low solubility of water in nominally anhydrous mantle silicates at ∼100 km depth (Hirth & Kohlstedt, 1996; Wallace, 1998) and the removal of K from the sub-solidus upwelling mantle by partitioning into small-fraction melts formed during deeper breakdown and fusion of hydrous silicates, if present (Wallace, 1998, fig. 1). There is, therefore, no reason to lower the estimate of the Tp of the convecting mantle beneath the Potrillo area significantly from a value of ∼1400°C. This view contrasts with that of Green et al. (2001), which was based on the behaviour of water in isobaric melting experiments, rather than considering decompression melting.

Comparison with the geophysics and magmatism of Hawaii supports the viewpoint taken here. Lithospheric thickness beneath Hawaii (∼70 km; Watson & McKenzie, 1991) is similar to that beneath the Potrillo area. When the Hawaiian mantle plume (Tp ∼ 1550°C) is directly beneath this lithosphere, its melt products are tholeiitic. Basanites only occur in Hawaiian magmatism at the space–time margins of the plume upwelling, such as North Arch and the Honolulu Volcanic Series (Dixon & Clague, 2001; Yang et al., 2003). This is consistent with the view that basanite genesis, especially when associated with alkali–olivine basalts, requires convecting mantle potential temperatures hotter than ambient but less than those of mantle plumes.

Before discussing possible sources of convecting sub-lithospheric mantle with Tp ∼ 1400°C beneath the PVF, it is first necessary to ask again whether this is ‘hotter than normal’, as McKenzie & Bickle (1988), McKenzie (1989) and McKenzie & O'Nions (1991, 1995) proposed, or within the ‘normal’ wide Tp range beneath mid-ocean ridges, as Anderson (2000), Green et al. (2001) and others argued. We remain convinced that the key factor in this debate is the remarkable constancy of MOR crustal thickness worldwide (White et al., 1992) and that alternative viewpoints tend to ignore this and its implications. Nevertheless, it should be noted that Wang et al. (2002) assembled arguments suggesting that ∼1300°C may be too low for the ambient Tp beneath most mid-ocean ridges. Taking the view that Tp ∼ 1400°C is ‘hotter than normal’, the next problem is to decide where this mantle originated. Wang et al. (2002) discussed this point in detail, following their modelling of high Tp values beneath the central Basin and Range area. Rather than duplicating their reasoning, its conclusions can be summarized briefly as follows.

The most obvious potential source for such hot convecting mantle is the Yellowstone plume. Although Yellowstone is ∼1400 km north of the PVF, the two are linked by the semi-continuous crustal extensional feature of the Rio Grande rift and its northward prolongation (Eaton, 1987; Gibson et al., 1992). Hot buoyant mantle flowing outwards from the Yellowstone plume would, if it encountered a channel formed by rifting (plate thinning) beneath the overlying lithospheric lid, flow preferentially along this route (Thompson et al., 1989; Leat et al., 1991; Thompson & Gibson, 1991). Humphreys et al. (2000) have subsequently shown by means of teleseismic images of the upper mantle that the mantle thermal structure beneath Yellowstone and the northwest USA is complex. They considered that a specific buoyant plume beneath Yellowstone is not essential to explain their observations and that convective rolls within the upper mantle may be a viable alternative. King & Anderson (1998) suggested that abrupt large-scale changes in lithospheric thickness, such as the boundary between the High Plains and thinner lithosphere to the west (Grand, 1987), might generate edge-driven convection, providing another local source of hot mantle beneath the PVF.

Timing and location of lithospheric veining and metasomatism

The concept that small-fraction melts leak continuously or discontinuously from the convecting mantle into the overlying lithospheric mantle has been firmly established for a generation (e.g. Frey & Green, 1974; Menzies & Hawkesworth, 1987; McKenzie, 1989; Yang et al., 2003). In many cases, the melt leakage may be dated as ‘geologically old’ relative to the later magmatism to which it is thought to contribute, because distinctive radiogenic isotopic ratios in the latter can best be modelled by a contributing source with distinctive elemental ratios that has remained outside the convecting mantle system for a long period of time. Nevertheless, there are also many examples of alkaline mafic magmatism, such as the PVF, where radiogenic isotope studies are consistent with their formation entirely as melts from convecting mantle of one sort or another (Roden et al., 1988; Ben Othman et al., 1990; Gibson et al., 1992). However, successful modelling of REE abundances and ratios, and supporting evidence from xenolith suites, frequently require a contribution from re-melting of veined metasomatized lithospheric mantle (McKenzie & O'Nions, 1995; Wilson et al., 1995; Smith, 2001; Späth et al., 2001; Yang et al., 2003; Johnson et al., 2005).

