| ||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||
Journal of Petrology | Volume 43 | Number 6 | Pages 963-981 | 2002
© Oxford University Press 2002
Environments of Crystallization and Compositional Diversity of Mauna Loa Xenoliths
UNIVERSITY OF WASHINGTON, DEPARTMENT OF EARTH AND SPACE SCIENCES, BOX 351310, SEATTLE, WA 98195, USA
Received November 3, 2000; Revised typescript accepted December 4, 2001
 |
ABSTRACT |
|---|
 TOP ABSTRACT INTRODUCTIONRESULTSDISCUSSIONCONCLUSIONSAPPENDIX: PETROGRAPHYREFERENCES |
|---|
Two picrite flows from the SW rift zone of Mauna Loa contain xenoliths of dunite, harzburgite, lherzolite, plagioclase-bearing lherzolite and harzburgite, troctolite, gabbro, olivine gabbro, and gabbronorite. Textures and olivine compositions preclude a mantle source for the xenoliths, and rare earth element concentrations of xenoliths and clinopyroxene indicate that the xenolith source is not old oceanic crust, but rather a Hawaiian, tholeiitic-stage magma. Pyroxene compositions, phase assemblages and textural relationships in xenoliths indicate at least two different crystallization sequences. Calculations using the pMELTS algorithm show that the two sequences result from crystallization of primitive Mauna Loa magmas at 6 kbar and 2 kbar. Independent calculations of olivine Ni–Fo compositional variability in the plagioclase-bearing xenoliths over these crystallization sequences are consistent with observed olivine compositional variability. Two parents of similar bulk composition, but which vary in Ni content, are necessary to explain the olivine compositional variability in the dunite and plagioclase-free peridotitic xenoliths. Xenoliths probably crystallized in a small magma storage area beneath the rift zone, rather than the large sub-caldera magma reservoir. Primitive, picritic magmas are introduced to isolated rift zone storage areas during periods of high magma flux. Subsequent eruptions reoccupy these areas, and entrain and transport xenoliths to the surface.
KEY WORDS: xenolith; Hawaii; volcano plumbing; mineral composition; picrite
|
INTRODUCTION |
|---|
|
TOPABSTRACTINTRODUCTIONRESULTSDISCUSSIONCONCLUSIONSAPPENDIX: PETROGRAPHYREFERENCES |
|---|
Xenoliths hosted in Mauna Loa’s tholeiitic-stage lavas provide a rare opportunity to glimpse into the volcano’s magma chambers and plumbing system. In Hawaii, xenoliths are much more abundant in alkalic lavas than in tholeiitic lavas, as a result of reconfiguration of the conduit and storage system as magma supply rate decreases (Clague, 1987
). Each xenolith records the chemical and physical environment in a small part of the magma chamber at a particular time. Similarly, the individual crystals that compose xenoliths hosted in a single flow or series of flows may reflect the changes in magma composition over time. Textural and compositional features of xenoliths preserve evidence for variation in geochemical, petrological and physical conditions within the magma system.
Compositions of Mauna Loa xenoliths range widely and include pyroxenite, peridotite, troctolite and dunite, as well as gabbroic lithologies. Many exhibit distinctive cumulate or laminated textures. Previous studies of Mauna Loa xenoliths are few (White, 1966; Jackson et al., 1982), so the initial objective of this study is to characterize the petrography and geochemistry of a representative collection of samples. The second objective is to analyze textures, mineral assemblages, phase relationships and compositions to identify genetic relationships among the xenoliths as well as with the host, to construct a crystallization history of the magmas that are parental to the xenoliths, and ultimately, to model the formation and emplacement of the xenoliths in relation to variability in magma supply rate.
|
RESULTS |
|---|
|
TOPABSTRACTINTRODUCTIONRESULTSDISCUSSIONCONCLUSIONSAPPENDIX: PETROGRAPHYREFERENCES |
|---|
Field relationships
Two xenolith-bearing flows erupted from Mauna Loa’s SW rift zone are exposed on the SE coast of Hawaii (Fig. 1). The Waikapuna Bay site (henceforth referred to as ‘Site A’) has a 14C age of 7360 BP, and the second site, near Kahilipali Point (‘Site B’), has been dated at 4080 BP (F. Trusdell, unpublished data, 1999). The xenolith-bearing parts of the flows are <400 m in lateral extent and have variable thicknesses. At Site B, the flow has a planar morphology, which is incised in several places by wave-cut amphitheaters, thus exposing the entire flow thickness (
1·7 m). At Site A, the flow occurs as multiple lobes on a steep slope; consequently, there are no well-preserved flow cross-sections. Flow lobes and tubes are of the order of 0·5–1 m in diameter. The xenoliths are distributed irregularly throughout the host layer and range in size from 1 to 25 cm in the longest dimension; most are in the 1–3 cm size range.
|
Petrography
Ultramafic xenolith types include dunite, troctolite, harzburgite, lherzolite, plagioclase-bearing harzburgite, and plagioclase-bearing lherzolite. Some dunites contain interstitial plagioclase, and one sample has interstitial orthopyroxene. Most dunites are coarse grained (Fig. 2a), but one sample is fine grained, with a granular texture. One dunite has a granoblastic texture, unlike any texture observed in other xenoliths. In the troctolite, plagioclase sub-poikilitically encloses olivine (Fig. 2b). A few clinopyroxene crystals are also present.
|
In subsequent discussion, when no distinction between these two lithologies is required, the harzburgite and lherzolite are grouped as ‘peridotite’ or ‘plag-free peridotite’. In these xenoliths, large (15 mm) pyroxene crystals poikilitically enclosing anhedral, amoeboid or resorbed olivine are typical (Fig. 2c). Only one sample has ol + cpx (Fig. 2d); the rest contain ol + opx. Plagioclase-bearing harzburgite and lherzolite are grouped as ‘plag-bearing peridotite’. These xenoliths lack the poikilitic texture that is characteristic of the plag-free peridotites. Olivine and pyroxene in these xenoliths have variable and typically irregular shapes. Plagioclase is primarily interstitial, but some euhedral crystals are present (Fig. 2e).
Gabbroic xenolith types include gabbro, gabbronorite, olivine gabbro, and olivine gabbronorite. Textures are variable among the gabbroic xenoliths, but most can be classified in one of two textural families: those exhibiting ‘open textures’ defined by vesicles, interstitial glass and/or fine-grained basalt incorporated in the xenolith (Fig. 2f), and those that are holocrystalline (Fig. 2g). Only one sample, a gabbronorite, is fine grained (Fig. 2h). A laminated texture, defined by aligned plagioclase crystals, is present in some holocrystalline xenoliths (Fig. 2g). The holocrystalline xenoliths typically are much more cohesive and larger than the open-textured gabbroic or any of the ultramafic xenoliths. With one exception, the holocrystalline gabbroic xenoliths lack olivine. Every xenolith larger than 5 cm is gabbroic, holocrystalline and typically laminated. Gabbroic xenoliths are generically termed ‘gabbros’, and subdivided by texture (holocrystalline vs open). Detailed petrographic descriptions of all xenoliths and host samples are given in the Appendix.
