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Journal of Petrology Volume 42 Number 9 Pages 1685-1704 2001
© Oxford University Press 2001
The Role of Tonalite and Diorite in Mauna Kea Volcano, Hawaii, Magmatism: Petrology of Summit-Region Leucocratic Xenoliths
DEPARTMENT OF MARINE, EARTH, AND ATMOSPHERIC SCIENCES, NORTH CAROLINA STATE UNIVERSITY, RALEIGH, NC 27695-8208, USA
Received January 10, 2000; Revised typescript accepted October 12, 2000
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONBACKGROUNDSUMMIT REGION LEUCOCRATIC...SUMMIT-REGION GABBRO AND...ANALYTICAL TECHNIQUESANALYTICAL RESULTSDISCUSSIONREFERENCES |
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A tonalite (
66 wt % SiO2; 2·4 wt % K2O) xenolith from the Mauna Kea summit region provides information on the origin of silicic liquids in mafic magma systems. This leucocratic rock has 40 vol. % quartz interstitial to and enclosing Ca-andesine, plus phlogopite (11 vol. %), clinopyroxene (mg-number 74), orthopyroxene, and Fe–Ti oxides. Also, it contains lithic fragments (
5 cm) of gabbro (MgO 7·2 wt %; An60–50; clinopyroxene mg-number 78–75; phlogopite). The tonalite has Sr–Nd–Pb isotopic ratios of 87Sr/86Sr 0·703610, 144Nd/143Nd 0·512976, 206Pb/204Pb 18·58, 207Pb/204Pb 15·49, and 208Pb/204Pb 38·15, which agree with the isotopic composition of Mauna Kea post-shield Hamakua Volcanics (tholeiitic and alkalic basalts). A positive Eu anomaly and poikilitic texture indicate a cumulate origin. Leucocratic diorite (50–53 wt % SiO2; 1–2·4 wt % K2O) xenoliths coexist with the tonalite. These have intergranular andesine–oligoclase (68–85 vol. % of An50–15), evolved clinopyroxene (mg-number 76–69), and biotite (± orthopyroxene, zircon). The tonalite represents SiO2-rich liquids that accumulated, perhaps in a small reservoir, after segregation from a hydrous basalt-magma solidification front; gabbro inclusions within the tonalite are probably remnants of that source. Hydrous conditions (implied by abundant phlogopite) were necessary for this silicic-liquid segregation; namely, relatively high H2O facilitated crystallization of SiO2-deficient phases (oxide; phlogopite) to leave SiO2-enriched residual liquids, and H2O vapor pressure helped segregate silicic liquids by gas-pressured filter pressing. The diorites are coarse-grained equivalents of hawaiite and mugearite produced by clinopyroxene-dominant crystallization of Hamakua alkalic basalt. On the basis of normative Ol–Di–Ne, diorite xenoliths originated under intermediate and high pressures, apparently reflecting the transition from low- to high-P crystallization regimes that are identified for Mauna Kea post-shield magmas (where the high-P regime is approximately at the crust–mantle boundary). Leucocratic xenoliths, particularly tonalite, provide additional insights into Hamakua post-shield magma evolution by manifesting differentiation that is more extensive and complex, approaching granitic compositions, than that represented by Hamakua lavas, and by showing that quartz can crystallize appreciably during Hawaiian magmatism. Quartz-bearing rock on Hawaii has global relevance in terms of providing better understanding of the circumstances for oceanic basalt–rhyolite magmatism. KEY WORDS: xenoliths; tonalite; diorite; gabbro; Mauna Kea, cumulate, quartz
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INTRODUCTION |
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In volcanic settings, subsurface plutonic environments can contain rocks that represent the compositional extremes to which the parental magmas evolved; that is, plutonic rocks of volcanic interiors may represent thermal minima and eutectic liquids that mark the final differentiation stages reached as subsolidus conditions were achieved. This appears to be the situation at Mauna Kea volcano, Hawaii. Among the varieties of plutonic-rock xenoliths associated with post-shield stage volcanism at Mauna Kea are leucocratic xenoliths—namely, tonalite and plagioclase-rich diorite—that were erupted at a cinder cone near the summit (Fig. 1). On the basis of their mineral compositions, these leucocratic xenoliths are compositionally evolved relative to Hawaiian lavas overall, and they represent some of Mauna Kea’s most extreme magmatic differentiation. The presence of these felsic rocks in an all-mafic oceanic volcanic province provokes investigation for their reconciliation with basaltic magmas.
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Mauna Kea has yielded a variety of ultramafic and gabbroic cumulate-origin xenoliths (e.g. Jackson et al., 1982
; Fodor & Vandermeyden, 1988; Fodor & Galar, 1997) that represent the magmas that erupted as post-shield lavas. These xenoliths have a basaltic modal mineralogy that is limited in compositional range to plagioclase with (
An50 and clinopyroxene with mg-number (76 (Fodor & Galar, 1997; Hoover & Fodor, 1997; Fodor, 2000). In contrast, the plutonic leucocratic xenoliths have modal andesine, oligoclase, sanidine, phlogopite, biotite, quartz, and zircon—minerals that are uncharacteristic of Mauna Kea post-shield lavas and, for the most part, of Hawaiian magmas in general. These minerals represent more evolved compositional paths than those manifested by post-shield volcanism. The leucocratic xenoliths are the focus of this study, and through their petrography and whole-rock and mineral compositions provide insights into reservoir crystallization processes that are not manifested by erupted magmas. Of particular interest is the presence of quartz-bearing tonalite in an intraplate oceanic environment, as it documents the origin of localized SiO2-rich liquids that could feasibly spawn granitoid continent-like crust within oceanic lithosphere.