We have shown above that it is possible to model the genesis of the PVF Main Series magmas satisfactorily, from both major-element and REE viewpoints, by simple small-degree decompression melting at the top of the convecting mantle beneath the sub-Potrillo lithosphere. Clearly, this is one possible conclusion to our study, but it leads to two difficulties.

(1) The first problem with such an outcome is the ‘ordinariness’ of the PVF Main Series magmas; they are alkali basalts and basanites that show no distinguishing geochemical characteristics from worldwide occurrences of such rock types. Their macrocryst populations are unusual but this has not led to any unique geochemical features. Any petrogenetic model that is satisfactory for the PVF Main Series magmas should therefore also be potentially applicable to other continental and oceanic occurrences of similar rock types. The problem that then arises is the occurrence of such lavas at localities where the underlying lithosphere is significantly thicker than 70 km. For instance, beneath the alkali basalts and basanites of both the Chyulu Hills, Kenya (Späth et al., 2001) and the Vitim Volcanic Field, Baikal rift zone, Siberia (Johnson et al., 2005), the lithospheric mantle is estimated to be ∼100 km thick, using PT equilibria calculated from garnet lherzolite xenoliths erupted with the lavas (Novak et al., 1997; Litasov et al., 2000). If the INVMEL program of McKenzie & O'Nions (1991) is used to attempt to model the REE in these melts by peridotite decompression fusion alone beneath a ∼100 km lithospheric lid, it fails (Smith, 2001; Johnson et al., 2005) because the base of the lithosphere is deeper than the spinel–garnet transition in mantle peridotite at the appropriate temperature (Klemme & O'Neil, 2000). Instead, either a multi-stage peridotite enrichment and subsequent melting model, such as the one summarized in Fig. 16, is required or it is necessary to implicate a garnet-pyroxenite, rather than a peridotite, magma source (Hirschmann et al., 2003; Kogiso et al., 2003; Johnson et al., 2005).

(2) The second problem is that any hypothesis which successfully explains the genesis of the PVF magmas must include both the Main Series and the Low-K lavas. Their identical Sr–Nd isotopic ratios (Table 4) logically require a single mantle source for both suites and, as explained above, genesis of the Low-K magmas by mantle fusion in the presence of residual amphibole is the most plausible explanation for their distinctive elemental features. Class & Goldstein (1997) summarized previous experimental studies to argue that amphibole is not stable anywhere within convecting mantle with Tp = 1300°C, let alone hotter. They concluded that amphibole can only be stable in mantle when it forms part of the lithosphere. No definite evidence about the PT stability of amphibole in peridotite has subsequently emerged to alter the Class & Goldstein viewpoint (S. F. Foley, personal communication, 2003).

In the southern Rio Grande rift at ∼32°N, there is now enough evidence, either already published or reported in this study, to begin to tightly constrain both the time-scale and location of a ‘geologically young’ lithospheric veining and metasomatism event. The Neogene tectonomagmatic evolution of the rift has taken place in two distinct phases (e.g. Keller et al., 1990). Between ∼30 and ∼20 Ma, rapid extension was accompanied by widespread magmatism (including the Potrillo area) that had calcalkaline characteristics and Sr–Nd–Pb isotopic ratios extending outside the MORB–OIB range. Although some of these features are probably related to patchy contamination of the convecting mantle by material subducted beneath the western USA (Leat et al., 1988), there is also no doubt that ‘geologically old’ veined and metasomatized lithospheric mantle also contributed (Gibson et al., 1992, 1993; Thompson & Gibson, 1994). After a 10 Myr lull, both extension and magmatism resumed, and the young magmatism shows no sign of this ‘geologically old’ component.

