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Journal of Petrology Advance Access originally published online on July 8, 2004
Journal of Petrology 2004 45(8):1663-1687; doi:10.1093/petrology/egh029
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Journal of Petrology 45(8) © Oxford University Press 2004; all rights reserved

Geochemical Constraints on the Role of Oceanic Lithosphere in Intra-Volcano Heterogeneity at West Maui, Hawaii

AMY M. GAFFNEY1,*, BRUCE K. NELSON1 and JANNE BLICHERT-TOFT2

1 DEPARTMENT OF EARTH AND SPACE SCIENCES, UNIVERSITY OF WASHINGTON, BOX 351310, SEATTLE, WA 98195, USA
2 LABORATOIRE DES SCIENCES DE LA TERRE, ÉCOLE NORMALE SUPÉRIEURE, 46 ALLÉE D'ITALIE, 69364 LYON, CEDEX 7, FRANCE

RECEIVED JUNE 1, 2003; ACCEPTED MARCH 5, 2004


   ABSTRACT
TOP
 ABSTRACT
INTRODUCTION
 SAMPLES AND ANALYTICAL...
 RESULTS
 DISCUSSION
 CONCLUSIONS
 APPENDIX
 REFERENCES
 
Stratigraphically well-constrained sequences of late shield-building stage lavas from West Maui volcano, Hawaii, show age-dependent compositional variability distinct from that seen in shield-stage lavas from any other Hawaiian volcano. These distinctions are defined by 206Pb/204Pb–208Pb/204Pb variation as well as 87Sr/86Sr correlation with 206Pb/204Pb and trace element compositions. The West Maui lavas from stratigraphically higher in the sequence have major and trace element and Sr–Pb–Hf–Nd isotopic compositions similar to Kea-type lavas sampled at the younger Mauna Kea and Kilauea volcanoes, indicating that the Kea compositional end-member of Hawaiian lavas has remained homogeneous over ~1·5 Myr. The 87Sr/86Sr–206Pb/204Pb variation in the stratigraphically lowest lavas is orthogonal to the all-Hawaii variation, indicating that it is not the result of mixing between components normally sampled by Hawaiian shield-stage magmas. We compare our West Maui data with observed compositions of Pacific oceanic basaltic and gabbroic crust, and predicted compositions for 2 Ga basaltic and gabbroic oceanic crust. The observed fine-scale compositional variability in the stratigraphically lowest West Maui lavas is consistent with 10–15% mixing of small-degree (2%) partial melts of the Pacific gabbroic oceanic crust with plume-derived, Kea-type magmas. Mass balance considerations indicate that this geochemical signal can be present in no more than 5% of the total volume of lavas erupted at West Maui.

KEY WORDS: Hawaii; Kea component; oceanic lithosphere; Sr–Pb–Hf–Nd isotopes


   INTRODUCTION
TOP
 ABSTRACT
 INTRODUCTION
SAMPLES AND ANALYTICAL...
 RESULTS
 DISCUSSION
 CONCLUSIONS
 APPENDIX
 REFERENCES
 
Investigations of the role of oceanic lithosphere in the geochemistry of Hawaiian magma sources has generated two main families of hypotheses. One invokes ancient oceanic lithosphere that has been subducted, stored in the mantle for some length of time and recycled into the plume source where it contributes to compositional variation in the plume-generated lavas (e.g. Hauri, 1996Go; Lassiter & Hauri, 1998Go; Blichert-Toft et al., 1999Go). Alternatively, plume-generated magmas may acquire the geochemical characteristics of oceanic lithosphere by interacting with or assimilating it as they rise to the surface (e.g. Chen & Frey, 1985Go; Eiler et al., 1996Go). Singly or together, these processes may contribute to the range of compositional variability expressed through the two primary compositional end-members of Hawaiian shield-stage magmatism: the relatively enriched Koolau component and the relatively depleted Kea component.

Kea-type lavas show the most depleted Sr and Nd isotopic compositions of all Hawaiian shield-stage lavas, and thus the Kea source material has been interpreted as having an association with, or derivation from, ambient depleted mantle or oceanic lithosphere. To date, the Kea component has been well described in late shield-stage lavas at only one volcano, Mauna Kea, where it has been thoroughly characterized through multi-disciplinary work on the Hawaii Scientific Drilling Project (HSDP) cores (e.g. Stolper et al., 1996Go, and included papers; Blichert-Toft & Albarède, 1999Go; Abouchami et al., 2000Go; DePaolo et al., 2001Go; Blichert-Toft et al., 2003Go; Eisele et al, 2003Go; Huang & Frey, 2003Go). From this characterization, broad conclusions have been drawn about the compositional structure of the Hawaiian plume and the generation of the compositional variations within the plume feeding Mauna Kea. Currently active Kilauea also has Kea-type compositions, and its historical summit lavas show rapid compositional changes (Pietruszka & Garcia, 1999Go). Kilauea lavas provide a point of contrast with West Maui and Mauna Kea, as they sample a window of time earlier in the shield-building stage than that sampled at West Maui or Mauna Kea. Loihi, Mauna Loa and Koolau are not dominated by Kea-type compositions as are Mauna Kea, Kilauea and West Maui; however, these volcanoes do show, over specific time intervals, time-dependent shifts towards Kea-type compositions (Staudigel et al., 1984Go; Garcia et al., 1993Go, 1995Go; Kurz et al., 1995Go; Tanaka et al., 2002Go). To date, the best characterizations of the Kea component over a large time window (i.e. over a significant stratigraphic section) are from the relatively young volcanoes Mauna Kea and Kilauea; there is so far no detailed characterization of the Kea component sampled over a comparably large time window at an older volcano.

Identification and characterization of oceanic lithosphere associations or other fine-scale compositional variations within Kea-type lavas require high-density sampling as well as stratigraphic (i.e. time) control. West Maui volcano has erupted Kea-type lavas throughout its sampled history, but there are few published geochemical data for this volcano (Macdonald & Katsura, 1964Go; Diller, 1982Go; Stille et al., 1986Go; Tatsumoto et al., 1987Go). The goals of this study are thus to (1) characterize geochemically the compositional range of lavas sampled by a Kea-type volcano at ~1·5 Ma, (2) compare characteristics of the Kea component as sampled at West Maui with that sampled at the younger volcanoes Mauna Kea and Kilauea, as well as Mauna Loa, Koolau and Loihi, (3) evaluate the oceanic lithosphere hypotheses as the probable cause of fine-scale compositional variations observed in West Maui lavas, and (4) determine the relationship, if any, of this variation to the Kea plume source.


   SAMPLES AND ANALYTICAL PROCEDURES
TOP
 ABSTRACT
 INTRODUCTION
 SAMPLES AND ANALYTICAL...
RESULTS
 DISCUSSION
 CONCLUSIONS
 APPENDIX
 REFERENCES
 
West Maui is one of six volcanoes that form Maui Nui, a bathymetric high comprising coalescing volcanoes (Haleakala, Kahoolawe, Lanai, West Maui, East and West Molokai) with a total area similar to that of the Big Island of Hawaii (Fig. 1). The K–Ar ages for shield-building activity in the Maui Nui volcanoes range from 1·84 ± 0·07 Ma at West Molokai to 0·83 ± 0·17 Ma at Haleakala (Naughton et al., 1980Go). Most exposed lavas at West Maui volcano are from the shield-building stage, and in places are capped by a thin veneer of lavas transitional between the shield-stage tholeiitic (Wailuku Basalts) and post-shield alkalic (Honolua Volcanics) basalts (Stearns & Macdonald, 1942Go; Diller, 1982Go). There are several post-shield stage alkalic flows, and four mapped occurrences of rejuvenated stage lavas (Lahaina Volcanics). Published K–Ar ages are few and range from 1·27 to 1·97 ± 0·96 Ma for the shield stage lavas (including transitional lavas), 1·15 to 1·50 ± 0·13 Ma for post-shield lavas, and 0·385 ± 0·012, 0·388 ± 0·009, 0·584 ± 0·010 and 0·610 ± 0·012 Ma for the rejuvenated stage lavas (McDougall, 1964Go; Naughton et al., 1980Go; Tagami et al., 2003Go).