Host lavas are highly variable in vesicle, glass and olivine content, although in general samples are typically highly vesicular and have abundant (up to 30%) olivine phenocrysts. Phenocrysts are generally euhedral, fractured and commonly show resorption along the edges. Olivine megacrysts (1–2·7 cm) are present in the host lavas and are fractured but typically very fresh. Crystals are dominantly euhedral but are embayed in places. Most megacrysts are kink-banded.
Analytical results
Analytical techniques
Whole-rock major and minor element compositions were analyzed using X-ray fluorescence (XRF), and trace element compositions using inductively coupled plasma mass spectrometry (ICP-MS), at the Washington State University GeoAnalytical Laboratory. Details of the XRF procedure have been described by Johnson et al. (1999). Precision for XRF analyses is based on repeat analyses of the BCR-P standard, and is within 0·4% for major elements, except MgO (1·4%) and Na2O (1·3%). ICP-MS analytical precision is calculated from repeat analyses of the same standard and is typically within 3%.
Secondary ion mass spectrometry (SIMS) analyses of clinopyroxene were completed on a Cameca IMS 4f ion microprobe at the University of New Mexico. Accelerating potential for analyses was 10 kV, ion beam current was 30–40 nA, and the focus diameter was 25–35 µm. For each analysis, repeated cycles of peak counting were performed on 30Si+, 88Sr+, 89Y+, 90Zr+, 139La+, 140Ce+, 146Nd+, 147Sm+, 151Eu+, 153Eu+, 163Dy+, 167Er+, and 174Yb+. Precision is typically better than 5–10% at chondritic concentrations. Data reduction utilized SiO2 content at the spot of analysis as determined in electron microprobe analyses [for detailed procedure, see Papike et al. (1995)].
Electron microprobe analyses were completed on a JEOL 733 Superprobe at the University of Washington, Department of Earth and Space Sciences. Reference standards are well-characterized natural minerals and glass. Analyses were carried out using a 15 kV accelerating potential, a 25 nA, fully focused beam and 40 s count times on all elements. Matrix corrections were made using the Armstrong (1988) correction routine. For pyroxene, analytical error is within 3% for SiO2, 3% for MgO, 7% for FeO, 7% for CaO, 8% for Al2O3, 6% for Cr2O3, 15% for TiO2, 50% for MnO, and 22% for Na2O, at concentrations typical of those observed in this study. Analytical error in olivine is within 2% for SiO2 and MgO, 3% for FeO, 4% for MnO, 8% for NiO and 20% for CaO, at concentrations typical of those observed in this study. In plagioclase, analytical error is within 1% for SiO2, 2% for Al2O3, 13% for CaO, 17% for K2O, 23% for Na2O, 53% for FeO, and 93% for MgO, at concentrations of 65·2%, 20·97%, 1·36%, 3·19%, 8·72%, 0·14% and 0·035%, respectively. All Fe is reported as FeO.
Whole- rock compositions
Major and minor elements. Five holocrystalline gabbro samples and eight host samples were analyzed (Table 1, Figs 3 and 4). Host lavas have 44–51 wt % SiO2 and 8–33 wt % MgO. The wide variation in SiO2 and MgO in the host samples is reflective of the highly variable fraction of olivine in the host. SiO2 in the gabbros ranges from 48 to 52% at 9–19 wt % MgO. The strong olivine control in the host lavas is evident in the linear pattern of variation in CaO, TiO2, Al2O3 and Na2O with respect to MgO content (Fig. 3). One gabbro sample plots on this linear array, but the rest of the gabbros show much more limited compositional variability and are distinct from the array defined by the host lavas.
|
|
|
The olivine accumulation trend defined by FeO–MgO variability encompasses all of the host samples, from both lava flows. In Fig. 4, steep lines labeled ‘Fo xx’ represent the loci of liquid compositions in equilibrium with olivine of Fo xx composition, as calculated using the exchange reaction: 0·3 = Kol–LFe–Mg = (FeO/MgO)ol/(FeO/MgO)liq (Roeder & Emslie, 1970). Compositions of all but one host lava lie on a mixing line between Fo88 olivine and a liquid with 12% MgO and in equilibrium with Fo88 olivine. The average phenocryst composition in the host lavas is Fo87. The one host lava sample that lies close to the Fo83 equilibrium line is an olivine-poor sample that has probably been affected by olivine fractionation. One gabbro sample falls on this olivine control line, but the bulk compositions of the rest of the gabbroic suite do not appear to be related by olivine accumulation. Although 3300 years separate the eruptions at the two sites, the inferred parental liquid Mg–Fe compositions for the host lavas from the two sites are effectively the same, and they are both controlled by a similar olivine-accumulation process. This indicates that consistent melt generation and transport processes operated over this time frame (Rhodes, 1983).
Trace elements. Chondrite-normalized rare earth element (REE) patterns in the host lavas are relatively uniform, with moderate light REE (LREE) enrichment relative to heavy REE (HREE) (La/Lu 30–33) (Fig. 5). The concentrations of La in the host lavas are 12–37 x chondrite values. None of the host lavas exhibits a Eu anomaly. Although the abundance of olivine in the host lavas contributes to the low REE concentrations in some samples, the shapes of the REE patterns are similar between sites A and B and are characteristic of Mauna Loa lavas (Tilling et al., 1987). All holocrystalline gabbro HREE patterns have a slope similar to the HREE slope of the host lavas. With one exception, sample 37, these samples have also have slightly negative LREE slopes, positive Eu anomalies, and concentrations lower than those observed in the hosts. These samples reflect pyroxene and plagioclase accumulation, and are not liquid compositions.
|
Rare earth element abundances determined for five spots in four different clinopyroxene grains from one gabbroic and three peridotitic xenoliths (78-2, 78f, 79d, 79a) are nearly identical (Table 2, Fig. 6), and all have a negative Eu anomaly. The liquid calculated to be in equilibrium with the clinopyroxene varies depending upon the chosen distribution coefficients. Liquids calculated using distribution coefficients from McKay et al. (1986) have slight LREE enrichment, a characteristic of Mauna Loa lavas (Tilling et al., 1987), but the REE slope in the calculated liquid is higher than characteristic of the lavas. Liquids calculated using the distribution coefficients of Hart & Dunn (1993), however, show a HREE slope characteristic of the lavas but have a slightly negative LREE slope, which is not observed in Hawaiian melts.
|
|
Regardless of choice of distribution coefficient, the REE concentrations in the liquids calculated to be in equilibrium with the clinopyroxene in the xenoliths are higher than the REE concentrations in the host lavas. Figure 6 shows the range of host lava REE concentrations after correction for 20–30% accumulated olivine and 20–35% olivine crystallization before pyroxene crystallization [as consistent with Kol–LFe–Mg calculations for olivine accumulation, olivine mode in the samples, and assumed liquid MgO content at the point of clinopyroxene saturation (Montierth et al., 1995)]. Clinopyroxene-equilibrium liquids calculated using the McKay et al. (1986) distribution coefficients fall within the range of LREE for the corrected host magmas, but the liquids calculated using the Hart & Dunn distribution coefficients produce the best fit for the HREE. Within the error introduced by choice of distribution coefficients, the clinopyroxene appears to have crystallized from liquids with similar REE concentrations and patterns to the liquid that crystallized as the host lava.