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BACKGROUND |
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Geologic setting
Mauna Kea volcano is one of five shield volcanoes on the island of Hawaii (Fig. 1). It comprises a tholeiitic shield capped by post-shield lavas of both tholeiitic and alkalic affinities (West et al., 1988
; Frey et al., 1990, 1991; Wolfe et al., 1997) and reaches an altitude of 4206 m. Only post-shield Mauna Kea lavas are subaerially exposed and they are divided into two units: (1) Hamakua Volcanics (250–65 ka), which are tholeiitic, transitional, alkalic, and Fe–Ti basalts (e.g. Frey et al., 1990, 1991; Wolfe et al., 1997); (2) Laupahoehoe Volcanics (65–4 ka), comprising largely hawaiite and to a small extent mugearite and benmoreite (West et al., 1988; Wolfe et al., 1997). In Mauna Kea post-shield syntax, Hamakua lavas form the ‘basaltic substage’ and Laupahoehoe lavas form the ‘hawaiitic substage’ (Frey et al., 1990; Wolfe et al., 1997).
Laupahoehoe volcanism created dozens of cinder cones and lavas on the upper flanks from 2800 m elevation up to the peak. Many of the Laupahoehoe cones and lavas contain gabbroic and ultramafic xenoliths 4–30 cm in size that represent cumulates of the earlier Hamakua post-shield magmas (Jackson et al., 1982; Fodor & Vandermeyden, 1988; Fodor & Galar, 1997).
Mauna Kea xenoliths
Summit region
Xenoliths from the summit cone (4200 m) have been described by Fodor & Vandermeyden (1988). One type of summit-cone xenolith is olivine gabbro with an anhedral granular texture, some samples of which have subparallel plagioclase laths that define a planar, or laminated, texture. These olivine gabbros grade into anhedral-granular wehrlite with <5 vol. % interstitial plagioclase. A second type of summit-cone xenolith is opaque-oxide gabbro. This variety is composed essentially of plagioclase, clinopyroxene, and Fe–Ti oxides (olivine is rare), where subparallel plagioclase grains define a planar texture. Titaniferous magnetite and ilmenite are intergranular and occur in amounts of 10–30 vol. %. Both types of summit cone xenoliths have parentages in transitional and/or alkalic Hamakua basaltic magmas.
Southern flank
Xenoliths from the southern flank at 2900 m elevation have been described by Fodor & Galar (1997) and Hoover & Fodor (1997). They are dunites, wehrlites, gabbros, gabbronorites, and troctolites. Most ultramafic xenoliths have porphyroclastic textures and the gabbroic xenoliths are largely anhedral granular. Among the spectrum of gabbroic xenoliths, many are frameworks of olivine and clinopyroxene with interstitial plagioclase, and others are frameworks of plagioclase with intergranular olivine and clinopyroxene. Relatively fine-grained, equigranular mosaic-textured gabbros are also present. Flank xenoliths have tholeiitic (e.g. many contain orthopyroxene), alkalic, and transitional Hamakua magma parentages.
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SUMMIT REGION LEUCOCRATIC XENOLITHS |
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The leucocratic xenoliths are from one of five Mauna Kea summit-region cones that yield ultramafic and gabbroic xenoliths. This cone has no official name and is called cone ‘D’ for field-sampling purposes. It is at
3850 m elevation (Fig. 1).
Tonalite
The tonalite xenolith is 15 cm in size (Fig. 2a). Because the sample is composed of 30 vol. % dark, lithic fragments that range in size from 0·5 to 5 cm (Fig. 2a), it forms a composite xenolith of two rock types. Millimeter-sized veinlets of tonalite penetrate the margins of the dark lithic inclusions (Fig. 2a).
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Tonalite modal mineralogy is largely quartz and plagioclase, each
40 vol. % (Table 1). Microprobe analyses reveal that the plagioclase is mainly calcic andesine. The presence of andesine plus abundant modal quartz forms the basis for the tonalite classification. The texture is largely subhedral to anhedral granular among plagioclase grains, 0·2–1 mm, some of which occur in triple-juncture relationships (Fig. 2b). Some plagioclase is poikilitically enclosed by quartz, which is anhedral and interstitial to the plagioclase grains (Fig. 2b). Phlogopite makes up 11 vol. %, largely occupying interstitial spaces and containing inclusions of Fe–Ti oxide and, less commonly, plagioclase. Fe–Ti oxides, 5 vol. %, are intergranular among the plagioclase grains and also occur within biotite. Clinopyroxene and orthopyroxene are interstitial to plagioclase and make up <2 vol. % of the tonalite. Sanidine is a rare interstitial phase, identified more easily during electron microprobe analyses than optically.
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The dark lithic fragments within the tonalite (Fig. 2a) are fine-grained gabbro occurring as equigranular (0·1–0·3 mm) mosaics of plagioclase, clinopyroxene, orthopyroxene, phlogopite, Fe–Ti oxides, and rare amphibole (Fig. 2c). Plagioclase makes up 39 vol. % and clinopyroxene and orthopyroxene combined are 37 vol. % of the gabbro (Table 1). Fe–Ti oxides are 15 vol. %, and phlogopite, largely interstitial, is 8 vol. %.