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  		Fig. 1. Location of Papalaua Gulch and Mahinahina Well stratigraphic sections, as well as ‘deep’ and ‘shallow’ samples. Section depths are given in meters. Caldera location from Stearns & Macdonald (1942)Go.

 
The sample set for this study includes two stratigraphically controlled sequences of lavas, the Papalaua Gulch section (260 m thick) and the Mahinahina Well section (345 m thick), that sample late shield-building stage West Maui lavas (Fig. 1). The shorter Papalaua Gulch section is more proximal to the caldera margin and samples a smaller window of time than the more distal Mahinahina Well section. To characterize more completely the compositional range of variation of late shield-stage West Maui lavas, we also collected samples from a wide geographical distribution around the volcano. Based upon sample location, distance from the caldera rim and sample depth within the volcanic edifice, we have categorized all samples as either ‘deep’ or ‘shallow’. These categories refer to the depths of samples within the volcanic edifice, relative to one another. As these categories also have geochemical significance, we will use this terminology in the subsequent geochemical discussion. Figure 1 shows sampling locations for both deep and shallow samples; details of the sample categorization procedure are given in the Appendix.

Samples are dark to light gray, fine-grained basalts. Olivine phenocrysts are commonly present and a few samples also contain plagioclase phenocrysts or microphenocrysts. Clinopyroxene phenocrysts are only very rarely present. Some samples contain iddingsite-rimmed olivine phenocrysts. The samples from the Mahinahina Well cuttings show a higher degree of oxidation and Fe-staining than the surface-collected samples. Lithologically, the splits of cuttings are homogeneous, indicating little mixing between layers during drilling. For splits that showed some heterogeneity, we used only pieces that matched the dominant lithology of the split. For all isotopic analyses, we hand-picked 2–5 mm sample chips that were free of visible alteration.

Major element (X-ray fluorescence; XRF) and trace element (inductively coupled plasma mass spectrometry; ICP-MS) analyses were completed at the Washington State University GeoAnalytical Laboratory, according to procedures described by Johnson et al. (1999)Go (Table 1). The 2{sigma} analytical precision, based upon repeat analyses of the BCR-P standard and consistent with repeat analyses of samples and our UWBCR-1 internal standard, is as follows: SiO2, 0·11%; Al2O3, 0·29%; TiO2, 0·79%; FeOt, 0·31%; MnO, 1·09%; CaO, 0·57%; MgO, 2·8%; K2O, 0·91%; Na2O, 3·0%; P2O5, 0·53%; Ni, 11·8%; Cr, 7·2%; Sc, 15·4%; Ga, 8·7%; Zn, 3·2%; La, 3·7%; Ce, 0·2%; Pr, 1·9%; Nd, 3·5%; Sm, 4·3%; Eu, 4·7%; Gd, 2·4%; Tb, 1·7%; Dy, 2·8%; Ho, 2·8%; Er, 3·0%; Tm, 3·6%; Yb, 1·8%; Lu, 3·8%; Ba, 3·9%; Th, 4·9%; Nb, 4·4%; Y, 1·5%; Hf, 3·0%; Ta, 4·9%; U, 7·3%; Pb, 6·4%; Rb, 2·9%; Cs, 6·3%.


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  		Table 1: Major and trace element compositions

 
Pb, Sr and Nd isotope analyses were all made from the same initial sample dissolution. Prior to dissolution, we leached sample powders for 45 min in 6N HCl, using ~1 ml HCl per 100 mg sample. During leaching, we placed samples on a ~100°C hotplate for 20 min, in an ultrasonic bath for 10 min, on the hotplate again for 10 min and in the ultrasonic bath for an additional 5 min. We rinsed samples in sub-boiling distilled H2O prior to dissolution. We completed all sample preparation for Sr, Nd and Pb isotope analysis at the University of Washington (UW), Seattle, according to procedures described by Nelson (1995)Go (Table 2). We prepared Hf separations both at the École Normale Supérieure in Lyon (ENSL), France, and at UW following the procedures described by Blichert-Toft et al. (1997)Go. Thermal ionization mass spectrometry (TIMS) isotope analyses of Sr and Nd were carried out on the VG Sector TIMS at UW. Multi-collector ICP-MS isotope analyses of Pb and Hf were carried out using the VG Plasma 54 at ENSL (Blichert-Toft et al., 1997Go; White et al., 2000Go). Strontium NBS 987 standards run over the course of the West Maui analyses (n = 33 over 2 years) have an average 87Sr/86Sr = 0·710294 ± 0·000019 (2{sigma}). The average composition of Nd La Jolla standards run during this same time period (n = 25) is 143Nd/144Nd = 0·511845 ± 0·000020 (2{sigma}). This precision for the standard is similar to that obtained for multiple aliquots (n = 10) and analyses of our internal UWBCR-1 basalt powder (143Nd/144Nd = 0·512617 ± 0·000024, 2{sigma}). The 2{sigma} external reproducibility for Pb isotopic compositions is determined from repeat analyses of the 98B internal standard (n = 28 over two analytical sessions) and is 0·043%, 0·048% and 0·056% for 206Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb, respectively. The JMC 475 Hf standard gave 176Hf/177Hf = 0·282160 ± 0·000010 (2{sigma}) and was measured between every second or third sample to verify machine performance. Precision for 176Hf/177Hf is determined from replicate aliquots (n = 10) and analyses of our UWBCR-1 internal standard powder (176Hf/177Hf = 0·282865 ± 0·000016). Although our standard powder has been well homogenized, the quoted precision represents a maximum value, given that it incorporates all procedural sources of error and any sample heterogeneity that may exist.


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  		Table 2: Isotopic compositions

 

   RESULTS
TOP
 ABSTRACT
 INTRODUCTION
 SAMPLES AND ANALYTICAL...
 RESULTS
DISCUSSION
 CONCLUSIONS
 APPENDIX
 REFERENCES
 
Most samples have loss on ignition (LOI) <0·4% (Table 1), indicating minimal alteration. For most samples (21 of 27) K2O/P2O5 > 1, likewise indicating that the samples have not undergone significant post-cooling alteration, which would mobilize K relative to P. All but one of the samples with K2O/P2O5 < 1 are shallow samples. The samples that exhibit K loss also show loss of the mobile incompatible elements U and Rb relative to immobile incompatible elements such as Nb, Th and Zr. Cerium, which is mobile when oxidized to Ce4+ during alteration, is well correlated with immobile incompatible elements in all samples with one exception (LT22), indicating that alteration has affected only some elements in some samples. Of the samples that exhibit K, U or Rb loss, none has anomalous isotopic or major element compositions, with the exception of two samples with anomalously low SiO2 (LT27 and WA13). None of the indicators of element mobility (K2O/P2O5, K, U, Rb) correlates with 87Sr/86Sr or 206Pb/204Pb isotopic compositions that are potentially susceptible to alteration. Therefore, we include all samples in subsequent geochemical discussion.

Major elements
The West Maui shield-building stage samples are all tholeiitic basalts with 45–51 wt % SiO2, and 6–21 wt % MgO (Fig. 2). We also sampled three transitional-stage lavas, identified through mapping and description by Diller (1982)Go, and one each of post-shield alkalic and rejuvenated stage lavas. Data for these samples are included in Tables 1 and 2, but not in any figures or discussion. The major element variability for the deep and shallow groups shows only partially overlapping fields. The systematics of the sample suite as a whole correspond, for the most part, to expected differentiation trends, and are analogous to well-documented differentiation trends at Kilauea (Garcia et al., 2003Go). The two shallow West Maui samples (LT27 and WA13) with anomalously low SiO2 are also two of the most altered samples (K2O/P2O5 = 0·72 and 0·47, respectively). It appears that, in general, the shallow samples tend either to be more evolved than the deep samples, or to have undergone a higher degree of olivine accumulation (e.g. MgO > 15 wt %). The major element compositions of these lavas are broadly similar to tholeiites of the extensively studied Kea-type volcanoes Mauna Kea (Yang et al., 1994Go; Rhodes, 1996Go) and Kilauea (Chen et al., 1996Go; Garcia et al., 1996Go, 2000Go). Notably, the bend in Al2O3, CaO, CaO/Al2O3, and TiO2, vs MgO variation for West Maui (Fig. 2b, c, e and f) occurs at similar whole-rock compositions in the Kilauea lavas, suggesting common differentiation processes.