Mineral compositions
Mineral analyses were completed on xenolith, phenocryst and groundmass crystals in 14 samples. One to three crystal core and 1–3 crystal rim points were analyzed per crystal in the xenolith as well as for phenocrysts in the host lava. One crystal core point was analyzed per grain in the groundmass. Reported analyses represent single spots. Approximately 10–15 points for each type of mineral were analyzed per thin section. Olivine, pyroxene and plagioclase compositions are presented in 7–10 and the full dataset may be downloaded from the Journal of Petrology Web site at http://www.petrology.oupjournals.org.
|
|
|
|
Olivine Ni–Fo variability within any single dunite or peridotitic xenolith is much more restricted than the range of compositions represented by a given xenolith type. This contrasts with the wide range of olivine compositions present in the open-textured gabbros; no olivine was analyzed in the holocrystalline gabbros. Pyroxene Cr–mg-number variability shows this same pattern. With the exception of holocrystalline gabbronorites 38 and 70 and holocrystalline gabbro 65, the plag-bearing xenoliths have much wider ranges in pyroxene composition than the plag-free xenoliths. Pyroxene compositions within a single peridotitic or gabbroic xenolith are much more restricted than the range observed across a whole family, with the exception of the open-textured xenoliths. Plagioclase An content in the xenoliths of a single rock type is highly variable, and with the exception of the troctolite, single xenoliths commonly show a significant range in plagioclase compositions.
Normal zoning is observed in olivine in the host lavas, whereas olivine in open-textured gabbroic xenoliths exhibits both normal and reverse zoning. Clinopyroxene in gabbroic xenoliths also shows reverse or normal zoning, and orthopyroxene in these xenoliths is either reversely zoned or unzoned. Pyroxene in the plagioclase-bearing peridotites is normally zoned or unzoned. Zoning is not observed in olivine or pyroxene in the plagioclase-free peridotite, dunite or troctolite xenoliths, with the exception of the small amounts of orthopyroxene in dunite, which is normally zoned.
|
DISCUSSION |
|---|
|
TOPABSTRACTINTRODUCTIONRESULTSDISCUSSIONCONCLUSIONSAPPENDIX: PETROGRAPHYREFERENCES |
|---|
Source of xenoliths
Possible sources for Hawaiian xenoliths are mantle lithosphere, old oceanic crust, or crystalline remnants of a Hawaiian magma (e.g. Clague & Chen, 1986
; Sen & Leeman, 1991). Olivine in the mantle lithosphere is typically more magnesian (Fo90–92, Clague et al., 1995) than even the most primitive olivine compositions observed in Mauna Loa xenoliths (Fo87–88). Textures of lithospheric mantle xenoliths from Hawaii are porphyroclastic and coarse granoblastic (Sen & Leeman, 1991), textures observed in only one dunite xenolith in this study. Normal mid-ocean ridge basalt (MORB), such as that which composes the Pacific plate underlying Hawaii, has characteristic LREE depletion (Sun & McDonough, 1989), which contrasts strongly with LREE enrichment observed in the gabbroic xenoliths and reconstructed liquid compositions (Figs 5 and 6). On the basis of these comparisons, both mantle lithosphere and oceanic crust can be eliminated as a potential xenolith source. The remaining potential source is crystalline remnants of Hawaiian magma.
Xenoliths hosted in this tholeiitic-stage magma could be products of crystallization of an earlier tholeiitic-stage magma, remnants from the pre-shield alkalic stage of magmatism (Bohrson & Clague, 1988; Chen et al., 1992; Fodor & Galar, 1997), or rocks from an older, underlying volcano such as Hualalai or Mauna Kea. In general, tholeiitic-stage lavas have flatter REE patterns and lower REE concentrations than those from alkalic stages (Haleakala—Chen & Frey, 1985; Mauna Kea—Frey et al., 1991). Only one gabbroic xenolith has major and trace element compositions that indicate that it represents a bulk liquid composition (Figs 4 and 5), and this sample has the shallow, positive REE slope characteristic of tholeiitic-stage magmas [Mauna Loa tholeiitic stage: La 3–38 x chondrite; Mauna Kea alkalic stage: La 38–84 x chondrite (Rhodes, 1996)]. The rest of the gabbroic xenoliths have LREE patterns that are reflective of clinopyroxene accumulation, but the HREE slopes in these xenoliths are similar to the characteristic Mauna Loa HREE slopes. These observations are consistent with a xenolith source of Mauna Loa tholeiitic magmas.
The xenoliths in this study are hosted in two different flows, separated in age by 3300 years, so is it uncertain whether these lava flows and xenoliths came from the same magma chamber and conduit system. However, every xenolith type is represented at each site. The only exceptions are the single troctolite sample from Site A and the plagioclase-bearing peridotites, which are present only at Site B (the younger site). Xenolith mineral compositions also overlap between the two sites. Even if the two flows did not erupt from the same magma chamber, the xenolith-forming processes are similar at the two sites, so all xenoliths are considered together.
Mineral composition constraints on genetic relationships
Figure 11a and b illustrates Cr2O3 vs 100Mg/(Mg + Fe) [molar 100Mg/(Mg + Fe) = mg-number] variability in orthopyroxene and clinopyroxene as a function of xenolith type. For both types of pyroxene, the plagioclase-free peridotite xenoliths have the most primitive compositions (highest Cr2O3 and mg-number). The compositional overlap between clinopyroxene in the plagioclase-free and plagioclase-bearing peridotites contrasts with the two distinct fields for orthopyroxene compositions in these same two xenolith groups. Therefore, the orthopyroxene data, which indicate that the plagioclase-free peridotites formed from a liquid with a higher Mg content and that the plagioclase-bearing xenoliths formed from a lower-Mg liquid, contradict the clinopyroxene data, which suggest that the two peridotite groups crystallized from magma(s) with the same Mg content. The lower mg-number of the orthopyroxene in dunite compared with higher mg-number of orthopyroxene in plagioclase-free peridotite also suggests that the magmas that crystallized these two xenolith families saturated with orthopyroxene at variable Mg content.
|
These observations of mineral assemblages and compositions, as well as xenolith textures, require at least two different crystallization sequences. The primary sequence, which produces most of the observed assemblages, is ol, ol + opx, ol + opx + cpx, ol + opx + cpx + pl, opx + cpx + pl. These families are labeled in Fig. 11a and b. A second crystallization sequence, ol, ol + pl, ol + pl ± cpx ± opx, is required by the few xenolith types that are not represented in the primary crystallization sequence. There are several possible explanations for these two apparent crystallization sequences: (1) different parent compositions; (2) different pressures of crystallization of the same parent magma; (3) post-cumulate processes. These possibilities are evaluated in the following sections.
pMELTS models
Crystallization models from the pMELTS algorithm were used to assess the effects of pressure and starting liquid composition on the crystallization sequence and olivine composition. The pMELTS algorithm (Ghiorso & Hirschmann, 2002) is a revised version of the MELTS algorithm (Ghiorso & Sack, 1995) and is calibrated to an expanded database that includes a larger number of high-pressure experiments. It was designed to model phase stability and compositions in peridotite melting at P >10 kbar. Although the models discussed below were run at P <10 kbar, the starting composition used in the models is primitive and similar to those produced by peridotite melting and therefore acceptable for use with pMELTS (M. S. Ghiorso, personal communication, 2001).
