He, Sr, Nd, and Pb isotopic constraints on the origin of the Marquesas and other linear volcanic chains

Chemical Geology 240 (2007) 205 – 221 www.elsevier.com/locate/chemgeo

He, Sr, Nd, and Pb isotopic constraints on the origin of the Marquesas and other linear volcanic chains P.R. Castillo ⁎, P. Scarsi 1 , H. Craig ✠ Scripps Institution of Oceanography, UCSD, La Jolla, CA 92093-0212, USA Received 12 January 2006; received in revised form 8 February 2007; accepted 13 February 2007 Editor: R.L. Rudnick

Abstract The classic hotspot hypothesis [Morgan, W. J., 1971. Convection plumes in the lower mantle. Nature 230, 42–43], which posits that linear volcanic chains are traces of fixed plumes in the mantle on moving lithospheric plates, was instrumental in elevating the plate tectonics paradigm in the 1960s into a modern Earth Science theory. The hypothesis itself, however, remains conjectural because many of its predictions, particularly the simple age-progressive type of volcanism, are not observed in many linear volcanic chains. As an alternative explanation, it is proposed that linear volcanic chains are formed through magmatism along preexisting lines of weakness such as transform zones and old sutures, or along cracks created by stresses on lithospheric plates. The Marquesas linear volcanic chain in south-central Pacific has geologic features that are consistent with some of the predictions of both hypotheses. To better constrain the origin of this volcanic chain, we collected major and trace element and Sr, Nd, Pb, and He isotopic data from several Marquesan lavas. Our new analyses combined with literature data classify the samples into the well established tholeiitic to mildly alkalic, low 87Sr/86Sr, high 143Nd/144Nd, shield-building volcanic phase lava group and highly alkalic, high 87Sr/86Sr, low 143Nd/144Nd, post-shield phase group. Lead isotopes show generally higher 206Pb/204Pb ratios and suggest evidence of crustal assimilation for the shield-building phase lavas, consistent with the argument that the shield-building phase volcanism has a lithospheric source component. On the other hand, post-shield phase lavas that are predicted to represent the true composition of the mantle source by the hotspot hypothesis have higher 3He/4He ratios and these are coupled to other geochemical tracers. Thus our results show that the Marquesas volcanic chain, similar to many other linear volcanic chains, has a high 3He/4He component in its mantle source. The presence of such a distinct source component cannot be easily explained by dispersed upper mantle heterogeneities, but provides a powerful constraint for the hotspot origin of many linear volcanic chains. © 2007 Elsevier B.V. All rights reserved. Keywords: Radiogenic isotopes; Helium isotopes; Marquesas; Hotspot; Mantle plume

1. Introduction ⁎ Corresponding author. Tel.: +1 858 534 0383; fax: +1 858 822 4945. E-mail address: [email protected] (P.R. Castillo). 1 Now at: Istituto di Astrofisica Spaziale e Fisica Cosmica, IASF-Pa I.N.A.F. Via Ugo La Malfa 153, 90146 Palermo, Italy. ✠ Deceased. 0009-2541/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2007.02.012

Volcanic chain lavas have a wide range of radiogenic isotope compositions (i.e., Sr, Nd, Pb, Hf, and Os isotopes) suggesting that they come from a compositionally heterogeneous mantle source. The contemporary hotspot hypothesis claims that the heterogeneity of this source

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results from recycling of old (up to ∼b.y. old) crustal materials such as altered oceanic crust, sediments and delaminated subcrustal lithosphere in the mantle and from Earth's planetary differentiation. Many geochemists (e.g., White, 1985; Zindler and Hart, 1986; Hart, 1988; Hofmann, 2003) envisage that linear volcanic chain lavas originate from upwelling geochemically enriched mantle plumes consisting of recycled crustal materials such as high μ (high 238U/204Pb ratio; HIMU) and enriched mantles 1 and 2 (EM1 and EM2, respectively) and of portions of the mantle that still have close to primitive or bulk Earth composition. They further envisage that these plume components reside in the lower section of the mantle, perhaps at the core–mantle boundary, and are overlain by an upper section consisting of the geochemically depleted, mid-ocean ridge basalt mantle (DMM) source. Although other investigators agree that the mantle source of linear volcanic chain lavas is indeed compositionally heterogeneous because of crustal recycling, the recycling events do not have to be necessarily ancient and the crustal materials do not have to form deep-seated plumes (e.g., Anderson, 1996; Foulger and Natland, 2003; Meibom and Anderson, 2004; Foulger and Anderson, 2005). The recycled