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  		Fig. 2. MgO variation diagrams for West Maui shield-stage lavas in ‘deep’ and ‘shallow’ sample categories. Oxides given in wt %. The 2{sigma} external reproducibility is smaller than symbol size for all oxides except for MgO, for which 2{sigma} error bar is shown. Also shown for comparison are compositions of Mauna Kea and Kilauea shield-stage lavas. Data sources: Garcia et al. (1992Go, 1996Go, 2000Go); Yang et al. (1994)Go; Chen et al. (1996)Go; Rhodes (1996)Go.

 
Trace elements
To account for variable amounts of olivine accumulation or fractionation, we adjusted all sample compositions to 13 wt % MgO either by incrementally adding equilibrium olivine or by removing Fo87 olivine until the sample composition reached 13 wt % MgO, and adjusted incompatible trace element compositions accordingly (assuming ). Because no olivine composition data exist for these samples, we chose Fo87 olivine as an equilibrium composition for our assumed parental magma. Although primitive Hawaiian magmas with MgO up to 16 wt % have been documented at other volcanoes (Haleakala, Chen, 1993Go; Kilauea, Clague et al., 1995Go), we have no samples in our West Maui collection in the 14–16 wt % MgO range, and consequently we have chosen not to call upon a primitive composition that we did not observe. Samples with MgO <7·0 wt % have probably begun to fractionate clinopyroxene as well as olivine (Helz & Wright, 1992Go; Montierth et al., 1995Go), and therefore we exclude these samples from plots and discussion involving corrected trace elements.

Both the deep and shallow groups of samples show a similar absolute range in the rare earth element (REE) slope La/Yb, although the deep group generally has lower [La] (Fig. 3a–d). Both groups also show the same positive correlation between La/Yb and [La]. This REE compositional variation is similar to, but encompasses a narrower range than, the variation observed in Mauna Kea tholeiites (Albarède, 1996Go; Feigenson et al., 1996Go; Yang et al., 1996Go; Huang & Frey, 2003Go). Overall REE patterns are distinct between the deep and shallow groups of West Maui lavas. The samples within the shallow group have similar overall REE patterns, but the REE concentrations vary among the samples (Fig. 3a). This is a common pattern of intra-volcano variation for Hawaiian tholeiites (e.g. Tilling et al., 1987Go; Frey et al., 1991Go). These observations contrast with those for the deep samples as a group, which show less variability in H(heavy)REE compared with L(light)REE concentrations (Fig. 3c); this variation pattern is similar to that observed in the historical Kilauea summit lavas (Pietruszka & Garcia, 1999Go). The shallow West Maui samples generally have a steeper HREE slope (higher Sm/Yb) than the deep samples, for a similar range in [Sm]. As a group, the deep samples have lower average REE concentrations than the shallow samples.



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  		Fig. 3. Rare earth element (REE) and incompatible trace element compositions for deep (circles) and shallow (triangles) West Maui lavas. Trace element concentrations are corrected to 13 wt % MgO. REE concentrations are chondrite-normalized using the values of Sun & McDonough (1989)Go. Error bars represent 2{sigma} external reproducibility; percent variation = 100 x (standard deviation/average).

 
Variation in ratios of incompatible trace elements, particularly HFSE (high field strength element)/HFSE and LILE (large ion lithophile element)/HFSE, such as, respectively, Nb/Zr and La/Nb, have only partially overlapping ranges for the deep and shallow groups (Fig. 3g and h). This indicates variation in either the source compositions or the range of source melting sampled by the two groups. The deep samples extend to lower [Nb] and Nb/Zr and to higher Sr/Nb than do the shallow samples (Fig. 3).

Uncorrected concentrations of the compatible trace elements Ni and Cr are well correlated with MgO, and the Ni, Cr–MgO correlations define continuous trends that encompass both the deep and shallow groups of samples (Fig. 4). Scandium (uncorrected)–MgO variation defines two distinct groups: a strong correlation that extends to high MgO and a weaker correlation that includes the lower-MgO samples and extends to the highest [Sc]. We take the first group to be the result of olivine accumulation. The poorly defined bend at ~7·5 wt % MgO in the second group may reflect the onset of clinopyroxene crystallization. In Hawaiian tholeiites, initiation of clinopyroxene crystallization at 7–7·5 wt % MgO has been observed experimentally as well as in nature (Helz & Wright, 1992Go; Montierth et al., 1995Go).



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  		Fig. 4. MgO-compatible trace element variation for West Maui shield-stage lavas. Trace element compositions are not corrected for olivine control. MgO composition is in wt %, trace element composition is in ppm, and error bars represent 2{sigma} external reproducibility.

 
Sr, Nd, Hf and Pb isotopes
Comparison of the 87Sr/86Sr–143Nd/144Nd–176Hf/177Hf variation in the West Maui samples with Hawaiian compositions as a whole shows that these lavas, along with lavas from the Kea-dominated volcano Mauna Kea, and in agreement with the few previously published West Maui data (Stille et al., 1986Go; Tatsumoto et al., 1987Go), have some of the most depleted isotopic compositions of all the Hawaiian islands (Fig. 5). In Nd–Hf–Sr isotope space, there is little overlap between the fields defined by the deep and shallow groups. The deep group has the more depleted compositions, and in 87Sr/86Sr–143Nd/144Nd space is displaced slightly towards less radiogenic 87Sr/86Sr, and in the direction of Pacific mid-ocean ridge basalt (MORB), relative to typical Kea-type compositions (as defined by Mauna Kea and Kilauea). There is small overlap between the deep West Maui lavas and Mauna Kea and Kilauea (Fig. 5a). The 143Nd/144Nd–176Hf/177Hf variation for these samples shows a similar distinction between the deep and shallow groups, with the deep samples being more depleted than the shallow samples (Fig. 5b). There is also much more scatter in the compositions of the deep samples—they show nearly twice the variability in 143Nd/144Nd as do the shallow samples. This contrasts with the major and trace element compositions of the deep samples, which tend to be less variable than those of the shallow samples. The West Maui samples as a group fall slightly below the ocean island basalt (OIB) array, as has been observed for Mauna Kea but not Kilauea (Blichert-Toft et al., 1999Go). The West Maui and Mauna Kea samples fall farther below the OIB array than any other measured Hawaiian shield-stage lava.



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  		Fig. 5. (a) 87Sr/86Sr–143Nd/144Nd and (b) 143Nd/144Nd–176Hf/177Hf variation in the West Maui lavas. Shown for comparison are the fields for shield-stage lavas from Koolau, Mauna Loa, Loihi, Mauna Kea and Kilauea. Insets show detail of compositional variation for West Maui lavas only. OIB and Hawaii correlation lines from Blichert-Toft et al. (1999)Go. Error bar represents 2{sigma} external reproducibility. Data sources: Stille et al. (1986)Go; Garcia et al. (1993Go, 1995Go); Roden et al. (1994)Go; Kurz et al. (1995)Go; Chen et al. (1996)Go; Lassiter et al. (1996)Go; Blichert-Toft & Albarède (1999)Go; Blichert-Toft et al. (1999)Go; Pietruszka & Garcia (1999)Go; Tanaka et al. (2002)Go.