Diorite
Five leucocratic diorite xenoliths from cone D (Fig. 1) range from 3 to 7 cm in size. Plagioclase is andesine and oligoclase in composition and is generally present in amounts >70 vol. % (Table 1). Plagioclase (largely 0·5–3 mm), clinopyroxene, biotite, and Fe–Ti oxides occur in subhedral to anhedral granular and intergranular textures (Fig. 2d and e). One diorite has a planar plagioclase orientation (Fig. 2f), and another has local areas that are planar. Fe–Ti oxides occur as inclusions in plagioclase, clinopyroxene, and biotite, as well as intergranular to plagioclase. Apatite occurs as rods within plagioclase and biotite in volumes up to 0·3%. Two xenoliths contain zircon, up to 1 mm (Fig. 2g), and two diorites contain orthopyroxene.
One diorite xenolith has veinlets of vesicular, brown glass that appears to represent melt from an external source (Fig. 2g). Another diorite has glass interstitial to plagioclase grains that appears to represent liquid trapped amongst the grains. The biotite in most diorites appears to be primary in origin. It occurs as an intergranular phase in three diorites, largely within the glass veinlets of one, but appears to be secondary after clinopyroxene in another.
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SUMMIT-REGION GABBRO AND WEHRLITE |
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Because the tonalite and diorite compositions are best evaluated within the context of other xenolith types with which they occur near Mauna Kea summit, I also examined 25 summit-region mafic and ultramafic xenoliths from cone D and four neighboring cones: ‘summit’ cone (S), Poliahu cone (P), Puu Hau Kea cone (HK), and another unnamed cone labeled ‘E’ in Fig. 1. Their elevations range over 600 m, from the summit westward to
3550 m (Fig. 1). Among these xenoliths, 21 are anhedral-granular gabbros (Fig. 2h), with grain sizes largely 0·5–5 mm but some with grains up to 1 cm, and four are anhedral-granular plagioclase wehrlites. Textures in half of these gabbros are weakly overprinted by lamination of subparallel plagioclase laths. Many of these xenoliths represent rock types and textures similar to those reported for summit-cone xenoliths by Fodor & Vandermeyden (1988). The modal mineralogy of representative gabbros and wehrlites is given in Table 1.
One gabbro xenolith contains an 0·5 cm vein (Fig. 2h) comprising largely subparallel plagioclase, 0·5–2 mm, in intergranular relationship with biotite, and minor vesicular microcrystalline glass. It represents penetration by the host lava that entrained the xenolith, a Laupahoehoe hawaiite–mugearite.
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ANALYTICAL TECHNIQUES |
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Mineral compositions were determined using an ARL-SEMQ electron microprobe at North Carolina State University using olivine, clinopyroxene, orthopyroxene, amphibole, plagioclase, microcline, spinel, and ilmenite provided by the Smithsonian Institution and Ni-doped diopside as reference minerals. Phi-rho-Z matrix corrections were applied. Whole-rock samples were analyzed for major- and trace-element abundances on a Philips 1410 X-ray fluorescence spectrometer at NCSU. Major elements (except Na, P) were analyzed using glass disks; trace elements (except REE, Sc, Hf, Th) and Na and P were determined on pressed powders. Reference samples for calibration curves include standards of the US, Canadian, South African, and Japanese geological surveys. Precision for this instrumentation has been reported by Fodor et al. (1992)
. Rare earth elements, Sc, Hf, and Th were determined by neutron activation at the Oregon State University Radiation Center. Sr, Nd, and Pb isotope compositions for the tonalite xenolith were obtained at the University of North Carolina isotope facility, using a VG Sector 54 mass spectrometer. Conditions for analyses are as described by Dobosi et al. (1995). |
ANALYTICAL RESULTS |
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Mineral compositions
Table 2 presents the mineral composition data for the leucocratic xenoliths, the gabbro lithic fragments within the tonalite, and the plagioclase–biotite vein and its host gabbro (Fig. 2h). Various aspects of these mineral compositions, plus those in the 25 other summit-region xenoliths studied here, namely gabbro and wehrlite, are illustrated in Figs 3–6.
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Olivine
Olivine is present in only one leucocratic xenolith, diorite D1, and it is highly evolved, Fo63. For comparison, olivine compositions in summit-region gabbro and wehrlite xenoliths range between Fo68 and Fo82·5 (Fig. 3a). These higher Fo values are similar to those of other Mauna Kea gabbros and wehrlites, such as the xenoliths erupted along the southern flank (Fodor & Galar, 1997).
Clinopyroxene
The range for average mg-numbers for three clinopyroxene grains in the tonalite is 75·5–74·2 (Fig. 3b). These clinopyroxenes have low Al2O3, 1·4 wt %, and low TiO2, 0·40 wt %, values less than those in clinopyroxene in Hawaiian tholeiitic and alkalic lavas (Fig. 3b and c). Although the tonalite clinopyroxene is evolved in terms of Hawaiian magmatism (i.e. low mg-number), some Mauna Kea submarine lavas have clinopyroxene that is even more evolved, with mg-numbers 71–72 (Yang et al., 1999). That clinopyroxene, however, has higher Al2O3 (>2 wt %) and TiO2 (>0·7 wt %), abundances that are characteristic of clinopyroxenes in Hawaiian tholeiitic lavas (Fig. 3b and c).
Clinopyroxene in the gabbro lithic fragments within the tonalite is similarly to slightly less evolved than the tonalite clinopyroxene. Average mg-numbers for three grains range from 78·1 to 74·8, and they have comparably low Al2O3 and TiO2 abundances (Fig. 3b and c).