The intensity of the first magmatic episode in the Potrillo area is confirmed by detailed studies of the Kilbourne Hole lower crust and lithospheric mantle xenoliths. Bussod & Williams (1991) showed that the PT values calculated from mineral equilibria in the Kilbourne Hole xenoliths recorded a thermal climax sufficiently intense to cause lower crustal fusion during the 30–20 Ma magmatism. Scherer et al. (1997) used Hf isotopic ratios to date this event at 25 Ma, affecting crust originally formed at or before 1·6 Ga (Padovani & Reid, 1989). The logical interpretation of the observations listed above is that the intense first tectonomagmatic phase purged the lithosphere beneath the Potrillo area of essentially all its pre-existing ‘geologically old’, relatively fusible parts, hence the rarity of mica and amphibole in the Kilbourne Hole mantle xenolith suite (Padovani & Reid, 1989).

Clearly, two other possibilities can neither be explored nor eliminated: (1) that future sampling at Kilbourne Hole or elsewhere in the PVF will produce examples of hydrous veined mantle with OIB-like isotopic characteristics; (2) that such veins formerly existed in the mantle sampled by the Kilbourne Hole xenolith suite but that all traces of these veins and any metasomatized mantle enclosing them have been eliminated by selective fusion during the PVF magmatism. If these hypothetical undiscovered MBL veins were amphibole-bearing, they would also have to be geologically young, in order to prevent the development of radiogenic 87Sr/86Sr and 143Nd/144Nd ratios. The only logical alternative to explore is to accept and reconcile the conflicting evidence from the PVF lavas and the Kilbourne Hole mantle xenoliths.

If a veined mantle source for part of the PVF melts developed during the last 20 Myr or so and was not sited within the Mechanical Boundary Layer (MBL) lithosphere sampled at Kilbourne Hole, it must—by elimination—have developed within the Thermal Boundary Layer (TBL; sensuMcKenzie & Bickle, 1988) that re-grew beneath the MBL lithosphere since the Miocene extension and associated magmatism. This episode of TBL formation cannot be quantified because the potential temperature of the underlying convecting mantle at that time is unknown. Nevertheless, the Kilbourne Hole xenolith PT data of Bussod & Williams (1991) can be used to fix the temperature at the base of the local ∼70 km thick MBL lithosphere at 1050°C. The convecting mantle freezing beneath this ancient lithospheric lid during TBL growth would have a temperature gradient rising from 1050°C at its top to the temperature of the asthenosphere below.

A temperature of 1050°C at ∼70 km depth is well below the solidus of peridotite containing mica and amphibole, and slightly below the solidus of an olivine-poor vein dominated by clinopyroxene, Ca-amphibole and mica (Foley et al., 1999; S. F. Foley, personal communication, 2003). Therefore, the top few kilometres of such a lithospheric TBL would be an ideal site for incipient hydrous melts, rising from the asthenosphere below, to freeze and form abundant hydrous veins within the newly stagnant TBL peridotite (McKenzie, 1989). Such veins would react with the surrounding peridotite wall rocks as they solidified (e.g. Litasov et al., 2000). Subsequent reheating of the TBL mantle would take place if Tp rose in the convecting mantle below and further complex vein and wall-rock reactions would then take place during the fusion of the veined mantle zone (Yaxley & Green, 1998). Stage 3 of the multistage INVMEL model (Fig. 16) is a simplified representation of this process. Its advantage over other PVF genetic scenarios is that it might also be able to clarify the genesis of similar alkalic magmas elsewhere, such as the Vitim Volcanic Field (Litasov et al., 2000; Pearson et al., 2004; Johnson et al., 2005), the Chyulu Hills (Novak et al., 1997; Späth et al., 2001) and Hawaiian North Arch (Yang et al., 2003). In each of these examples, the precise balance between peridotite and pyroxenite magma sources is a topic of current debate (Hirschmann et al., 2003; Kogiso et al., 2003, 2004; McKenzie et al., 2004; Johnson et al., 2005).

Wilson et al. (1995) also suggested a very similar hypothesis to explain the genesis of Tertiary–Quaternary melilitites in western and central Europe. The major difference between their proposal and the one outlined here concerns the thickness of the TBL at the base of the lithosphere. Following Anderson (1994), they took the viscous–elastic boundary isotherm (650 ± 100°C) as the upper boundary of the TBL, and argued that lithospheric mantle hotter than this would be susceptible to convective overturn on a relatively short geological time-scale, and therefore unable to develop radiogenic isotopic characteristics significantly different from those of the underlying convecting mantle. Obviously, this viewpoint is not applicable in cratonic areas, where subcontinental lithospheric mantle is inherently buoyant because of its composition (e.g. Pearson & Nowell, 2002; Boyd et al., 2004). The evidence of the Kilbourne Hole mantle xenoliths, as summarized earlier in this paper, is that stable MBL lithospheric mantle there also extends much deeper than the 650°C isotherm, even in a region where it initially stabilized only 1·6 Gyr ago (Meisel et al., 2001), and has a relatively fertile composition throughout (e.g. the Kilbourne Hole lherzolite xenolith KLB-1, which has been extensively used in phase-equilibria studies; see Fig. 13).