 
The sample suite as a whole shows variation in 207Pb/204Pb of only 4–5 times analytical error, and there is no distinction in the 206Pb/204Pb–207Pb/204Pb trend between the deep and shallow groups (Fig. 6a). 208Pb/204Pb is strongly correlated with 206Pb/204Pb in the set of samples as a whole, and the deep vs shallow samples show separate correlations (Fig. 6b). The shallow and deep samples were analyzed at ENSL in random order and over the same periods of time. Thus, the contrast in Pb isotope ratios between the two sample groups is not due to a shift in instrument performance between multiple analytical sessions. The shallow samples have a slightly lower degree of correlation than the deep samples, and also show slightly lower 208Pb/204Pb for the same range of 206Pb/204Pb exhibited by the deep samples. High-precision Pb isotope studies of Mauna Kea HSDP-1 lavas (680 m thick; Abouchami et al., 2000Go) also identify two compositional groups, although these groups appear to lack any correlation with stratigraphy. Interestingly, the two groups identified in Mauna Kea lavas are distinct in both 206Pb/204Pb–207Pb/204Pb and 206Pb/204Pb–208Pb/204Pb space, whereas the two West Maui groups show no distinction in 206Pb/204Pb–207Pb/204Pb space. Eisele et al. (2003)Go used high-precision Pb isotope compositions, specifically 206Pb/204Pb–208Pb/204Pb variation, to define three distinct compositional groups in the HSDP-2 lavas (2790 m thick). All of our West Maui samples, as well as the HSDP-1 Mauna Kea samples, correspond to the HSDP-2 array with the lowest 208Pb/204Pb.



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  		Fig. 6. (a) 206Pb/204Pb–207Pb/204Pb and (b) 206Pb/204Pb–208Pb/204Pb variation in West Maui lavas. Error bar represents 2{sigma} external reproducibility.

 
The 206Pb/204Pb–87Sr/86Sr variation in the West Maui lavas shows a distinction between the deep and shallow lavas (Fig. 7). The deeper samples show a positive correlation that trends to less radiogenic 206Pb/204Pb and 87Sr/86Sr than shallow West Maui or Mauna Kea. Discussion of this trend is given in the following section. This contrasts with the 208Pb/204Pb–87Sr/86Sr variation (not shown), where the shallow samples display a better correlation than the deep samples.



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  		Fig. 7. 87Sr/86Sr–206Pb/204Pb variation for West Maui lavas. Shown for reference are postulated compositions of the ‘Kea’ and ‘Koolau’ end-members (Stille et al., 1986Go; West et al., 1987Go; Eiler et al., 1996Go).

 
The 176Hf/177Hf and 206Pb/204Pb stratigraphy for the Papalaua Gulch and Mahinahina Well sections shows minimal correlation between the two isotope systems, or of either isotope ratio with depth (Fig. 8). The average 206Pb/204Pb compositions and standard deviations for the two sections are nearly identical. The average 176Hf/177Hf of the deeper section (0·28312) is slightly more radiogenic than that of the shallower section (0·28311), and the range of values in these two sections show the same standard deviation. Although the average 176Hf/177Hf values of these two sections are within error of each other, they correspond in the expected sense to different average 143Nd/144Nd compositions (0·51299 and 0·51301 for the deep and shallow sections, respectively), and therefore we believe that the 176Hf/177Hf compositional difference between the two sections is real. Blichert-Toft et al. (1999)Go described the all-Hawaii 176Hf/177Hf–206Pb/204Pb variation as a mixing hyperbola between the Kea end-member with relatively radiogenic 176Hf/177Hf and 206Pb/204Pb and the Koolau end-member with relatively unradiogenic 176Hf/177Hf and 206Pb/204Pb. Not unexpectedly, our West Maui data fall on the upper limb of this hyperbola, near Mauna Kea.



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  		Fig. 8. 176Hf/177Hf and 206Pb/204Pb isotope stratigraphy for the Papalaua Gulch and Mahinahina Well stratigraphic sections. Also shown are the average compositions for each section, ±1 SD.

 

   DISCUSSION
TOP
 ABSTRACT
 INTRODUCTION
 SAMPLES AND ANALYTICAL...
 RESULTS
 DISCUSSION
CONCLUSIONS
 APPENDIX
 REFERENCES
 
At West Maui, we observe fine-scale geochemical variation defined by 206Pb/204Pb–87Sr/86Sr, 206Pb/204Pb–208Pb/204Pb and 87Sr/86Sr–trace element compositions. These correlations are present at West Maui only in lavas that erupted over a particular time interval. Studies of Pb isotope composition have identified fine-scale Pb isotope variability within other Kea-dominated volcanoes (Kilauea, Pietruszka & Garcia, 1999Go; Mauna Kea, Abouchami et al., 2000Go; Eisele et al., 2003Go), but the particular 87Sr/86Sr-correlated trends we identify at West Maui have not been observed in tholeiites of any other Hawaiian volcano. The 206Pb/204Pb–87Sr/86Sr trend in the deep West Maui lavas is oblique to the trend of pure mixing between Kea and Koolau components as they are typically sampled by Hawaiian shield-stage magmas. It is not uncommon to observe Hawaiian shield-stage lava compositions that deviate from the pure Kea–Koolau mixing trend, and in fact it is the rule rather than the exception that at least three end-member compositions are required to describe compositional variations within one volcano or across the whole chain. However, the specific signal that we have identified in the 87Sr/86Sr–206Pb/204Pb–trace element compositions is unique to this volcano both in its compositional range and in its stratigraphic dependence.

There exists a long history of associating Hawaiian magma source regions with oceanic lithosphere (e.g. Hofmann & White, 1982Go; Chen & Frey, 1985Go; Eiler et al., 1996Go; Hauri, 1996Go; Lassiter & Hauri, 1998Go), and several studies associate oceanic lithosphere with both the Kea and Koolau end-members. The Sr–Nd–Pb–O isotopic compositions of Kea-type magmas have supported hypotheses for the origin of the Kea component in small-degree melts of ambient depleted MORB-source mantle (Chen & Frey, 1985Go) or hydrothermally altered Pacific crust (Eiler et al., 1996Go) that mixes with plume-derived magmas. Blichert-Toft et al. (1999)Go used Pb–Hf isotopic correlations to argue against Pacific lithosphere as the source of the Kea component, and Lassiter & Hauri (1998)Go used O–Os isotopic correlations to argue that the Kea source is the gabbro + harzburgite/dunite segment of ancient (>1 Ga) oceanic crust that has been subducted and recycled into the Hawaiian plume. Koolau-type magmas have been interpreted, on the basis of Hf–Pb isotopic correlation (Blichert-Toft et al., 1999Go) and O–Os compositions (Lassiter & Hauri, 1998Go), to have their source in recycled basaltic oceanic crust ± pelagic sediment in the Hawaiian plume.

Although mixing between the Kea and Koolau end-members serves as a general framework to explain inter-volcano isotopic and elemental variability among Hawaiian volcanoes on the scale of the entire archipelago, it does not account for the fine-scale variability we observe in West Maui volcano. In the following, we examine the origin of the fine-scale, intra-volcano chemical trends in West Maui, and evaluate whether we can relate these variations to the hypotheses for involvement of oceanic crust discussed previously. Specifically, we compare the expected compositional variation for ambient vs ancient upper (hydrothermally altered basaltic) and lower (gabbroic) oceanic crust, and model the contributions of these compositional reservoirs to Kea-type lavas. In doing this, we evaluate whether the source of the fine-scale, intra-volcano West Maui variation is inherent to the Kea source or instead represents a distinct process or source that operates or exists independently of the origin or identity of the Kea source. Finally, we seek to identify similar trends in other volcanoes, and determine whether processes that we recognize at West Maui operate at other Hawaiian volcanoes as well.

Sr–Pb–Hf isotope variation
The 206Pb/204Pb–87Sr/86Sr variation observed in the shallow West Maui lavas is similar to that observed in shield-stage lavas in other Kea-dominated volcanoes (e.g. Anders & Nelson, 1996Go; Lassiter et al., 1996Go; Pietruszka & Garcia, 1999Go; Abouchami et al., 2000Go). However, the deep lavas are unique among Kea-type volcanoes in that they define a positive 206Pb/204Pb–87Sr/86Sr trend that extends to less radiogenic 87Sr/86Sr than observed in shield-stage lavas from any other Kea-type volcano (Fig. 9a). As mentioned above, this variation is oblique to the overall all-Hawaii shield-stage trend, and hence cannot be the result of mixing between Hawaiian plume components as they are represented at other volcanoes.