Clinopyroxene in the diorites is also evolved, where average mg-numbers for most are within the range 73–76, and is 68·7 in sample D3 (Fig. 3b and c). Diorite clinopyroxene also has low Al2O3 and TiO2 contents, similar to the values for the tonalite clinopyroxene (Fig. 3b and c).
The other summit-region xenoliths, gabbro and wehrlite, have clinopyroxene mg-numbers between 75 and 85 (Fig. 3b and c). These clinopyroxenes have Al2O3 largely between 2 and 4·5 wt % and TiO2 between 0·6 and 1·5 wt %, consistent with Al2O3 and TiO2 in clinopyroxene of tholeiitic and alkalic Hawaiian lavas (Fig. 3b and c), as well as with the clinopyroxene of other Mauna Kea gabbro and wehrlite xenoliths (Fodor & Galar, 1997).
Orthopyroxene
Orthopyroxene occurs in tonalite, its gabbro inclusions, some diorites, and in some summit-region gabbro xenoliths. Orthopyroxenes in the leucocratic xenoliths have relatively evolved mg-numbers that are lower than those in corresponding clinopyroxenes (Fig. 3d). These evolved compositions plot at the end of a weak compositional trend formed by orthopyroxene–clinopyroxene pairs in summit-region xenoliths overall, and at the end of the compositional field for orthopyroxene–clinopyroxene in gabbro xenoliths elsewhere on Mauna Kea (Fig. 3d).
Plagioclase
Average An and Or contents for plagioclase in all xenoliths are indicated in Fig. 4a and, for reference, compared with the compositional field for plagioclase in Kilauea tholeiitic lavas and gabbros. Plagioclase in the tonalite compositionally straddles Ca-andesine and Na-labradorite (i.e. An55–45), whereas plagioclase in the gabbro inclusions within the tonalite is Na-labradorite (Fig. 4a). Both tonalite and its gabbro inclusions plot at the low-An end of a compositional continuum formed by plagioclase in P- and E-cone gabbros that begins at An85. Point-analyses for several grains in six E-cone gabbros, in the tonalite, and in its gabbro inclusions show the details of this continuum (Fig. 4b). The An–Or continuum is relatively Or rich (compared with Kilauea field), but overlaps some plagioclase observed for Mauna Kea lavas (Fig. 4b).
The diorites have plagioclase with average compositions that are oligoclase (Fig. 4a). They form a continuum extending from the Kilauea field, gradually increasing in Or to achieve Or15 at An15 (Fig. 4a). Point-analyses in Fig. 4c show that the compositional continuum among the diorite xenoliths extends beyond Or15. Plagioclase of the plagioclase–biotite vein (Fig. 2h) is largely oligoclase and it falls on the compositional continuum of the diorites (Fig. 4c).
The D-cone gabbro xenoliths that co-erupted with the tonalite and diorite xenoliths also fall on the Kilauea plagioclase trend (i.e. lower Or; Fig. 4a). Plagioclase in gabbro and wehrlite from two other summit-region cones, S and HK, is >An80 and therefore cannot be assigned to a particular An–Or trend (Fig. 4a).
Opaque oxides
The tonalite and its gabbro inclusions contain titaniferous magnetite (77 wt % FeO*) and titanohematite, an oxidization product of the former containing lower FeO*, 56–66 wt %, but higher TiO2, 30 wt %. The diorites have Fe-rich (>80 wt % FeO*), Al-poor, and Cr2O3-depleted (<0·05 wt %) titaniferous magnetites, and some also have ilmenite. Application of the Fe–Ti oxide geothermometer (e.g. Buddington & Lindsley, 1964) yields temperatures of 820–840°C at relatively oxidizing fO2 (10-12). Figure 5 shows the compositional continuum formed by the high-FeO*, Cr2O3-depleted Fe–Ti oxides in the diorites and the lower FeO*, Cr2O3-enriched oxides in the summit-region gabbro and wehrlite xenoliths.
Mica, amphibole, and apatite
Mica in the tonalite is phlogopite, with mg-number 79, and that in its gabbro lithic fragments has mg-number 80 (Fig. 6a). These mg-numbers are higher than those of coexisting clinopyroxenes, and equal to and higher than the mg-numbers of phlogopite observed in three summit-region gabbros, 75–79 (Fig. 6a). The phlogopite in the tonalite and its gabbro inclusions is primary, whereas that in the summit-region gabbros is secondary after clinopyroxene.
The diorite xenoliths and the plagioclase–biotite vein in gabbro xenolith D16 (Fig. 2h) contain biotite with evolved mg-numbers, 52–65 (Fig. 6). The most mafic biotite, mg-number 65, occurs sparsely in diorite D13 as a secondary phase after clinopyroxene, but the more evolved biotites in the other diorite xenoliths are primary grains. Compared with phlogopite in the tonalite, biotite in diorite has higher Al2O3 and TiO2 (Fig. 6b). Compared with rare primary biotite observed elsewhere on Hawaii, such as phenocrystal biotite in rhyodacite, the diorite biotite is more evolved and has lower Al2O3 and higher TiO2 (Fig. 6c). All mica analyzed in the tonalite, its gabbro inclusions, and in the diorites has only small amounts of Cl, 0·05–0·25 wt %, and F that is barely above detection limits.
The amphibole in the gabbro inclusions is hornblende. Its mg-number (76·2) is consistent with that of the coexisting clinopyroxene. Apatite in diorites (Table 3) has 3·6 wt % F and 1 wt % Cl, characteristic of apatite elsewhere in the Hawaiian system (e.g. Fodor & Vandermeyden, 1988; Spengler & Garcia, 1988).