Crystallization of several primitive Mauna Loa near-liquid compositions (Montierth et al., 1995; Rhodes, 1995; Norman & Garcia, 1999) was modeled in pMELTS to assess the dependence of the crystallization sequence on starting composition. At a given pressure, the three starting compositions all produced the same crystallization sequence. The starting composition TLW67-61 (Table 3; Montierth et al., 1995) is the starting composition for all subsequently discussed models. TLW67-61 was used in an experiment by Montierth et al. (1995), and the results of this experiment serve as a test of the ability of pMELTS to model crystallization of this liquid. pMELTS is able to reproduce closely the experimentally determined crystallization sequence, phase compositions and evolution of liquid composition for this starting composition when run at conditions consistent with the experiment [1 kbar and nickel–nickel oxide (NNO) fO2 constraint; Fig. 12 and Table 4]. During the early stages of crystallization, pMELTS underestimates the Fo content of olivine by 1 Fo unit, but as the system moves to lower temperatures, the pMELTS-predicted olivine compositions are nearly identical to the experimentally determined compositions. Olivine mode is also nearly identical between the experiment and the pMELTS prediction.
|
|
|
In the following discussion, pMELTS predictions are compared with observed crystallization sequences and Fo–Ni composition of olivine in the xenoliths. pMELTS does not account for Ni partitioning into pyroxene or plagioclase, and Ni content of the liquid depends upon Ni partitioning into all the phases present. Therefore, Ni partitioning between the liquid and crystals needs to be calculated independently. Using pMELTS-generated liquid major element compositions, crystallization sequence and phase proportions, a bulk crystal–liquid distribution coefficient for Ni was calculated independently at each crystallization increment (1°C) and used to model the Ni content of olivine during crystallization of the melt. DNiol–liq shows a strong inverse correlation with liquid MgO content and is calculated using the experimentally determined method described by Kinzler et al. (1990). Two other methods for calculating DNiol–liq (Hart & Davis, 1978; Beattie, 1991) were compared with the Kinzler et al. (1990) results, and the difference in the end results is minor and does not affect the conclusions. DNiopx–liq shows a much smaller dependence on the MgO content of the liquid and was calculated using the thermodynamic determination of Beattie (1991). Values used for DNicpx–liq and DNipl–liq are two and 0·001 (Steele & Lindstrom, 1981).
pMELTS calculations show that crystallization of the same starting composition at different pressures (2 kbar vs 6 kbar) will result in distinct crystallization sequences. The olivine compositions observed in the plag-bearing peridotite and open-textured gabbroic xenoliths fall within the range of olivine compositions predicted by pMELTS models and independent Ni-partitioning calculations of equilibrium crystallization at pressures ranging from 2 to 6 kbar (Fig. 13b). The pMELTS-predicted 2 kbar sequence is ol, ol + pl, ol + pl + cpx, ol + pl + cpx + opx; this is consistent with the sequence observed in low-pressure experiments of Montierth et al. (1995) on this same starting composition. The predicted 6 kbar sequence is ol, ol + opx, ol + opx + cpx, ol + opx + cpx + pl. The early crystallization of orthopyroxene, which is evident in the poikilitic relationship between olivine and orthopyroxene in the harzburgites, is not apparent in Mauna Loa magmas that have passed through the shallow summit reservoir (Rhodes, 1988). pMELTS models predict that orthopyroxene will be an early-forming phase only at pressures of 6 kbar or greater. This is consistent with Green & Ringwood’s (1967) experiments on Hawaiian tholeiite that show olivine crystallization followed by orthopyroxene at pressures greater than 8 kbar. A second effect of pressure is that at higher pressures, the liquid saturates with the second crystallizing phase at higher liquid MgO content. These models show that neither the variation in crystallization sequence nor the spread in pyroxene compositions requires more than one parent for the plag-bearing peridotite or gabbroic xenoliths.
|
However, the spread in Ni–Fo content of the olivine in plag-free peridotites, dunites and troctolite is greater than that predicted by crystallization from a single parent. In pMELTS models for equilibrium crystallization of a magma saturated only in olivine, the Fo–Ni compositional evolution of olivine has minimal dependence on pressure (Fig. 13a). Some Hawaiian magmas are known to be saturated in sulfides that contain up to 4% Ni (Stone & Fleet, 1991). A total of five small (5–10 µm) sulfide globules were identified in three different xenoliths. Calculations show that 0·01% abundance (Larocque et al., 2000) of a sulfide phase containing 4% Ni will affect the Ni concentration in the magma by <1%. The two crystallization trends illustrated in Fig. 13a are best explained by olivine crystallization from two magmas that have the same major element compositions, but vary in the Ni content (650 ppm Ni vs 500 ppm Ni). Each predicted curve passes through individual xenoliths of both plag-free peridotite and dunite (Figs 7 and 13a), illustrating that the crystallization processes that formed the different xenolith types are consistent across the two batches of parent magmas.
Mineral thermobarometry
Several of the xenoliths contain two pyroxenes, which makes them potentially suitable for the determination of temperature estimates. The geothermometer of Lindsley (1983) was used to calculate temperatures of two holocrystalline gabbroic and four peridotitic xenoliths, based on average orthopyroxene and clinopyroxene compositions within a single xenolith. The temperatures range from 1000 to 1200°C (Table 5). Most of these temperatures are within the range of or slightly lower than the temperature ranges of cpx + opx saturation predicted by pMELTS, which are 1145–1128°C at 2 kbar, and 1264–1212°C at 6 kbar. In the Montierth et al. (1995) experiments, cpx + pig starts crystallizing at 1140°C, at 1 bar. Because the temperatures show such a small range within a given xenolith type, the pyroxene compositions probably are in equilibrium. Most of the xenoliths chosen for the temperature estimates have pyroxenes that show restricted compositional range. The pyroxene compositions for the two plag-bearing lherzolites are variable (maximum range is mg-number 77–86), yet these two samples give the same temperature.
|
Geobarometry can be used to assess the validity of the conclusions about pressure of formation that were drawn from the pMELTS analyses. The Herzberg (1978) and Gasparik (1984) geobarometers both utilize the pressure dependence of the CaTs component in clinopyroxene in the equilibrium system of cpx–opx–ol–pl. However, only two of the xenoliths, both plagioclase-bearing peridotites, have the appropriate assemblages for these calculations. With average clinopyroxene compositions for sample 78-2, the Herzberg geobarometer yields 3 kbar and the Gasparik geobarometer gives pressures of 2 kbar at 1125°C. Pressures for sample 84h are 4·5 kbar according to the Herzberg geobarometer, and 3 kbar by the Gasparik geobarometer (Table 6). The pressures calculated with these geobarometers fall within the range of pressures of crystallization that are consistent with pMELTS predictions of crystallization sequence and olivine compositions.
|
Environments of crystallization
The restricted range of pyroxene compositions and lack of zoning in pyroxene or olivine in any single plagioclase-free peridotite xenolith suggest formation of any one of these xenoliths from a magma that was not substantially modified in any way (such as by magma replenishment) other than by crystallization. Two possible scenarios are: (1) the xenoliths crystallized over such a short time period that the magma was not subjected to any replenishment events; or (2) they crystallized in a small, isolated magma storage area where a batch of magma, once introduced, would crystallize without subsequent reintrusion. The poikilitic texture in these xenoliths, as well as in the troctolite, is consistent with cumulus crystallization in an environment that does not experience frequent physical disruption and chemical resetting such as is common in the large sub-caldera magma chamber (Rhodes, 1988).