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  		Fig. 9. (a) 87Sr/86Sr–206Pb/204Pb and (b) 206Pb/204Pb–176Hf/177Hf variation in West Maui lavas, with tholeiitic lava compositions from Koolau, Mauna Loa [(a) only], Loihi, Mauna Kea, Kilauea and Hawaiian rejuvenated-stage lavas [(a) only] from Kauai, Koolau and Haleakala shown for comparison. Oceanic crustal end-member compositions are given in Table 3 and discussed in text. Mixing line between West Maui and Pacific oceanic crust gabbro represents mixing between 2% melt of oceanic gabbro and plume-derived Kea-type magma; tick marks indicate 10% mixing increments. Mixing line between 2 Ga gabbro and 2 Ga upper oceanic crust (UOC) is for solid–solid mixing between these end-members; tick marks along this line indicate 5% increments of mixing. Data sources for other volcanoes: Stille et al. (1983)Go; West & Leeman (1987)Go; Reiners & Nelson (1998)Go; Abouchami et al. (2000)Go; Lassiter et al. (2000)Go; Blichert-Toft et al. (2003)Go; as well as those listed in Fig. 5 caption.

 
The two possibilities that we explore are contributions of very small-degree partial melts of ambient Pacific lithosphere (~110 Ma) or within-plume melts of ancient (~2 Ga) oceanic crust that was hydrothermally altered at a mid-ocean ridge, subducted, stored in the mantle, and recycled into the Hawaiian plume. On a global scale, the details of subduction zone processes and their effect on the chemical budget of the subducting slab are poorly quantified, and potentially can have highly variable effects (Stracke et al., 2003Go). For this reason, we have chosen to make no assumptions about subduction zone processing and instead focus on the first-order chemical variations we would expect in aged oceanic crust, based on the range of available observations and datasets from the literature.

Different lithological layers in the oceanic lithosphere develop distinct radiogenic isotope signals over time, and thus we specifically consider two different segments of oceanic crust: the hydrothermally altered basaltic upper crust and the gabbroic cumulates in the lower crust. These, in combination with other lithologies (pelagic sediment + basalt; gabbro + depleted lithosphere), have been used to explain the genesis of Kea and Koolau compositional end-members (Lassiter & Hauri, 1998Go; Blichert-Toft et al., 1999Go). In our models, we investigate whether altered basalt and gabbro, independently or combined, can explain the fine-scale compositional variation distinct from Kea–Koolau mixing that we observe within just one Kea-type volcano.

For the 110 Ma Pacific upper oceanic crust end-member, we use 87Sr/86Sr and 206Pb/204Pb measured for upper oceanic crust from Ocean Drilling Program (ODP) Site 843 (King et al., 1993Go), located 300 km from Maui. For the 110 Ma gabbroic end-member, we use 87Sr/86Sr and 206Pb/204Pb of gabbroic xenoliths from Hualalai volcano, interpreted to be Pacific oceanic crust, measured by Lassiter & Hauri (1998)Go. We use [Sr] and [Pb] measured in oceanic crust drill cores (Staudigel et al., 1995Go; Hart et al., 1999Go). There currently exist no 176Hf/177Hf measurements for oceanic crust of the appropriate age, so we calculated a possible range of compositions based upon the 176Hf/177Hf composition of modern East Pacific Rise MORB (Chauvel & Blichert-Toft, 2001Go). We calculated the 176Hf/177Hf composition of this depleted mantle MORB-source reservoir at 110 Ma, and then forward-modeled the 176Hf/177Hf composition of the basaltic and gabbroic portions of oceanic crust generated from this depleted mantle reservoir, based upon Lu/Hf measured in oceanic crust drill-cores (Staudigel et al., 1995Go; Hart et al., 1999Go). For the 2 Ga end-members, we calculated the 87Sr/86Sr, 206Pb/204Pb and 176Hf/177Hf of the depleted mantle MORB-source reservoir at 2 Ga, again based upon modern compositions of East Pacific Ridge MORB (Chauvel & Blichert-Toft, 2001Go), and then forward-modeled the composition of basaltic and gabbroic crust derived from this reservoir at 2 Ga, in the same way as we did for the 176Hf/177Hf composition of the 110 Ma gabbro crust. For calculating the elemental concentrations in the 2% gabbro melt, we use a gabbro mode of 20% orthopyroxene, 20% clinopyroxene, 20% plagioclase and 40% olivine, and partition coefficients from McKenzie & O'Nions (1991)Go and Hauri & Hart (1995)Go. These results are presented in Table 3.


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  		Table 3: Model results

 
Of the modeled 2 Ga source components, hydrothermally altered upper oceanic crust has radiogenic 87Sr/86Sr and 206Pb/204Pb compositions, whereas gabbroic crust evolves to unradiogenic 87Sr/86Sr and 206Pb/204Pb compositions (Fig. 9a). Neither of these components alone lies on a trajectory of mixing with the shallow lavas that could result in the observed variability in the deep samples. However, mixing between a melt from an approximately 25:75 upper oceanic crust:gabbro solid–solid mixture and plume-derived Kea-type magmas would reproduce the range of compositions in the deep lavas.

Alternatively, we consider the composition of ambient Pacific oceanic crust. Isotopic compositions of upper oceanic crust from ODP Site 843 (King et al., 1993Go), located 300 km from Maui, have unradiogenic 87Sr/86Sr, but 206Pb/204Pb too radiogenic to cause the variability observed in the deep West Maui samples (Fig. 9a). Some gabbro xenoliths from Hualalai volcano, which have been interpreted to represent the oceanic crust beneath Hawaii (Clague & Chen, 1986Go), have unradiogenic 87Sr/86Sr and 206Pb/204Pb and Sr/Pb such that a mixing hyperbola between small-degree melts (2%) of these gabbros and average shallow West Maui lavas describes the compositional range in the deep West Maui lavas (Okano & Tatsumoto, 1996Go; Lassiter & Hauri, 1998Go). Mixing of small-degree melts of this ocean crust layer with plume-generated Kea-like magmas would result in hybrid magmas similar to what we observe at West Maui. The Sr/Pb as well as [Sr] and [Pb] of the gabbro melt depend upon the fraction of melt generated (F). Higher-F melts (>2–3%) will have lower [Sr] and [Pb] and therefore a larger proportion of this melt would be required to mix with the plume-generated magmas to produce the observed trends. A smaller-degree melt would also provide the observed mixing trend, but such a small-degree melt may not have sufficient melt connectivity to separate from the solid residue.

Within the range of 176Hf/177Hf variability at West Maui, the deep samples tend to have only slightly higher than average compositions (Fig. 9b). Thus the source of this variation in the deep samples must be close enough in composition to the shallow lavas so that the extent of mixing required by the 87Sr/86Sr compositions of the deep samples does not result in dramatic shifts in 176Hf/177Hf. Whatever caused the isotopic shift in the deep samples must have a 176Hf/177Hf composition similar to or higher than ‘Kea’, rather than lower. Magma generated through mantle melting (i.e. MORB) will have Lu/Hf lower than its mantle source, such that {varepsilon}Hf (as defined in the footnote of Table 2) of the melt will decrease relative to the residual source as it ages. The 110 Ma oceanic crust that we model has 176Hf/177Hf higher than Kea-type lavas, but progressively older model crust will have 176Hf/177Hf less radiogenic than the Kea-type lavas. Thus, both of the modeled 2 Ga end-members, or any mixture of them, will have 176Hf/177Hf significantly lower than Kea-type lavas. However, as with the 206Pb/204Pb–87Sr/86Sr variability, the 206Pb/204Pb–176Hf/177Hf variation in the deep samples is consistent with modeled contributions of small-degree melts from ambient Pacific gabbros, with the same mixing proportions as identified in the 87Sr/86Sr–206Pb/204Pb variation (Fig. 9b).