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Glass compositions
Table 3 lists the compositions for brown glass that permeates diorite D9 as veinlets (Fig. 2g), and sparsely occupies interstices in diorite D3. Figure 7 shows the glass compositions in terms of SiO2 and K2O + Na2O, where veinlet glass is benmoreite composition and interstitial glass is phonolite. Overall, these glass compositions resemble those previously reported for some S-cone gabbros (Fodor & Vandermeyden, 1988).
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Whole-rock compositions
Whole-rock compositions for tonalite, its gabbro inclusions, diorite, and some summit-region gabbro and wehrlite xenoliths (Table 1) are illustrated in Figs 7 and 8. For comparision, Fig. 8 includes compositional fields for post-shield Hamakua (tholeiitic and alkalic) and Laupahoehoe (hawaiite, mugearite, benmoreite) lavas, and both Figs 7 and 8 include compositions of the glass in two diorites and the composition of Hawaii rhyodacite.
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The tonalite, with SiO2 66·9 wt % (dacite field in Fig. 7) and MgO 2·5 wt %, is more compositionally evolved than any Hamakua and most Laupahoehoe lavas, but its relatively low Al2O3, P2O5, Sr, and Nb abundances are not extensions of trends created by the Hamakua and Laupahoehoe lava compositions. Its SiO2, MgO, and P2O5 contents are like those of Hawaiian rhyodacite, but its TiO2, CaO, and Zr are higher and its Al2O3, K2O + Na2O, and Nb are lower. Perhaps most significant is that the tonalite REE pattern (Fig. 9a) shows a high La/Yb ratio and low total REE abundances relative to Mauna Kea basalts, and a positive Eu anomaly.
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In contrast, the gabbro within the tonalite has major- and trace-element abundances that plot in the compositional fields of Hamakua tholeiitic and alkalic basalts (Figs 7 and 8). Also, the gabbro REE pattern is like those of some Mauna Kea post-shield tholeiitic basalts (Fig. 9a). In spite of its ‘basaltic’ composition, the gabbro has a modal mineralogy that differs from that of typical Hamakua basalt—namely, phlogopite (
8 vol. %), abundant Fe–Ti oxide (15 vol. %), a trace of amphibole, and no olivine. The diorite xenoliths have SiO2 50–53 wt %. Three plot within the mugearite field and one is transitional to the hawaiite field in the SiO2 vs total alkalis diagram (Fig. 7). The overall compositions of the diorites resemble those of Laupahoehoe lavas, as they largely overlap the MgO-variation diagram fields for hawaiite, mugearite, and benmoreite (Fig. 8). Diorite D13, however, which is transitional to hawaiite in Fig. 7, has compositional aspects that are transitional between Hamakua and Laupahoehoe post-shield lavas, such as relatively high MgO (5 wt %), CaO, K2O, and Sr (Fig. 8). Also, P2O5 is higher and Zr lower in D13 than in any post-shield lavas. The REE patterns for all diorites are like those characteristic of hawaiite, mugearite, and benmoreite (Fig. 9b), and none has positive Eu anomalies despite containing >50 vol. % plagioclase.
The summit-region gabbroic xenoliths have whole-rock compositions unlike those of Mauna Kea lavas (Table 1; Fig. 8). Most notable are their low incompatible-element abundances, with some exceptions for plagioclase-rich xenoliths P1, P2, and P5 (e.g. Zr, Nb, P). Overall, abundances of plagioclase-compatible Al2O3, CaO, and Sr are equal to or higher than those in Hamakua lavas. Other diagnostic characteristics are the low total REE abundances relative to basaltic lavas and positive Eu anomalies (Fig. 9c). All of these whole-rock compositional features are characteristic of Mauna Kea gabbroic xenoliths in general (e.g. Fodor & Galar, 1997).
Isotope composition
The tonalite 87Sr/86Sr ratio is 0·703610 ± 10, and along with 143Nd/144Nd of 0·512976, 206Pb/204Pb of 18·58, 207Pb/204Pb of 15·49, and 208Pb/204Pb of 38·15, its isotopic composition overlaps the compositional fields established for both shield and post-shield Hamakua lavas of Mauna Kea (Kennedy et al., 1991; Yang et al., 1994). Only the Sr isotopic composition is available for the gabbro inclusions in the tonalite; it is 0·703589 ± 8, similar to that of the tonalite.
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DISCUSSION |
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Origin of tonalite and associated gabbro fragments
The tonalite has interstitial quartz poikilitically enclosing plagioclase (Fig. 2b) and a positive Eu anomaly (Fig. 9a), suggesting that this leucocratic xenolith has a cumulate origin. The isotopic compositions identify the tonalite as belonging to Mauna Kea shield or to Hamakua post-shield magmatism. In either case, the quartz, plagioclase of An55–45, clinopyroxene with mg-numbers
75, and the presence of phlogopite all identify the tonalite as derived from a liquid that was more evolved than nearly all basaltic magmas that erupted as shield and Hamakua post-shield lavas (e.g. Frey et al., 1990, 1991) or crystallized to gabbro and wehrlite cumulates (i.e. xenoliths of Fodor & Galar, 1997). The exception is a group of submarine shield lavas that Yang et al. (1999) noted to have clinopyroxene evolved to mg-number 71. Because the tonalite contains orthopyroxene, clinopyroxene with low Al2O3, and quartz, its parent magma was probably tholeiitic. The plagioclase An–Or continuum created by the tonalite and the E- and P-cone gabbros (Fig. 4a and b), some of which also have orthopyroxene, suggests a genetic relationship among these rock types: namely, the tonalite cumulate can represent an evolved liquid on the line of descent defined by the relatively mafic magma(s) that crystallized E- and P-cone gabbros.