The major improvement in understanding lithospheric mantle structure and evolution that has taken place since the study of Wilson et al. (1995) is the use of Os isotopic data to ‘map’ the model ages of initial lithospheric stabilization through its associated Re depletion, relative to Os. In places where appropriate data are available, such as the Archaean cratonic and surrounding Proterozoic platform areas of southern Africa (Pearson et al., 2002; Carlson et al., 2005), Kilbourne Hole (Meisel et al., 2001) and Vitim, Siberia (Pearson et al., 2004), the age of local lithospheric mantle stabilization (Re depletion) is close to the oldest radiometric dates for crustal rocks in that area. Furthermore, such ancient model Re-depletion ages characteristically extend to samples that originated far deeper than the 650°C isotherm: to >200 km beneath the Kaapvaal Craton; 150 km beneath southern Namibia; 70 km beneath Kilbourne Hole; 100 km beneath Vitim (see references above). Thus, the concept of the role of TBL mantle in alkalic magma genesis that Wilson et al. (1995) advocated can be focused in the case of PVF magmatism to give a detailed magma genesis model, summarized in Fig. 17.

Fig. 17.

Summary of the results and interpretations in this study. The left-hand panel summarizes both new results and relevant published information (Bussod & Williams, 1991; Huckenholz et al., 1992; Meisel et al., 2001). McKenzie & Bickle (1988) explained why it is preferable to express the amounts of melt produced during mantle decompression as thicknesses, rather than volume percent. The right-hand panel is a schematic illustration showing our preferred interpretation of all the data. See text for details. The stippled round and oval ‘magma chambers’ in the right-hand panel represent magma cumulate and reservoir zones beneath the PVF, emphasizing their vertical range and concentration immediately beneath the Moho. Thompson et al. (1972, p. 247) previously described such a magmatic plumbing model beneath the Palaeocene lava field of Skye, NW Scotland, as: ‘A ramifying plexus of subvolcanic conduits and fissures; resembling a sponge but without good lateral connections.’ This concept also seems appropriate for the PVF.

Fig. 17.

Summary of the results and interpretations in this study. The left-hand panel summarizes both new results and relevant published information (Bussod & Williams, 1991; Huckenholz et al., 1992; Meisel et al., 2001). McKenzie & Bickle (1988) explained why it is preferable to express the amounts of melt produced during mantle decompression as thicknesses, rather than volume percent. The right-hand panel is a schematic illustration showing our preferred interpretation of all the data. See text for details. The stippled round and oval ‘magma chambers’ in the right-hand panel represent magma cumulate and reservoir zones beneath the PVF, emphasizing their vertical range and concentration immediately beneath the Moho. Thompson et al. (1972, p. 247) previously described such a magmatic plumbing model beneath the Palaeocene lava field of Skye, NW Scotland, as: ‘A ramifying plexus of subvolcanic conduits and fissures; resembling a sponge but without good lateral connections.’ This concept also seems appropriate for the PVF.

The Os-isotopic data of Meisel et al. (2001) also contradict the hypothesis of Kil & Wendlandt (2004), based on mineralogical and textural studies of Kilbourne Hole peridotite xenoliths, that the base of the sub-PVF lithosphere is at only 48 km. As beneath Kaapvaal (Pearson et al., 2002; Carlson et al., 2005), changes in peridotite texture with mineral equilibration depth in the Kilbourne Hole xenoliths clearly do not mark the lithosphere–asthenosphere boundary.

SUMMARY AND CONCLUSIONS

The <80 ka alkali–olivine basalt–basanite lavas of the PVF, New Mexico, mostly have MgO contents >8 wt %, ranging up to 12·5 wt %. Conventional wisdom would assume that such Mg-rich lavas were very suitable material for a study of the mantle sources of continental extension-related mildly alkaline magmatism because they had probably undergone nothing but fractionation of some olivine since their genesis. The mineralogy and geochemistry of the lavas shows a more complex situation, involving extensive post-genesis modification of the magmas. For instance, fractional crystallization appears to have affected all these lavas and this process involved removal of Ca-rich clinopyroxene, in addition to olivine, from all compositions with MgO < 10·7 wt %. The clinopyroxene removal has significantly affected both the Na2O contents and REE pattern slopes (e.g. La/Lu) of magmas where it took place. Therefore, only PVF samples with MgO > 10·7 wt % can be used with confidence for petrogenetic modelling involving Na2O or REE.