87Sr/86Sr and trace element constraints on end-member compositions
The deep West Maui lavas show correlations of 87Sr/86Sr with some trace elements (Fig. 10). Among the deep samples, those with less radiogenic 87Sr/86Sr trend to higher La/Yb, Nb/Zr, Nb/Hf, Nb/La, [Nb] and [Sr]. Although these elemental indicators are dependent in part on degree of melting, the correlation with 87Sr/86Sr indicates that the variations they describe are fundamentally related to the source. The high degree of correlation for these parameters also implies that the melt end-members involved are very homogeneous, and therefore there is only a small range in extent of melting of this end-member source as it contributes to the magmas. If this source melted over a wider range of F as it contributed to the West Maui lavas, then the melt end-member would have a wider range in elemental composition for a given isotopic composition, and there would be much more scatter in the West Maui trends. With 87Sr/86Sr vs Nb/Zr and Nb/Hf, the deep samples define a linear trend that is distinct from the trend defined by the shallow samples. Although the absolute spread in Nb/Zr of the deep samples is close to the analytical error for this ratio, the difference in this ratio between the two sample groups is greater than the analytical uncertainty. It appears that three end-members (at least) are required to explain the observed trends. Whereas one end-member has high 87Sr/86Sr and two end-members low 87Sr/86Sr, they are distinct in trace element composition. In these mixing scenarios, the deep samples form well-correlated arrays implying that, at least for the chemical signals under consideration, the contributing end-members are homogeneous. It therefore follows that the larger range in compositions observed for the shallow compared with the deep samples stems from heterogeneity in the low-87Sr/86Sr end-member that contributes to the array defined by the shallow samples.



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  		Fig. 10. Trace element–87Sr/86Sr variation diagrams. All trace element concentrations are normalized to 13 wt % MgO. Error bar represents 2{sigma} external reproducibility. Shown for comparison are fields for tholeiitic lavas from Mauna Kea, Kilauea and Loihi. Data sources: Garcia et al. (1993Go, 1995Go); Albarède (1996)Go; Chen et al. (1996Go); Hofmann & Jochum (1996)Go; Lassiter et al. (1996)Go; Pietruszka & Garcia (1999)Go.

 
For La/Yb, Nb/Zr, Nb/Hf and Nb/La vs 87Sr/86Sr, the shallow West Maui samples fall within the range defined by Kilauea, Mauna Kea and Loihi (Garcia et al., 1993Go, 1995Go; Albarède, 1996Go; Chen et al., 1996Go; Hofmann & Jochum, 1996Go; Lassiter et al., 1996Go; Pietruszka & Garcia, 1999Go). For these trace element ratios, Kilauea historical and prehistoric lavas together define a broader range and extend to higher values than West Maui, Mauna Kea or Loihi. Only for Nb/La–87Sr/86Sr do the Kilauea historical lavas define a trend that is distinct from prehistoric Kilauea lavas. The deep West Maui lavas fall either on the edge of or outside the fields defined by Mauna Kea, Kilauea and Loihi. None of these volcanoes shows La/Yb, Nb/Zr, Nb/Hf or Nb/La vs 87Sr/86Sr trends as clearly defined as the deep West Maui lavas.

Placed in the context of the preceding discussion, we can consider whether any of the proposed isotopic reservoirs can also produce a unique trace element signature. Gabbros from the Gabal Gerf Ophiolite complex (Zimmer et al., 1995Go), as well as ODP Site 735B (Hart et al., 1999Go), show an extremely wide range in Nb/La (0·19–1·5), Nb/Zr (0·03–0·14), La/Yb (1·24–4·28) and Nb/Hf (0·5–5·22). Because in all of these ratios the numerator is more incompatible than the denominator, the ratios will be greater in the melt relative to its source. The absolute increase of the ratio in the melt relative to the source depends upon F, partition coefficients and the mineralogical mode of the source. Within the range of variability introduced by these parameters (see discussion below), it is possible to calculate small-degree melts of oceanic gabbro that have trace element ratios appropriate for end-members in the mixing scheme proposed here.

Partial melting of a gabbro can also potentially generate magmas with Eu and Sr anomalies. Gabbro from oceanic crust drill cores as well as ophiolites (Zimmer et al., 1995Go; Hart et al., 1999Go) show both positive and absent Eu and Sr anomalies. As shown in Fig. 11, the predicted presence or absence of a Eu anomaly is dependent upon the starting gabbro composition and partition coefficients. We show REE compositions of two representative gabbros (Zimmer et al., 1995Go), and 2% melts derived from these using two different sets of partition coefficients (Schnetzler & Philpotts, 1970Go; McKenzie & O'Nions, 1991Go; Hauri & Hart, 1995Go). La/Yb in these melts ranges from 0·5 to 7·9. The range of resultant melts shows positive, negative or absent Eu anomalies, and therefore it is difficult to use absence or presence of a Eu anomaly in the West Maui lavas as an argument either for or against mixing with small-degree gabbro melts. The West Maui lavas do not show any Eu anomaly, and both deep and shallow groups of samples show approximately the same range of [Sr]. However, the low-87Sr/86Sr end-member that contributes to the deep samples also has higher [Sr] than the high-87Sr/86Sr end-member contributing to the deep samples. This is consistent with additions of small-degree gabbro melts to the deep West Maui lavas.



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  		Fig. 11. Range of possible REE concentrations for 2% melts of oceanic gabbro. ‘gabbro A’ and ‘gabbro B’ represent the typical range of ophiolite gabbro compositions (Zimmer et al., 1995Go; see text for discussion). Melts shown are for 2% melting of ‘gabbro A’ and ‘gabbro B’ using partition coefficients from Hauri & Hart (1995)Go for olivine, orthopyroxene and clinopyroxene (HH), and Schnetzler & Philpotts (1970)Go for plagioclase (SP), or from McKenzie & O'Nions (1991)Go for all minerals (MO).

 
Perhaps the primary uncertainty in modeling isotopic and elemental compositions of the oceanic crust end-members is the effect of variation in choice of partition coefficients. Wherever possible, we avoid introducing this uncertainty by using elemental concentrations as are actually observed in rocks (e.g. ODP cores, xenoliths, young MORB); however, this is not always possible, and in Fig. 11 we illustrate the range of effects of two different sets of partition coefficients on modeling small-degree gabbro melts. The effect of this uncertainty is especially enhanced at the low degrees of partial melting (2%) that we are considering here. The second largest variability in end-member choice or calculation is the variation inherent to the composition of the rocks themselves. In some instances (e.g. 87Sr/86Sr and 206Pb/204Pb of 110 Ma gabbro and upper oceanic crust) we are able to use data from actual rocks that we believe represent the end-member compositions. In these cases, we use an average of multiple analyses for the end-member composition. In other cases, where modeling of the composition of aged components (e.g. 176Hf/177Hf) amplifies variation introduced by starting composition, we show a range of possibilities for the modeled end-members (Fig. 9, Table 3). In Fig. 11, we also show melts generated from two different gabbro starting compositions. It is evident that the combination of gabbro starting composition and choice of distribution coefficient can result in a wide range of melt compositions. Although not all compositions provide results that are consistent with our hypothesis and data, we do find that there are geologically reasonable combinations of partition coefficient and gabbro composition that are consistent with our hypothesis.

Age-correlated compositional shifts in other Hawaiian volcanoes
Several other Hawaiian volcanoes show age-correlated compositional shifts in the isotopic and trace element composition of shield-stage, tholeiitic lavas (e.g. Kurz et al., 1995Go; Pietruszka & Garcia, 1999Go; Tanaka et al., 2002Go). In Koolau and Mauna Loa, volcanoes known to be dominated primarily by the Koolau component, these shifts are evident as Sr–Nd–Pb isotopic compositions that lie on a trajectory towards Kea-type compositions, as defined by Mauna Kea and Kilauea (Kurz et al., 1995Go; Tanaka et al., 2002Go). A subset of Koolau lavas overlaps with Kilauea and Mauna Kea fields (Fig. 9a), whereas none of the Mauna Loa lavas show compositional shifts of this magnitude. In both of these volcanoes, lavas showing the Kea-like compositions are the oldest ones sampled for each volcano: from the submarine Nuuanu slide block for Koolau, and from the submarine SW rift zone, Ninole and Kahuku Basalts of Mauna Loa.