A fractionation origin for the tonalite cannot be quantitatively modeled because it does not represent a liquid composition, but there is evidence that SiO2-rich liquids are produced during various stages of Hawaiian magmatism. For example, on a lava-flow scale, there is rhyodacite (66 wt % SiO2) in the Wainaea Range of Oahu (Bauer et al., 1973). Small-scale occurrences are more common. These are SiO2-rich (60–70 wt %) segregation veins (1 cm to 1 m) and oozes (fracture fillings) in lava lakes (e.g. Helz, 1980, 1987), and rhyolitic glass as interstitial material (e.g. Keil et al., 1972) and vesicle linings and fillings in basalt, some with micro-grains of pure SiO2 (e.g. Fodor et al., 1993).
Helz (1987) suggested on the basis of lava-lake drill cores that local SiO2-rich accumulations originate from filter pressing driven by gas pressures that increase during basalt magma differentiation (Anderson et al., 1984). Marsh (1996) presented a supportive model for silicic liquids segregating from basaltic magmas that he called solidification front instability (SFI). This SFI process refers to the sagging of a lining of mafic crystallization products from a reservoir roof and its ultimate collapse into the reservoir solidification front, followed by ‘pressing’ of evolved (silicic) interstitial and fracture-filling liquids from the front into overlying voids left by the collapse. The accumulation of the localized silicic liquids creates SiO2-rich lenses.
These models for segregation and accumulation of SiO2-rich liquids, each substantiated by field and thin-section evidence (Helz, 1987; Fodor et al., 1993; Marsh, 1996), can explain silicic liquid represented by the tonalite; that is, interstitial SiO2-rich liquid is inferred to have been filter pressed from a post-shield basaltic solidification front to coalesce into a lens. Adequate H2O vapor pressure, which is implied by abundant phlogopite in the tonalite–gabbro composite, probably facilitated the segregation process (Anderson et al., 1984). Whether or not the SFI process as detailed by Marsh (1996) was involved in that accumulation of the silicic liquid into a lens cannot be assessed from the xenolith alone. But to account for the tonalite’s cumulate texture, the segregated liquid must have occupied a small reservoir and crystallized in situ, or inward from a wall (e.g. Langmuir, 1989), largely precipitating plagioclase, Fe–Ti oxide, and mica, and trapping residual SiO2-oversaturated liquid.
Primary phlogopite is rare in the Hawaiian system and has been reported only for a Kahoolawe basalt that represents substantial differentiation (Fodor et al., 1998). The significance of phlogopite here is that it identifies hydrous magmatic conditions during the formation of the tonalite–gabbro composite, and hydrous conditions may be crucial for forming SiO2-rich liquids from basaltic magmas in volumes large enough to segregate (e.g. Kinzler et al., 2000). Specific to the tonalite example of quartz-saturated segregation on Hawaii, hydrous conditions were important because they: (1) allowed crystallization of phlogopite, a SiO2-undersaturated phase, to utilize K that would otherwise form K-feldspar, thereby leaving SiO2 to enrich in the residual liquid to form modal quartz; (2) created an oxidizing environment to encourage Fe–Ti oxide crystallization, which enriches the residual liquid in SiO2; (3) elevated vapor pressure to facilitate segregation of silicic liquid from a basaltic solidification front by gaseous filter-pressing (e.g. Anderson et al., 1984).
The tonalite illustrates that the Hawaiian magma system can produce appreciable modal quartz, and it accounts for the largest volume of unequivocally igneous quartz reported for Hawaiian magmatism. On a grander scale, its oceanic occurrence is fundamental to the concept that ancient oceanic environments yielded SiO2-rich magmas that led to nuclei for early continental crust formation. In particular, the tonalite supports the model for how small-scale, essentially routine production of interstitial silicic liquids (within basalt) can undergo accumulative processes to ultimately create bimodal oceanic environments. McCormick & Marsh (1997) believe that this process accounts for 0·1 km3 of rhyolite on Easter Island (71 wt % SiO2; Haase et al., 1997). The Galápagos Islands offer another rhyolite–basalt occurrence (Geist et al., 1995). The SiO2-rich (70–71 wt %) rock there is relatively voluminous (1 km3), and it exemplifies extreme differentiation from parental basalt. Iceland, too, is an example of bimodality, although the copious amounts of rhyolite there involved partial melting of altered mafic crust in addition to silicic segregations (e.g. Marsh, 1996; Gunnarsson et al., 1998).
Mineral compositions of the gabbro inclusions in the tonalite are slightly more primitive than those of the tonalite. This and the An–Or continuum (Fig. 4b) suggest that the gabbro is related to the tonalite. Hydrous and oxidizing conditions, indicated by phlogopite in the gabbro, can account for the high volume of Fe–Ti oxide (15%) in the gabbro inclusions. Because both oxide and mica crystallizations lead to SiO2 enrichment in residual liquids, the gabbro inclusions, then, probably represent the basaltic magma from which the tonalite parental liquid was ‘pressed’ (e.g. the gabbro inclusions are cognate). The veinlets of tonalite that penetrate the gabbro (Fig. 2a) are consistent with this cognate relationship. The rather fine, essentially equigranular texture of the gabbro (Fig. 2c) may also relate to H2O because saturation decreases crystal growth rates, although it also decreases nucleation rates (Fenn, 1977).