There is also evidence that other processes took place to complicate the magmatic evolution. For instance, most of the lavas contain rounded and partially resorbed macrocryst fragments of olivine, clinopyroxene, spinel (pleonaste) and plagioclase that have compositions completely different from those of equilibrium phenocrysts in the same lava. These are evidence that magmas either resorbed earlier cumulates (also found as cognate xenoliths in PVF lavas) or mixed with phenocryst-rich remnants of earlier unerupted melt batches.

New Sr–Nd isotopic analyses of Mg-rich PVF lavas are tightly clustered around a mean of 87Sr/86Sr = 0·70308 (SD = 0·00004) and 143Nd/144Nd = 0·512952 (SD = 0·000025). Thus, in terms of Sr and REE, there is no evidence of interaction between these melts and the ∼1·6 Ga granite crust through which they were erupted, which has 87Sr/86Sr up to 0·79 and 143Nd/144Nd down to 0·5119 (Padovani & Reid, 1989). Nevertheless, Os isotopic analyses of four Mg-rich lavas show a subtly different picture. One of the samples contains relatively abundant Os (209 ppt; approximately four times the average crustal values; Saal et al., 1998), whereas the Os contents of the other three are extremely small (4·1–14·1 ppt), making them highly susceptible to crustal contamination. Their 187Os/188Os ratios increase with decreasing Os content, as would be predicted to result from crustal Os contamination (crustal 187Os/188Os = 0·15–2·5; Saal et al., 1998). The low Os contents of the Os-poor lavas probably resulted from early sulphide loss; small inclusions of sulphide occur in PVF clinopyroxene phenocrysts and megacrysts.

Whilst Os isotopic evidence marks the shallowest levels at which post-genesis processes affected the PVF magmas, the aluminous sub-calcic clinopyroxene megacrysts found at Kilbourne Hole and elsewhere in the PVF show how deep such processes began. The megacrysts formed in equilibrium with PVF magmas at depths of 50–60 km, within the subcontinental lithospheric mantle beneath the Potrillo area (Fig. 17). Kaersutitic amphibole megacrysts also precipitated at similar depths, from trapped melt batches where water contents built up locally.

There is a considerable amount of published information about the lithospheric mantle beneath the PVF, based on studies of mantle xenoliths found at Kilbourne Hole (e.g. Irving, 1980; Padovani & Reid, 1989; Bussod & Williams, 1991; Kil & Wendlandt, 2004). Its lithology is spinel lherzolite, with pyroxenite veins only in its shallower parts and both amphibole and phlogopite scarce. Pre-incorporation PT conditions of the xenoliths, calculated from lherzolite mineral equilibria, range from 850°C at 35 km to 1050°C at 67 km (Bussod & Williams, 1991). The equilibration temperatures for pyroxenite-veined xenoliths are in the range 850–1000°C (Bussod & Williams, 1991; Kil & Wendlandt, 2004). Extensive Os isotopic studies show that all this lithosphere was initially stabilized at ∼1·6 Ga and has undergone no subsequent event sufficient to disturb its Re–Os isotope systematics (Meisel et al., 2001). Therefore, the long-term-stable MBL lithosphere beneath the PVF is at least 70 km thick—a value within a few kilometres of the lithosphere base estimated by teleseismic travel-time delay tomography by Achauer & Masson (2002).

If the CIPW norms of Mg-rich PVF lavas are compared with the compositions of the melts produced during isobaric melting experiments on fertile mantle compositions, such as KLB-1, the PVF compositions are appropriate for small-degree melts (a few percent) of anhydrous lherzolite mantle fusion at pressures in the range 1·5–3·0 GPa.