Kilauea, which is dominated by Kea-type compositions in its historical lavas, has prehistoric lavas (Hilina) that extend to the most radiogenic 206Pb/204Pb observed in any Hawaiian shield-stage lavas, but do not extend to significantly lower 87Sr/86Sr. In the ‘Kea-like’ lavas from Mauna Loa and Koolau, as well as in the entire range of lavas from Mauna Kea and Kilauea, the compositional variation is expressed with a wide range for 206Pb/204Pb (~18·2–18·9), which corresponds to very limited variation in 87Sr/86Sr (~0·70358–0·70367). In contrast, the compositional shifts we see in the deep West Maui lavas are specifically defined by lower 87Sr/86Sr for a given range in 206Pb/204Pb. Although we see short-term compositional changes in other Hawaiian volcanoes that indicate a complexity in source components and the time periods over which they contribute to the magmas, none of the absolute senses of variation observed in Kilauea, Mauna Loa or Koolau is consistent with interaction of magmas with oceanic crust, using the end-members that we define for the oceanic crust beneath West Maui.

Tholeiites from Loihi, which is still in its early stages of growth, have Sr–Nd–Pb isotopic compositions comparable with Mauna Kea and Kilauea (Staudigel et al., 1984Go; Garcia et al., 1993Go). A subset of Loihi tholeiitic lavas shows a positive 206Pb/204Pb–87Sr/86Sr correlation similar to what we observe in West Maui. Although the Loihi compositions do not actually overlap with West Maui compositions, and do not extend to 87Sr/86Sr as low as those we observe in West Maui deep lavas, these lavas may still reflect a process similar to what we propose for West Maui. Primarily, the lower magma fluxes at the early and late stages of shield growth allow for the preservation of the geochemical signals of interaction between plume-generated magmas and the oceanic crust.

As this low-87Sr/86Sr end-member is sampled at West Maui, and possibly at Loihi, it does not constitute an identifiable end-member in Kea-type lavas as they are sampled at Mauna Kea or Kilauea, or in the variability in the compositional range of Hawaiian volcanoes as a whole. It therefore does not appear that this compositional end-member is a fundamental component of the Hawaiian plume, which otherwise would be sampled in a more regular or identifiable manner. Rather, it may reflect transitory processes related to early (Loihi) and late (West Maui) stages of Hawaiian volcano growth. The existence of this end-member at West Maui is identifiable only through dense sampling and stratigraphic control, and therefore lack of observation of this signal at other volcanoes does not necessarily preclude its existence.

Rejuvenated stage lavas, which erupt after a period of quiescence following the post-shield stages of volcano growth, occur on Kauai (Koloa Volcanics), Koolau (Hololulu Volcanics), East Molokai (Kalaupapa Volcanics), West Maui (Lahaina Volcanics) and Haleakala (Hana Volcanics). These lavas have more depleted 87Sr/86Sr, 143Nd/144Nd and 176Hf/177Hf than observed in any Hawaiian shield-stage lavas, and are commonly interpreted as having a source external to the Hawaiian plume (e.g. Clague & Frey, 1982Go; Reiners & Nelson, 1998Go). Although they are compositionally distinct from the local Pacific oceanic crust, both the asthenosphere and Pacific lithosphere beneath Hawaii are possible sources for the rejuvenated lavas (Lassiter et al., 2000Go). The rejuvenated lavas have less radiogenic 206Pb/204Pb and 87Sr/86Sr relative to the Kea end-member, as we also observe in the deep West Maui lavas, but they are offset from the trajectory suggested by the 206Pb/204Pb–87Sr/86Sr variation in the deep West Maui lavas, and have 206Pb/204Pb that is too unradiogenic to share a source with the deep West Maui lavas.

Constraints on contributions of Pacific oceanic lithosphere to West Maui lavas
The model that is most consistent with our observations from West Maui is that compositional variations measured in the deep samples are the result of mixing small-degree melts of Pacific oceanic crust with the primary, plume-generated West Maui magmas. Trace element ratios and Sr concentrations in the deep samples point to a contribution from the gabbroic layer of oceanic crust, and Sr–Pb–Hf isotope variations suggest that this is the ~110 Ma Pacific crust, rather than ancient recycled oceanic crust within the plume itself.

The generation and incorporation of these melts have implications for the heat budget and mass balance of the overall system. We used the MELTS algorithm (Ghiorso & Sack, 1995Go) to model isenthalpic assimilation of oceanic crust gabbro into a primitive West Maui magma (sample WA17). For all model runs, the pressure is 2 kbar, starting temperature of the assimilating magma is 1325°C, and starting temperature of the assimilant is either 100 or 200°C. We modeled assimilation of three starting compositions with olivine:orthopyroxene: clinopyroxene:plagioclase proportions of 40:20:20:20, 30:0:40:30 and 20:20:20:40, to understand the influence of the initial mineralogical mode of the assimilant. For each model run we report r = mass assimilated/mass crystallized, and in Fig. 12 we show this parameter as a function of total mass assimilated. In all models, for the first 16 wt % assimilation (relative to initial mass of the assimilating magma) olivine is the only crystallizing phase. After 16 wt % assimilation, clinopyroxene and eventually plagioclase join the crystallizing assemblage. However, the isotope and trace element data presented in the previous sections indicate involvement of only small-degree (~2%) melts of oceanic gabbro, so we limit the following discussion exclusively to a system with low degrees of assimilation. Within the range of assimilant composition and starting temperature that we explored in these models, r ranges from 0·73 to 0·9 for small degrees of assimilation, which translates into a maximum decrease in the total volume of liquid in the system of 10% over the range of assimilation required to result in the geochemical shifts observed in the deep lavas. Therefore, the assimilation model supported by the geochemical data is also thermally plausible.



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  		Fig. 12. Results of MELTS (Ghiorso & Sack, 1995Go) models of assimilation of gabbro into primitive shallow West Maui magma. Different runs are labeled by mode of olivine:orthopyroxene:clinopyroxene:plagioclase and starting temperature of assimilant. (See text for discussion of parameters.)

 
The major element changes that accompany assimilation are slight increases in SiO2, Al2O3 and CaO, and decreases in MgO and FeO. These shifts are similar to those that accompany olivine fractionation without any assimilation of gabbro, making it difficult to use major element composition alone as an indicator of assimilation of oceanic crust. Ten to 15% assimilation of a small-degree gabbro melt into a primitive West Maui magma (~13 wt % MgO) will produce a resultant melt with ~8·5–10 wt % MgO. Of the deep West Maui lavas, only one sample has MgO >10 wt %. We do not observe any significant differences between the deep and shallow groups with respect to major element variability as a function of 87Sr/86Sr. However, we would not necessarily expect to observe significant variation, as the major element changes with assimilation are small, and it is difficult to distinguish them from olivine control.

As the starting composition in these models we used the most primitive (MgO = 13·61 wt %) shallow West Maui composition that did not show significant olivine accumulation (i.e. high olivine abundance in hand sample). In other Hawaiian volcanoes, primary magmas with MgO ≥15 wt % have been identified (Chen, 1993Go; Clague et al., 1995Go), and a hotter, more primitive starting composition may produce a different result in the assimilation calculations. However, we have not observed any primitive magmas such as these at West Maui, and although they may exist, we prefer to use the observed primitive composition, to keep our model consistent with our observations at this volcano.

Generation of these geochemical signals requires extraction of small degrees (2%) of melt from the lower oceanic crust. Assuming that the West Maui total volume is 6·4 x 103 km3 (Bargar & Jackson, 1974Go), the presence of this oceanic crust-derived geochemical signal throughout the volcano would require 640–1280 km3 of small-degree melt of oceanic crust, which would in turn imply depleting nearly 50 000 km3 of oceanic crust of a small degree of melt. When comparing this with the volume of gabbro underlying the entire volcano, ~3000 km3 (volcano radius 20 km; gabbro thickness 2·5 km), it is apparent that no more than 5% of the total volume of erupted lavas can exhibit this geochemical signal, which derives from 10–15% mixing of low-degree gabbro partial melt with plume-derived magmas. It is reasonable to envisage that this signal would be present during times when the magma pathways through the oceanic crust reconfigure in response to the motion of the Pacific plate over the Hawaiian hotspot. Then, as new magma pathways become established and conduit walls armored or reacted, the imprint of the gabbro geochemical signal decreases. The lavas we sampled represent the latest stages of volcano growth, which could imply that the oceanic crust signal may be a response to a decreasing magma flux from the plume. West Maui is also the least voluminous of all the major Hawaiian volcanoes (Bargar & Jackson, 1974Go), which may explain why the geochemical signals of this process are preserved at this particular volcano, whereas at other, larger volcanoes, where the geochemical fingerprint is swamped or diluted by the relatively more substantial volumes of magma moving through the oceanic crust, they are missing. Also, because few tholeiitic shield-stage volcanoes have been sampled in this much stratigraphic detail, this process may be as yet unrecognized in other shields.