Origin of diorite
The four diorite compositions (Table 1; Figs 7–9) suggest that these rocks are coarse-grained equivalents of hawaiite and mugearite. Additionally, their andesine–oligoclase plagioclase compositions are consistent with plagioclase in the most evolved Hawaiian lavas—hawaiites, mugearites, benmoreites, and trachytes. For example, mugearite, benmoreite, and trachyte from Maui and Kohola have plagioclase in the range An40–15 (Keil et al., 1972; Spengler & Garcia, 1988). More mineral information for evolved Hawaiian magmas comes from the plagioclase–biotite vein dissecting gabbro D16 (Fig. 2h), which originated from the hawaiite–mugearite lava that entrained the xenolith. It has An15 plagioclase (Fig. 4c) and a biotite composition consistent with biotite in the diorites (Fig. 6; Table 2). These evolved mineral compositions are unrelated to the mineral compositions of the vein’s gabbro xenolith host, D16 (Fig. 2h), which are Fo80, An83, and mg-number 84 clinopyroxene (Table 2). I conclude, then, that holocrystalline equivalents of hawaiite and mugearite, namely diorites, have <An50 plagioclase zoned from andesine through oligoclase and biotite with mg-number 65–60.
The Mauna Kea post-shield hawaiitic substage, Laupahoehoe Volcanics, includes significant mugearite (e.g. Wolfe et al., 1997), and the bulk compositions of the diorite xenoliths suggest that these rocks (the diorites) are coarse-grained equivalents of Laupahoehoe lavas. Modeling by Frey et al. (1990) demonstrates that Laupahoehoe hawaiite liquid (and mugearite and benmoreite compositions derived therefrom; West et al., 1988) originated from magma represented by the earlier Hamakua basaltic substage, but with the notable distinction that Hamakua composition magmas produced hawaiite (Laupahoehoe) magmas at relatively high pressures. These regimes were below the shallow reservoir(s) occupied by the bulk of precursor Hamakua magmas, probably at the crust–mantle interface.
The change to higher-pressure crystallization during post-shield magmatism (i.e. from Hamakua to Laupahoehoe) reflects diminishing magma-production rates. It is manifested by compositional gaps in MgO diagrams (Fig. 8), but perhaps most obviously in normative Ne–Di–Ol diagrams that depict 1 bar and 8–30 kbar cotectics (Frey et al., 1990; Wolfe et al., 1997). Diorites D1, D3, and D10 have overall compositions like those of Laupahoehoe mugearites (Figs 7–9), and these coarse-grained equivalents accordingly plot closer to the high-pressure cotectic than to the 1 kbar cotectic, at the Ne end of the Ne–Di–Ol field for Laupahoehoe Volcanics (Fig. 10). Because of compositional similarity between Laupahoehoe Volcanics and these diorite xenoliths, the Frey et al. (1990) Hamakua fractionation model for generating Laupahoehoe hawaiite and mugearite at relatively high pressures also applies to the origins of D1, D3, and D10. As holocrystalline samplings, these diorites represent unerupted, degassed portions of magmas existing shortly after the change from Hamakua to Laupahoehoe post-shield substages (as identified by the lava compositions). They were dislodged and entrained by subsequently erupting Laupahoehoe magmas.
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Diorite D13 generally plots transitionally between the compositional fields for Hamakua and Laupahoehoe lavas in MgO variation diagrams (Fig. 8; e.g. K2O, CaO). Also, it plots closer to the low-P cotectic in the Ne–Di–Ol diagram than the other diorites, or in an ‘intermediate’ pressure region (Fig. 10). To test its compositional relationship with Hamakua magma, I used least-squares modeling for Hamakua alkalic basalt KI-2, the parent composition used by Frey et al. (1990) to show that Laupahoehoe hawaiites originated by segregation of clinopyroxene-rich gabbroic assemblages. Table 4 shows that KI-2 can also yield D13 by clinopyroxene-rich segregation. The overall crystallization assemblage totals 58% as olivine, plagioclase, clinopyroxene, magnetite, and ilmenite. Because Zr in D13 is unusually low (70 ppm; Table 1), zircon may have additionally segregated. Among the four diorite xenoliths analyzed for their whole-rock compositions, then, three conform to the Frey et al. (1990) model of subvolcanic high-P fractionation regime for Laupahoehoe magmas, and one (D13) suggests a more ‘intermediate’ P region of formation—but all are consistent with an origin by fractionation from Hamakua composition magmas.
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Polybaric crystallization for Mauna Kea post-shield magmas was also reported by Wolfe et al. (1997), but for a different compositional group: namely, Hamakua Fe–Ti lavas that have Al2O3/CaO ratios >1·6 have Ne–Di–Ol compositions that place their origins close to the crust–mantle interface, whereas those with Al2O3/CaO <1·6 evolved in shallower reservoirs, or at low pressures represented by Hamakua lavas as a whole (Fig. 10). Considering this example of polybaric origins for Hamakua Fe–Ti lavas (Wolfe et al., 1997) along with the intermediate- to high-P range indicated for the diorite xenoliths (Fig. 10), it appears that the change from ‘low’- to ‘high’-pressure regimes (i.e. the transition from Hamakua to Laupahoehoe) for Mauna Kea post-shield magma crystallization was a gradual process, probably coupled to gradually diminishing magma production rates.