The wide range in pyroxene and olivine compositions in many single plag-bearing peridotite or gabbroic xenoliths, the overlap in pyroxene and olivine compositions between these xenolith types, and variable degrees of mineral zoning (both normal and reverse) indicate that either the crystallization sequence or magma composition was disrupted during crystallization (Figs 7–9 and 11). These observations may reflect magma replenishment, or cumulate or post-cumulate processes. Crystal sorting during crystallization has been described as the cause of layering and modal variations at Duke Island (Irvine, 1974). Observations at mafic layered intrusions such as Stillwater or Skaergaard have shown that crystal compositions may not remain in equilibrium with the liquid from which they were formed but may be modified as they reside within the cumulate pile (e.g. McBirney, 1995; Boudreau, 1999). Trapped interstitial liquid can evolve in small pockets and affect the composition of only the crystals with which it is in contact. Evolved interstitial liquid can also migrate through the cumulate pile, forming channels by reaction and mixing with the main magma body to change the liquid composition in the magma chamber (McBirney, 1995). Boudreau (1999) has shown that fluids resulting from inter-cumulus crystallization can flux overlying cumulates and induce remelting. This secondary melt then can recrystallize as a lithology not represented in the primary crystallization sequence. In their study of Kilauea Iki lava lake drill cores, Helz et al. (1989) concluded that diapiric melt transfer of low-density liquids resulted in the transfer of as much as 40 wt % melt out of the crystal mush at the base of the magma lens.
Although processes such as these are probably responsible for some of the variability in mineral compositions in the plag-bearing peridotite and gabbroic xenoliths, it is extremely difficult to evaluate whether these processes are the cause of lithologic diversity across the suite of xenoliths as a whole. There is no stratigraphic framework for the xenoliths, there are no samples of preserved interstitial glass in any of the holocrystalline xenoliths, and the xenoliths themselves are all very small, and give no information on the thickness or lateral continuity of the crystal pile or layer from which they came. Crystal sorting may have caused some of the lithologic diversity, but petrographic and compositional observations still necessitate two crystallization sequences. The location of orthopyroxene in the crystallization sequence (as the second or fourth phase to crystallize) is the primary observation that leads to the conclusion of multiple pressures of crystallization. In the plag-free peridotite xenoliths, the primitive orthopyroxene compositions and the poikilitic relationship between olivine and orthopyroxene indicate that this orthopyroxene is an early-forming phase, and that the ol + opx xenoliths represent a primary assemblage. Observation of very evolved orthopyroxene compositions in xenolith 38, a holocrystalline gabbronorite, suggests that orthopyroxene was a late-forming phase, and the restricted compositional variability and lack of zoning suggest that these orthopyroxene compositions have not been modified subsequent to crystallization by crystal sorting or other cumulate or post-cumulate processes (Fig. 8b). Evolved compositions, limited compositional variability and lack of zoning in clinopyroxene in this xenolith support this interpretation (Fig. 9b).
Compositions of olivine and pyroxene in the open-textured gabbroic xenoliths span a wider range than in any other xenolith group. Abundant interstitial groundmass material and vesicles in these xenoliths suggest that interstitial liquid was present in these xenoliths at the time of eruption, and that these xenoliths formed from a crystal–liquid mush zone in the storage area. Because mineral compositions in the open-textured gabbros show such variability, it is unlikely that the open-textured xenoliths represent infiltration of host magma into a gabbro similar to the holocrystalline gabbros, which typically have more restricted ranges of mineral compositions. A range of processes such as magma replenishment or reaction with interstitial liquid may have contributed to the variability in crystal composition in these xenoliths.
Xenolith formation in the context of Mauna Loa volcanism
The observed range of phase assemblages and compositions in the xenoliths implies complex crystallization, mixing and magma transport processes in Mauna Loa. A hypothesis for xenolith formation must explain distinct crystallization sequences, present a mechanism to separate these magmas in time or space, and incorporate the crystalline remnants into subsequent magmas moving to the surface.
Mineral compositions and crystallization sequences and thermobarometry results are consistent with pMELTS models that indicate crystallization at multiple pressures. This is physically expressed as a network of multiple interconnected levels or ‘tiers’ of magma storage areas. It is unlikely that the primary conduit could be an environment for crystallization at different depths, because this is thought to be occupied by dense, olivine-rich magmas (Garcia et al., 1995), and is inconsistent with the degree of evolution exhibited in some xenoliths.
The magmas parental to the xenoliths may be created or related by a variety of processes. At least one of these magmas must have high MgO (>14%), and they must have variable Ni content. They may represent different primary magmas from the mantle, which form by different degrees of partial melting (Norman & Garcia, 1999), or differentiates from deeper storage areas (Chen, 1993).
The textural and physical requirements indicate magma storage isolated from the primary conduit and summit magma chamber. Rhodes (1988) proposed that Mauna Loa is underlain by a shallow (3 km beneath the summit), long-lived, steady-state magma chamber that maintains homogeneity by buffering through continuous replenishment with more primitive magma. Mauna Loa lavas remain remarkably homogeneous over eruptions of large volume (220 x 106 m3), which take place at the summit caldera as well as along the rift zone [e.g. the 1984 eruption (Rhodes, 1988)]. A large steady-state magma chamber does not provide an environment conducive to the crystallization of the diverse xenoliths.
Small, isolated magma storage area(s) located under Mauna Loa’s SW rift zone and within the mid- to lower oceanic crust are environments in which the xenoliths probably crystallize. Storage areas isolated from primary magma conduits are subject to infrequent replenishment and so are free from physical disruption, which would disturb cumulate textures, and chemical disruption, which would prevent magma differentiation.
Dense, Mg-rich magmas reside at depth within Mauna Loa’s conduit system and occupy storage areas in the oceanic crust and rift zone (Garcia et al., 1995). During episodes of lower magma flux, these storage areas remain isolated and undisturbed as the liquid crystallizes, until the next episode of increased magma flux. Episodes of high magma flux are associated with eruption of picritic magmas such as those that host the xenoliths (Rhodes, 1995). Picritic magmas abandoned in storage areas during subsequent decrease in eruption rate can crystallize into another generation of xenoliths.