Is there still a requirement for 2 Ga recycled oceanic crust in Kea-type lavas?
Because none of the trace element signals that we attribute to Pacific oceanic lithosphere are dependent on the age of the source, it is possible that they could derive from an ancient subducted and recycled oceanic crust gabbro layer, as has been previously suggested for Hawaii and other ocean island volcano systems (Lassiter & Hauri, 1998Go; Chauvel & Hémond, 2000Go; Geldmacher & Hoernle, 2000Go; Sobolev et al., 2000Go). So far, we have not considered the possible role of 2 Ga lithospheric mantle (harzburgite/lherzolite) that is residual to the process of MORB generation. We calculated the isotopic composition of 2 Ga residual lithosphere that would be complementary to the 2 Ga gabbro and basalt compositions calculated above. For parent/daughter ratios in the depleted lithosphere, we used the residue of a 20% batch melt of depleted mantle, using partition coefficients from McKenzie & O'Nions (1991)Go and Hauri & Hart (1995)Go, initial element concentrations from Sun & McDonough (1989)Go, and a starting mineralogical composition of 52% olivine, 30% orthopyroxene, and 18% clinopyroxene (Hirschmann & Stolper, 1996Go). Variation in mineralogical mode of the depleted residue (e.g. clinopyroxene content) has only a minor effect relative to the magnitude of the modeled isotopic compositions, and therefore does not influence our conclusions. This reservoir will develop extreme isotopic compositions (176Hf/177Hf = 0·2836, 87Sr/86Sr = 0·7025, 206Pb/204Pb = 14·22) as will the gabbroic and basaltic parts of the oceanic crust (modeling and results discussed previously). However, such depleted lithosphere will sensibly affect bulk isotopic composition only when present in 90–95% abundances. A hybrid source of 90–95% 2 Ga depleted lithosphere and 5–10% of a 2 Ga 25:75 basalt:gabbro mixture will produce a source that could give rise to the Sr–Pb–Hf isotopic variations seen in the deep West Maui samples. This source would have relatively depleted major element compositions, and, therefore, we would expect magmas derived from a source dominated by residual mantle to show major element-isotope variability distinct from the typical Kea-type magmas. As this is not apparent, we argue against contributions of an ancient, hybrid, depleted lithosphere:basalt:gabbro source as the cause of the fine-scale variations we observe in the deep West Maui lavas.

The differences in isotopic composition among the modeled ancient depleted lithosphere, basalt and gabbro end-members are so great that minor variations in the relative mixing proportions could have significant effects on the bulk composition of the mixture. It is therefore conceivable that a mixture of depleted lithosphere: basalt:gabbro could act as the source of the ‘Kea’ plume component as defined by its Hf–Pb–Sr–Nd isotopic composition. However, Kea-type magmas that are sampled during the ~1·5 Myr time span over which Mauna Kea, Kilauea and West Maui erupted are very homogeneous in isotopic composition. If the Kea source is the result of mixing between recycled lithosphere, basalt, and gabbro components, then the mixing processes that generate the Kea source have remained extraordinarily reproducible over this time period. This implies consistent processes of oceanic crust generation and alteration at the mid-ocean ridge as well as sampling of the recycled oceanic lithosphere package in the plume.


   CONCLUSIONS
TOP
 ABSTRACT
 INTRODUCTION
 SAMPLES AND ANALYTICAL...
 RESULTS
 DISCUSSION
 CONCLUSIONS
APPENDIX
 REFERENCES
 
In the West Maui lavas, we see contributions from three components. Mixing between two of these components describes the variability inherent to Kea-type magmas. These two ‘sub-components’ are defined by their 87Sr/86Sr variability and a number of trace elements. They show only small distinctions in Hf, Pb or Nd isotopic compositions, and, therefore, it is likely that they are closely related genetically. The third component is not required to account for the compositional variability of Mauna Kea, Kilauea or the shallow lavas at West Maui. Our interpretation of the source of this component, and also its apparent absence in the Mauna Kea, Kilauea and shallow West Maui lavas, indicates that it is not intrinsic to Kea-type lavas or to the genesis of their plume source.

In the deep West Maui lavas, we have identified the chemical fingerprint of Pacific oceanic crust gabbro overprinted onto the plume-derived Kea-type magma compositions. These compositional variations, which we attribute to mixing with small-degree melts of oceanic gabbro, are present only over short time scales, and, based on mass balance calculations, can be present in no more than ~5% of the total volume of lavas erupted. These characteristic compositional trends have not been observed in shield-stage lavas at other Hawaiian volcanoes. Their absence at other volcanoes may be, in some cases, a consequence of less dense sampling or lack of stratigraphic control, or may indicate that this geochemical signal is not generally characteristic of Kea-type volcanoes or of the plume source that generates Kea-type magmas.

The process that we observe in the West Maui deep lavas, namely the overprinting with signals of partial melts from gabbroic oceanic crust, is unlikely to be unique to this volcano. It is the fundamental homogeneity of Kea-type lavas that allows the identification of variation from plume-derived Kea-type compositions. The geochemical signal of this process is so subtle that it would be extremely difficult to resolve in the more heterogeneous Koolau-type volcanoes or even at more voluminous Kea-type volcanoes where this signal would be diluted by the higher magma flux.


   APPENDIX
TOP
 ABSTRACT
 INTRODUCTION
 SAMPLES AND ANALYTICAL...
 RESULTS
 DISCUSSION
 CONCLUSIONS
 APPENDIX
REFERENCES
 
Most of the ‘breadth’-reaching samples we collected either on ridges that closely represent original flow surfaces (Stearns & Macdonald, 1942Go) or at stream level in traverses of deep canyons that dissect the volcano. Taking into account distance from the caldera, depth within the volcanic edifice and local dips of flows, we have divided these samples into two groups: ‘deep’ and ‘shallow’. This terminology describes the collection depth of a given sample relative to the surface of the volcano. In most of the canyons that we traversed, lavas are dipping more steeply than the canyon floor, such that samples collected farther inland are older than samples collected closer to the coast. Therefore, within each canyon, we were able to assign relative ages to the samples. Correlating relative ages between canyon traverses is more difficult, because it requires making assumptions about the volcano morphology, eruption frequency and distance traveled by individual flows, relative to the caldera and rift zones as mapped by Stearns & Macdonald (1942)Go and Diller (1982)Go, and therefore is not strictly quantifiable in the absence of actual ages. In general, the same thickness of stratigraphic section represents more absolute time the farther away it is from the caldera. As a guideline, at 2 km away from the caldera, we use 300 m as the dividing depth between deep and shallow, and at 10 km from the caldera, we use 75 m as the dividing depth. Based on these guidelines, the ‘deep’ and ‘shallow’ categories show some geochemical distinction, and we use this terminology in geochemical discussion. Table A1 gives geographic coordinates, distance from caldera and depth within the volcanic edifice for all samples.


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  		Table A1: Sample locations

 

   ACKNOWLEDGEMENTS
 
We thank Glenn Bauer for providing the Mahinahina Well samples and logistical information, and David Sherrod for collaboration in the field. Jerry Hinn and Philippe Télouk always kept the mass spectrometers running smoothly. We thank J. Bryce, A. Pietruszka and H.-J. Yang for helpful and constructive reviews. This work was supported by a Geological Society of America Harold T. Stearns grant, a DOSECC Internship grant and a UW Graduate Research grant to A.M.G. J.B.T. acknowledges financial support from INSU.


   FOOTNOTES
 

* Corresponding author. Telephone: 206-685-3326. Fax: 206-543-0489. E-mail: agaffney{at}u.washington.edu


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 APPENDIX
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