The vesicular glass veinlets of benmoreite composition (Table 3) in diorite D9 (Fig. 2g) probably represent Laupahoehoe liquids that entrained and infiltrated diorite during eruptive ascent, cooling rapidly en route. The interstitial phonolitic glass in D3 (Table 3) probably represents liquid trapped during diorite crystallization.
Origin of summit-range gabbro and wehrlite
These mafic samples, by and large, are petrographically, mineralogically, and compositionally similar to most of the xenoliths discussed in detail by Fodor & Galar (1997). In brief, their textures, low incompatible-element abundances, and positive Eu anomalies (Figs 8 and 9) all point to cumulate origins for summit-range gabbro and wehrlite. Their mineral compositions and their entrainment by Laupahoehoe lavas and tephra are consistent with these xenoliths forming as cumulates from magmas similar to Mauna Kea post-shield Hamakua tholeiitic and alkalic basalts.
Paths and environments of compositional evolution: tonalite vs diorite
The diorites, with An <50 mol %, represent liquids differentiated from basaltic magmas in equilibration with An75 plagioclase (Fig. 4). Similarly, the tonalite and its gabbro inclusions, with An40–50, also have parentage in basaltic magma in equilibrium with An75. However, the An–Or relationship among the diorites suggests a lineage different from that for the tonalite. Specifically, Fig. 4a and c shows that diorites define an An–Or evolutionary trend extending from the compositional reference field for plagioclase in Kilauea lavas. [Plagioclase in the Laupahoehoe hawaiite host lavas for all Mauna Kea xenoliths also plots on this trend, as does plagioclase in mugearites and benmoreites from Maui and Kohala (Keil et al., 1972; Fodor & Vandermeyden, 1988; Spengler & Garcia, 1988).] The presence of orthopyroxene in some of the diorites and its absence in others suggest that the liquid compositions resulting from differentiation were largely ‘balanced’ between SiO2 over- and under-saturation. This is equivalent to the ‘saddle’ region along the alkali-feldspar join in the Q–Ne–Ks system (e.g. Morse, 1994),
In contrast, plagioclase in the tonalite, its gabbroic inclusions, and in gabbro xenoliths from summit-region cones E and P follows a relatively Or-enriched An–Or trend (Fig. 4a and b). Some of these samples also contain orthopyroxene, suggesting that if orthopyroxene is an indicator of SiO2 saturation, then such an Or-enriched plagioclase trend is not necessarily describing an alkalic, SiO2-undersaturation magmatic trend. Rather, the variations in An–Or trends may reflect particular conditions. For example, residual reservoir liquid may remain isolated to behave under ‘perfect’ fractional crystallization conditions, concentrating K2O to enrich the crystallizing plagioclase. Alternatively, residual liquid may mix with less-evolved reservoir liquids, as when pressed from cumulate layers (e.g. in situ crystallization; Langmuir, 1989), and have its K2O diluted. In the latter, K2O (Or) does not increase in products representing progressive crystallization as rapidly as under isolated fractional crystallization conditions. Accordingly, the plagioclase continuum (Fig. 4) with higher Or may represent an isolated, static, and rapidly shrinking reservoir, whereas the diorites may represent crystallization where interstitial liquids sustained ‘communication’ with a relatively large, dynamic reservoir of convecting and mixing liquid or with flowage related to dike emplacement (e.g. laminated texture of diorite D10; Fig. 2f).
Alternatively, a distinctively different explanation for the two Or trends among the summit-region xenoliths (Fig. 4a and b) is that crystallization pressures may have been influential. Experimental studies show that higher K2O in plagioclase correlates with higher pressure over the range 1–5 kbar (Fuhrman & Lindsley, 1988). In this case, tonalite would reflect a deeper environment than diorite. It may be unrealistic, however, to place the tonalite origin at an even higher pressure than the substantially high pressures indicated for the diorites (Fig. 10).
Summary
Leucocratic xenoliths from near the summit of Mauna Kea exemplify how lavas may not fully report magma differentiation histories by not representing the extreme mineral phases (e.g. quartz, oligoclase) and H2O conditions (e.g. phlogopite) that can be achieved by differentiation. In one example, a quartz-bearing tonalite xenolith suggests that basaltic magma can locally segregate liquids approaching granite compositions. The mechanics of such segregations and their accumulations as reservoirs have implications for the origins of ‘continental’ crust from oceanic magma. In another example, several diorite xenoliths are composed of largely andesine–oligoclase plagioclase with intergranular evolved clinopyroxene (mg-number 75–69). Although these diorite phase compositions are unlike any observed in Hamakua lavas, they probably represent the plutonic crystallization products of Hamakua magmas that fractionated at relatively high pressures—from deep within the volcano down to the crust–mantle interface. They largely represent the transition during post-shield development to Laupahoehoe (high-P) from Hamakua (low-P) magmatism.
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ACKNOWLEDGEMENTS |
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This project was supported in part by National Science Foundation grant EAR9104927. I also acknowledge Oregon State University Radiation Center for providing neutron activation services through their Department of Energy Reactor Sharing grant, and I thank P. Galar for assistance with some microprobe analyses. I owe the isotopic analyses to assistance from S. A. Goldberg, University of North Carolina isotope laboratory. F. A. Frey’s review comments were, as usual, invaluable.
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FOOTNOTES |
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*Telephone: 001-919-515-7177. Fax: 001-919-515-7802. E-mail: rfodor{at}.ncsu.edu
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REFERENCES |
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