|
CONCLUSIONS |
|---|
|
TOPABSTRACTINTRODUCTIONRESULTSDISCUSSIONCONCLUSIONSAPPENDIX: PETROGRAPHYREFERENCES |
|---|
This suite of xenoliths is the result of crystallization of a Mauna Loa tholeiitic-stage magma. The range in olivine compositions in the xenoliths is consistent with crystallization from liquids with initial MgO contents of
>14 wt %, but variable Ni content. The variability in orthopyroxene and clinopyroxene composition across multiple xenolith mineral assemblages indicates that at least two crystallization sequences are necessary to produce the entire assemblage. pMELTS models of crystallization of the same starting composition at 2 and 6 kbar will result in the two different observed crystallization sequences and different evolution paths of the olivine compositions. Slight variability in parent composition and crystallization over a range of pressures (2–6 kbar) produced the spread of olivine and pyroxene compositions and an array of xenolith assemblages. This, along with the textural diversity, implies that the xenoliths are products of complex recharge, mixing and crystallization processes. The crystallization and emplacement of these xenoliths are related to periods of increased magma supply and eruption rate. As the magma supply rate increases, dense, MgO- and olivine-rich magma that normally resides in the lower regions of the magma plumbing system moves to higher elevations, and fills isolated magma storage areas. As the magma supply rate decreases, fresh magma no longer enters the isolated storage area, and magma that remains in the abandoned storage area crystallizes, forming the observed range of xenolith types. Repetition of this cycle of introduction of magma, isolation from the main conduit, and crystallization produces the range of mineral compositions and assemblages seen in the xenoliths. Subsequent eruptions scour the xenoliths from the magma storage area and carry the xenoliths to the surface.
|
APPENDIX: PETROGRAPHY |
|---|
|
TOPABSTRACTINTRODUCTIONRESULTSDISCUSSIONCONCLUSIONSAPPENDIX: PETROGRAPHYREFERENCES |
|---|
Gabbroic xenoliths
Rock types include gabbro, gabbronorite, olivine gabbro, and olivine gabbronorite. Textures vary significantly within all the gabbroic xenoliths, but most can be classified in one of two textural families: those exhibiting open textures with abundant interstitial glass and/or fine-grained basalt, and those that are holocrystalline. Some holocrystalline xenoliths also exhibit a laminated texture. The holocrystalline xenoliths typically are much more cohesive and larger than the open-textured gabbroic or any of the ultramafic xenoliths. Every xenolith larger than
5 cm is gabbroic, holocrystalline and typically laminated.
Gabbro (samples 19a, 65, 71)
Gabbros have both open and holocrystalline textures. Sample 19a has blocky, rectangular, or lath-shaped plagioclase, 1–2 mm in length, that exhibits resorbed textures. Some plagioclase crystals enclose clinopyroxene, which is subhedral and 0·6–1·2 mm long. This sample has an open texture with pockets of host (groundmass and glass) and vesicles throughout the xenolith.
In samples 65 and 71, plagioclase exhibits a spongy texture. Crystals are rectangular to blocky and <1 mm long. Clinopyroxene is anhedral, with embayed edges and irregular fractures. There is a faint foliation as a result of alignment of crystals. A small amount of Cr-spinel is present.
Olivine gabbro (105)
In this open-textured sample, plagioclase is euhedral to anhedral, and has rectangular to blocky shapes, and sizes from 2·3 to 0·4 mm. Clinopyroxene is 1 mm long, fractured, and has irregular shapes. Olivine is 0·8–2 mm long, and is variably fractured, fragmented, rounded or embayed.
Gabbronorite (31, 38, 70, 73, 78c, 78f)
The two open-textured gabbronorites (78c, 78f), have plagioclase crystals that are euhedral to subhedral, rectangular to blocky, and 0·6–2 mm long. Clinopyroxene crystals are weathered, embayed and amoeboid (rounded and irregularly shaped), and 0·2–2 mm long. Orthopyroxene grains are 0·2–1·2 mm long, anhedral, fractured and altered. Pockets of microcrystalline glass are enclosed within both xenoliths. In places the xenoliths have disaggregated, and small crystal fragments were incorporated in the groundmass.
The holocrystalline gabbronorites (31, 38, 70) have plagioclase that is rectangular to blocky, 0·5–3 mm long, and exhibits a spongy texture. Clinopyroxene is 0·6–2 mm long, fractured, irregularly shaped, and contains melt inclusions. Orthopyroxene crystals are amoeboid, fractured, 0·4–0·8 mm long, and are filled with melt inclusions. Some xenoliths are weakly laminated, as defined by oriented plagioclase (Fig. 2g).
The single fine-grained gabbronorite (73) has uniform grain size of 0·2 mm. Plagioclase is euhedral to subhedral, and blocky to rectangular. Pyroxene is subhedral and rounded to blocky (Fig. 2h).
Olivine gabbronorite (37, 47p, 48a, 51, 76, 79b)
All of these samples are open textured. Plagioclase is subhedral to anhedral, rectangular to lath shaped, has corroded edges, and is 0·1–1·2 mm long. Pyroxene is embayed, irregularly shaped, 1–2·5 mm long, fractured and has abundant melt inclusions. Olivine size is variable (0·5–3 mm) and shapes range from euhedral to fractured, embayed and resorbed. Pockets of groundmass and abundant vesicles are enclosed in the xenoliths (Fig. 2f). The grains that are in contact with the host generally have more irregular grain boundaries.
Ultramafic rocks
Dunite (47d, 47s, 56, 66, 78b)
Olivine crystals range in size from 0·3 to 2·2 mm and have shapes that vary from rounded, amoeboid and embayed to euhedral or fractured (Fig. 2a). Crystals are kink-banded to different extents. Olivine commonly encloses Cr-spinel (0·1 mm). Several samples have interstitial plagioclase, and one sample (56) has interstitial orthopyroxene. Sample 47d is fine grained (0·05–0·3 mm) with a granular texture. Olivine may be rounded, resorbed or euhedral. Interstitial glass is locally present. Most of the xenoliths are disaggregated at the edges or have reacted with the melt.
Sample 66 has a granoblastic texture and olivines have a bimodal size distribution; large crystals are 0·4–1·5 mm, strongly kink-banded, elongate, slightly aligned and have scalloped edges, whereas small crystals are 0·02–0·1 mm, rounded, exhibit a granular texture and commonly are kink-banded. Cr-spinel in this xenolith is not enclosed by olivine crystals.
Troctolite (48c)
Olivine in the troctolite ranges from 0·5 to 4 mm in length, and is rounded, resorbed, euhedral or fractured. Some is kink-banded. Plagioclase is interstitial, but can be very large (up to 6 mm) and have a sub-poikilitic relationship with olivine (Fig. 2b). The few clinopyroxene crystals are 4 mm long, and poikilitically enclose olivine.
Harzburgite or lherzolite (48-3, 48g, 48i, 78q, 79a, 79c, 79d, 79e)
In these xenoliths one or several orthopyroxene crystal(s) poikilitically enclose olivine (Fig. 2c). Olivine grains are 0·2–4 mm in the long dimension; some are strained or fractured, and most are anhedral, amoeboid or resorbed. Orthopyroxene crystals are large, up to 15 mm x 14 mm, fractured, and have rounded grain boundaries. In some cases, the orthopyroxene crystal constitutes the whole xenolith, and is probably a remnant of an even larger crystal. Only one sample, 79a, contains clinopyroxene (Fig. 2d).
Plagioclase-bearing harzburgite or lherzolite (78-2, 84h, 84l, 85a)
Xenolith 78-2 lacks the poikilitic texture characteristic of the other ultramafic xenoliths. Small olivine crystals (0·5–1 mm) are rounded and enclosed in clinopyroxene; larger (3 mm) crystals are more angular, fractured, and euhedral. Clinopyroxene and orthopyroxene are large (1·5–7 mm), very fractured and with variable, but mostly irregular, shapes. Plagioclase is primarily interstitial, but there are some larger euhedral crystals that range in size from 1·5 to 2·5 mm.