An average parental PVF magma can be derived by taking the mean of chemical analyses with MgO > 10·7 wt % and then calculating the result of adding equilibrium olivine to it in small increments, until the equilibrium olivine is Fo89 (e.g. Larsen & Pedersen, 2000; Wang et al., 2002). This composition contains 13·4 wt % MgO. If 10% of the total Fe in this composition is allocated to Fe2O3, the decompression melting modelling approach of Langmuir et al. (1992) and Wang et al. (2002) suggests that it resulted from ∼7% melting in a decompressing upwelling mantle column with its base at ∼95 km, top at ∼70 km and a potential temperature (Tp) of 1400°C. Thus, the value of ∼70 km for the present thickness of the long-term-stable Proterozoic lithosphere beneath the PVF can be derived independently from LKP modelling and from lherzolite xenolith mineral PT equilibria. In turn, this can be used as an input constraint, enabling the INVMEL program of McKenzie & O'Nions (1991) to calculate a best-fit inverse model of mantle decompression melting to fit the REE and other trace-element abundances in PVF parental magmas. The simplest successful model is one-stage melting of convecting mantle with Tp ∼ 1400°C. The melt in this scenario is generated in two spatially consecutive stages: (1) <1 vol. % of hydrous melt produced between 90 and 75 km depth; (2) ∼2 vol. % of anhydrous melt produced between 75 and 70 km.

Despite the encouraging consistency between the results of two geochemical decompression–melting models and also mineral PT equilibria in the Kilbourne Hole mantle xenoliths, there is one crucial flaw. This is the occurrence of a variant within the PVF lavas (∼15% of the total analysed samples and randomly distributed amongst the ‘normal’ flows), with the same petrography as the majority but some geochemical differences, notably relatively low K2O contents. These low-K lavas have the same Sr–Nd isotopic ratios as the majority and the most plausible evidence for their elemental features is that they were melts from the same mantle source as all PVF magmas but with amphibole locally present in the residuum. This necessitates that the source mantle for all the PVF was lithospheric at the time of magma genesis, because convecting mantle with Tp ∼ 1400°C is too hot below 70 km depth for amphibole to be stable (e.g. Class & Goldstein, 1997). Nevertheless, the value of 187Os/188Os in the single Os-rich lava located by this study is 0·145 (γOs = +13·6)—a value well above the measured range in Kilbourne Hole lherzolite xenoliths (187Os/188Os 0·117–0·130; Meisel et al., 2001) and typical of OIB in oceanic magmatism derived from mantle plumes. These data, in turn, rule out peridotitic lithospheric mantle from above ∼70 km as the PVF magma source.

Therefore, the probable source of all of the PVF magmas is:

  1. located at the base of the lithosphere (below ∼70 km) and was re-melted during regional extension and re-heating in the last few million years;

  2. lithospheric (not convecting) during the immediately preceding period and, therefore, able to cool significantly by conduction to the surface to stabilize amphibole, at least locally; this mantle was probably a product of veining by small-fraction melt invasion from below;

  3. isotopically (Sr–Nd–Os) similar to the convecting mantle sources of OIB and not the Proterozoic MBL mantle sampled by xenoliths at Kilbourne Hole.

The source is thus identified as a thin TBL (sensuMcKenzie & Bickle, 1988) of convecting mantle that froze onto the base of the sub-PVF lithosphere immediately after early Basin and Range intense extension and magmatism ceased at ∼20 Ma. A three-stage calculation using INVMEL can model this complicated succession of lithospheric TBL formation below 70 km, veining and subsequent re-fusion satisfactorily. Figure 17 attempts to illustrate all these points.

SUPPLEMENTARY DATA

Supplementary data for this paper are available on Journal of Petrology online.

The fieldwork and INAA were funded by NERC Research Grant GR3/5299. Agents of the US Border Patrol and Potrillo ranchers helped greatly with advice on access and other issues. The microprobe analyses were made at Manchester University, using the NERC/MEMF instruments. Mike Henderson enabled this and Dave Plant gave excellent technical support. The majority of XRF analyses were made at Birmingham University, with assistance from Graham Hendry. Our particular thanks go to Nick Marsh for undertaking some final XRF major-element analyses at Leicester University. Paul Asimow, Steve Foley, Vicky Hards, Dimitri Ionov, Dan McKenzie, Glen Milne, Geoff Nowell, Terry Plank, Dave Sales and Bill Seager, together with Glen Izett, Peter Lipman, Dallas Peck and Ren Thompson (USGS), gave us all manner of invaluable advice and assistance. The perceptive comments of Terry Plank, Peter Reiners, Julia Shaw, Marjorie Wilson and an anonymous referee improved the manuscript immensely.

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