Olivine in 84l is 0·3–3·5 mm in diameter, rounded and irregular, and locally fractured. Orthopyroxene and clinopyroxene is 1·5–2 mm, altered and fractured. Plagioclase is present as several blocky crystals (Fig. 2e).
Sample 84h is a unique xenolith with the same mineral assemblages as 78-2, 79a and 84l but a range of textures. Olivine, which is rounded, amoeboid, irregularly shaped and fractured, is poikilitically enclosed by pyroxene in one half of the xenolith and by plagioclase in the other half. The olivine enclosed in pyroxene is generally more rounded than that enclosed in plagioclase. A few olivine crystals are strained, and range in size from 0·6 to 2 mm. Orthopyroxene and clinopyroxene are 7 mm, fractured, and with irregular edges. Plagioclase is mostly interstitial, although some shows a blocky shape and encloses some olivine. Cr-spinel is enclosed in pyroxene, olivine and plagioclase.
Olivine in 85a has linear trains of melt or vapor inclusions that cut across multiple crystals. There are also linear concentrations of small (0·03–0·3 mm) orthopyroxene in olivine crystals. Olivine is 1·8 mm across and rounded to angular, although all olivine is anhedral. Some of the olivine crystals are strained. Orthopyroxene is 2 mm, anhedral and some encloses olivine. Plagioclase is 0·2–1 mm, interstitial and anhedral.
Olivine megacrysts (48o, 48f)
These single crystals of olivine are up to 2·7 cm long, and have inclusions of Cr-spinel and glass. The crystals are fractured and overall are very fresh. Crystals are dominantly euhedral, but are embayed in places. Most megacrysts are kink-banded.
Host basalt, olivine basalt and picrite (19c, 26, 27, 28, 48, 82, 83, 84, 86)
Host samples are highly variable in vesicularity, glass and olivine content, depending on location within the flow, although samples are typically highly vesicular and have abundant olivine phenocrysts. Phenocrysts are generally euhedral, but are very fractured and can show resorption along the edges. They range in size from 1 to 6 mm. Glass content ranges as a function of proximity to the edge of the flow. Plagioclase and clinopyroxene phenocrysts or xenocrysts are rare. Clinopyroxene is very similar to olivine in shape and size. Plagioclase is <2 mm and very corroded. Groundmass minerals are plagioclase, clinopyroxene, Cr-spinel and glass. Groundmass in some samples is slightly trachytic. Others show dendritic patterns in the groundmass spinel resulting from rapid cooling of the lava.
|
ACKNOWLEDGEMENTS |
|---|
The author is grateful to Frank Trusdell, Don Swanson, and the staff at the Hawaiian Volcano Observatory. The manuscript has benefited greatly from discussions with Stu McCallum and Bruce Nelson, and thorough reviews by Wendy Bohrson, Dennis Geist, Mike Rhodes and Gautam Sen. Field work for this project was completed while the author was a volunteer at HVO and was supported by a Geological Society of America Harold T. Stearns Fellowship, a University of Washington Geological Sciences Graduate Student Research Fund grant and the University of Washington Geological Sciences P. Misch Fellowship.
|
FOOTNOTES |
|---|
Extended dataset can be found at http://www.petrology.oupjournals.org
*Telephone: 206-543-1975. Fax: 206-543-3836. E-mail: agaffney{at}u.washington.edu
|
REFERENCES |
|---|
|
TOPABSTRACTINTRODUCTIONRESULTSDISCUSSIONCONCLUSIONSAPPENDIX: PETROGRAPHYREFERENCES |
|---|
Armstrong, J. T. (1988). Quantitative analysis of silicate and oxide materials: comparison of Monte Carlo, ZAF, and phi(rho z) procedures. In: Newbury, D. E. (ed.) Microbeam Analysis. San Francisco, CA: San Francisco Press, pp. 239–246.
Beattie, P. (1991). Partition coefficients for olivine–melt and orthopyroxene–melt systems. Contributions to Mineralogy and Petrology 109, 212–224.
Bohrson, W. A. & Clague, D. A. (1988). Origin of ultramafic xenoliths containing exsolved pyroxenes from Hualalai Volcano, Hawaii. Contributions to Mineralogy and Petrology 100, 139–155.[Web of Science]
Boudreau, A. (1999). Fluid fluxing of cumulates: the J-M Reef and associated rocks of the Stillwater Complex, Montana. Journal of Petrology 40, 755–772.
Chen, C.-Y. (1993). High-magnesium primary magmas from Haleakala Volcano, east Maui, Hawaii: petrography, nickel, and major-element constraints. Journal of Volcanology and Geothermal Research 55, 143–153.
Chen, C.-Y. & Frey, F. A. (1985). Trace element and isotopic geochemistry of lavas from Haleakala Volcano, East Maui, Hawaii: implications for the origin of Hawaiian basalts. Journal of Geophysical Research 90, 8743–8768.
Chen, C.-H., Presnall, D. C. & Stern, R. J. (1992). Petrogenesis of ultramafic xenoliths from the 1800 Kaupulehu flow, Hualalai volcano, Hawaii. Journal of Petrology 33, 163–202.
Clague, D. A. (1987). Hawaiian xenolith populations, magma supply rates, and development of magma chambers. Bulletin of Volcanology 49, 577–587.
Clague, D. A. & Chen, C.-H. (1986). Ocean crust xenoliths from Hualalai Volcano, Hawaii. Geological Society of America, Abstracts with Programs 18, 565.
Clague, D. A., Moore, J. G., Dixon, J. E. & Freisen, W. E. (1995). Petrology of submarine lavas from Kilauea’s Puna Ridge, Hawaii. Journal of Petrology 36, 299–349.
Fodor, R. V. & Galar, P. (1997). A view into the subsurface of Mauna Kea volcano, Hawaii: crystallization processes interpreted through the petrology and petrography of gabbroic and ultramafic xenoliths. Journal of Petrology 38, 581–624.
Frey, F. A., Wise, W. S., Garcia, M. O., West, H., Kwon, S.-T. & Kennedy, A. (1991). Evolution of Mauna Kea volcano, Hawaii: petrologic and geochemical constraints on postshield volcanism. Journal of Geophysical Research 95, 1271–1300.[Web of Science]
Garcia, M. O., Hulsebosch, T. P. & Rhodes, J. M. (1995). Olivine-rich submarine basalts from the southwest rift zone of Mauna Loa Volcano: implications for magmatic processes and geochemical evolution. In: Rhodes, J. M. & Lockwood, J. P. (eds) Mauna Loa Revealed: Structure, Composition, History and Hazards. Washington, DC: American Geophysical Union, pp. 219–239.
Gasparik, T. (1984). Two-pyroxene thermobarometry with new experimental data in the system CaO–MgO–Al2O3–SiO2. Contributions to Mineralogy and Petrology 87, 87–97.[Web of Science]
Ghiorso, M. S., Hirschmann, M. M., Reiners, P. & Kress, V. C. (2002). pMELTS: A revision of MELTS for improved calculation of phase relations and major element partitioning related to partial melting of the mantle to 3 GPa. Geochemistry, Geophysics, Geosystems in press.
Ghiorso, M. S. & Sack, R. O. (1995). Chemical mass transfer in magmatic processes IV. A revised and internally consistent thermodynamic model for the interpolation and extrapolation of liquid–solid equilibria in magmatic systems at elevated temperatures and pressures. Contributions to Mineralogy and Petrology 119, 197–212.[Web of Science]
Green, D. H. & Ringwood, A. E. (1967). The genesis of basaltic magmas. Contributions to Mineralogy and Petrology 15, 103–190.
Hart, S. R. & Davis, K. E. (1978). Nickel partitioning between olivine and silicate melt. Earth and Planetary Science Letters 40, 203–219.[Web of Science]
Hart, S. R. & Dunn, T. (1993). Experimental cpx/melt partitioning of 24 trace elements. Contributions to Mineralogy and Petrology 113, 1–8.[Web of Science]
Helz, R. T., Kirschenbaum, H. & Marinenko, J. W. (1989). Diapiric transfer of melt in Kilauea Iki lava lake, Hawaii: a quick, efficient process of igneous differentiation. Geological Society of America Bulletin 101, 578–594.
Herzberg, C. T. (1978). Pyroxene geothermometry and geobarometry: experimental and thermodynamic evaluation of some subsolidus phase relations involving pyroxenes in the system CaO–MgO–Al2O3–SiO2. Geochimica et Cosmochimica Acta 42, 945–957.
Irvine, T. N. (1974). Petrology of the Duke Island ultramafic complex, southeast Alaska. Geological Society of America, Memoir 138.
Jackson, E. D., Clague, D. A. & Beeson, M. H. (1982). Miscellaneous Hawaiian xenolith localities. US Geological Survey Open-File Report 82-304.
Johnson, D. M., Hooper, P. R. & Conrey, R. M. (1999). XRF analysis of rocks and minerals for major and trace elements on a single low dilution Li-tetraborate fused bead. Advances in X-Ray Analysis 41, 843–867.
Kinzler, R. J., Grove, T. L. & Recca, S. I. (1990). An experimental study on the effect of temperature and melt composition on the partitioning of nickel between olivine and silicate melt. Geochimica et Cosmochimica Acta 54, 1255–1265.
Larocque, A. C. L., Stimac, J. A., Keith, J. D. & Huminicki, M. A. E. (2000). Evidence for open-system behavior in immiscible Fe–S–O liquids in silicate magmas: implications for contributions of metals and sulfur to ore-forming fluids. Canadian Mineralogist 38, 1233–1249.[Web of Science]
Lindsley, D. H. (1983). Pyroxene thermometry. American Mineralogist 68, 477–493.[Abstract]
McBirney, A. R. (1995). Mechanisms of differentiation in the Skaergaard Intrusion. Journal of the Geological Society, London 152, 421–435.
McKay, G. A., Wagstaff, J. & Yang, S.-R. (1986). Clinopyroxene REE distribution coefficients for shergottites: the REE content of the Shergotty melt. Geochimica et Cosmochimica Acta 50, 927–937.[Web of Science]
Montierth, C., Johnston, A. D. & Cashman, K. V. (1995). An empirical glass-composition-based geothermometer for Mauna Loa lavas. In: Rhodes, J. M. & Lockwood, J. P. (eds) Mauna Loa Revealed: Structure, Composition, History and Hazards. Washington, DC: American Geophysical Union, pp. 207–217.
Norman, M. D. & Garcia, M. O. (1999). Primitive magmas and source characteristics of the Hawaiian plume: petrology and geochemistry of shield picrites. Earth and Planetary Science Letters 168, 27–44.
Papike, J. J., Spilde, M. N., Fowler, G. W. & McCallum, I. S. (1995). SIMS studies of planetary cumulates: orthopyroxene from the Stillwater Complex, Montana. American Mineralogist 80, 1208–1221.[Abstract]
Rhodes, J. M. (1983). Homogeneity of lava flows: chemical data for historic Mauna Loan eruptions. Journal of Geophysical Research 88A, 869–879.
Rhodes, J. M. (1988). Geochemistry of the 1984 Mauna Loa eruption: implications for magma storage and supply. Journal of Geophysical Research 93, 4453–4466.
Rhodes, J. M. (1995). The 1852 and 1868 Mauna Loa picrite eruptions: clues to parental magma compositions and the magmatic plumbing system. In: Rhodes, J. M. & Lockwood, J. P. (eds) Mauna Loa Revealed: Structure, Composition, History and Hazards. Washington, DC: American Geophysical Union, pp. 241–262.
Rhodes, J. M. (1996). Geochemical stratigraphy of lava flows sampled by the Hawaii Scientific Drilling Project. Journal of Geophysical Research 101, 11729–11746.
Roeder, P. L. & Emslie, R. F. (1970). Olivine–liquid equilibrium. Contributions to Mineralogy and Petrology 29, 275–289.[Web of Science]
Sen, G. & Leeman, W. P. (1991). Iron-rich lherzolitic xenoliths from Oahu: origin and implications for Hawaiian magma sources. Earth and Planetary Science Letters 102, 45–57.
Steele, I. M. & Lindstrom, D. J. (1981). Ni partitioning between diopside and silicate melt: a redetermination by ion microprobe and recognition of an experimental complication. Geochimica et Cosmochimica Acta 45, 2177–2183.
Stone, W. E. & Fleet, M. E. (1991). Nickel–copper sulfides from the 1959 eruption of Kilauea Volcano, Hawaii: contrasting compositions and phase relations in eruption pumice and Kilauea Iki lava lake. American Mineralogist 76, 1363–1372.[Abstract]
Sun, S.-s. & McDonough, W. F. (1989). Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: Saunders, A. D. & Norry, M. J. (eds) Magmatism in the Ocean Basins. Geological Society, London, Special Publications 42, 313–345.
Tilling, R. I., Wright, T. L. & Millard, H. T., Jr (1987). Trace-element chemistry of Kilauea and Mauna Loa in space and time: a reconnaissance. In: Decker, R. W., Wright, T. L. & Stauffer, P. H. (eds) Volcanism in Hawaii. US Geological Society, Professional Paper 1350, 641–689.
White, R. W. (1966). Ultramafic inclusions in basaltic rocks from Hawaii. Contributions to Mineralogy and Petrology 12, 245–314.
CiteULike 
Connotea 
Del.icio.us What's this?
This article has been cited by other articles:
|
|
 |
|
  M. Gaeta, T. Di Rocco, and C. Freda Carbonate Assimilation in Open Magmatic Systems: the Role of Melt-bearing Skarns and Cumulate-forming Processes J. Petrology, February 1, 2009; 50(2): 361 - 385. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
T. Rehfeldt, D. E. Jacob, R. W. Carlson, and S. F. Foley Fe-rich Dunite Xenoliths from South African Kimberlites: Cumulates from Karoo Flood Basalts J. Petrology, July 1, 2007; 48(7): 1387 - 1409. [Abstract] [Full Text] [PDF]
|
|
|
|
|
|
P. J. SHAMBERGER and J. E. HAMMER Leucocratic and Gabbroic Xenoliths from Hualalai Volcano, Hawai'i J. Petrology, September 1, 2006; 47(9): 1785 - 1808. [Abstract] [Full Text] [PDF]
|
|
| ||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||||