materials instead comprise the heterogeneities in a “perisphere” layer in the shallower part of the mantle that overlies the DMM (Anderson, 1996) or are small- to moderate-scale heterogeneities ubiquitously distributed in the upper mantle (Meibom and Anderson, 2004). Linear volcanic chain lavas are formed through sampling of such heterogeneities when and where the oceanic plate is breached by stress. Plume detractors further claim that plume believers adapt the hypothesis as a priori model in a variety of settings by the addition of posteriori model supplements. As a result, the plume hypothesis has been made to account for essentially any observation or the lack of expected observations at many linear volcanic chains (Foulger and Anderson, 2005). Thus the origin of linear volcanic chains is the object of recent healthy debates (e.g., Foulger and Natland, 2003; DePaolo and Manga, 2003; Sheth, 2003; Foulger et al., 2004). The Marquesas Archipelago, consisting of 8 main islands and several islets and seamounts on top of an elongated bathymetric swell, is the easternmost member of a group of Neogene–Quaternary linear volcanic chains in French Polynesia, south-central Pacific (Fig. 1). The swell is oriented sub-parallel to the current Pacific plate motion and the islands and seamounts show a general

Fig. 1. Generalized bathymetric map of the Marquesas Archipelago from (Liotard et al., 1986; Dupuy et al., 1987; Vidal et al., 1987; Caroff et al., 1999; Legendre et al., 2005a). Islands and major seamounts are shown in black; 3000 and 4000 m bathymetric contour lines are also shown. Samples analyzed in the study are from the islands of Nuku Hiva, Hiva Oa and Fatu Hiva. Inset shows the general location of the archipelago (square) in the Pacific Ocean basin.

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age-progressive volcanism from ∼1 Ma in the southeast to ∼6 Ma in the northwest (Duncan and McDougall, 1974; Duncan et al., 1986; Desonie et al., 1993). These have led some people to believe that it is a trace of a hotspot. However, the slower calculated rate of motion of Marquesan volcanism relative to the Pacific plate motion, absence of a volcanically active hotspot and presence of a crustal root below the central portion of the archipelago seem to indicate that the Marquesan lavas may be coming from the crust rather than from the deeper mantle (McNutt and Bonneville, 2000). Hence, it is equally possible that the Marquesas Archipelago is being built along a fissure created by tectonic stress (Gutscher et al., 1999; McNutt and Bonneville, 2000; Guille et al., 2002). In order to constrain the origin of the Marquesas volcanic chain, we have collected He, Sr, Nd, and Pb isotope and major and trace element data on olivinebearing lavas collected from the Marquesas Archipelago (Fig. 1). Using the combined major element, trace element and isotope data, we want to define more clearly the geochemical characteristics of the source of their parental magmas. Specifically, we want to determine the presence or absence of a distinct, high 3He/4He mantle source and how such a source fits in the volcanologic and geochemical evolution of the Marquesan volcanic chain. The presence of such a source, as is proposed for many other linear volcanic chains, would support the hotspot hypothesis although not necessarily the presence of a primitive or bulk Earth mantle (e.g., Class and Goldstein, 2005; Parman et al., 2005). In contrast, the absence of such a distinct source would support the (henceforth called) ‘plate stress’ hypothesis for the origin of linear volcanic chains. 2. Previous work Marquesan subaerial volcanic rocks are mainly mafic basalts and lesser amounts of differentiated rocks that belong to both tholeiitic and alkalic rock series. The lavas are compositionally heterogeneous and the heterogeneity is observed not only between different volcanoes, but also within individual volcanoes. They are mildly to highly enriched in incompatible trace elements and have a wide range of 87Sr/86Sr ratios despite a moderate range in both 143 Nd/144Nd and 206Pb/204Pb values (e.g., Liotard et al., 1986; Dupuy et al., 1987; Vidal et al., 1987; Desonie et al., 1991; Woodhead, 1992; Desonie et al., 1993; Caroff et al., 1999; Legendre et al., 2005a,b). Two major phases of volcanism are observed in the Marquesas and these are equated to the two major stages of volcanism that figure prominently in the evolution of volcanic islands — an initial shield-building phase and a

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later post-shield phase (e.g., Duncan and McDougall, 1974; Duncan et al., 1986; Woodhead, 1992; Desonie et al., 1993). As in many other islands, the shield-building phase is often separated from the post-shield phase by a hiatus. Moreover, the Marquesan shield-building phase generally belongs to the tholeiitic to mildly alkalic rock series and is compositionally more homogeneous and more mafic than the post-shield phase, which belongs to the alkalic rock series. Accordingly, the shield-building Marquesan lavas have lower highly incompatible/moderately incompatible trace element, lower 87Sr/86Sr (≤0.7041) and higher 143 Nd/144Nd (≥0.51285) ratios than the post-shield lavas (e.g., Woodhead, 1992; Desonie et al., 1993). The difference in the ranges of chemical and radiogenic isotope composition of the two phases of volcanism clearly indicates that they originate from different sources, but the details of the petrogenesis of Marquesan volcanic rocks are still the object of debate. Although there is the general question of whether the Marquesas was formed by plate stress or by hotspot volcanism, the majority of current petrogenetic arguments assume that the islands were formed by a hotspot and so the debate centers mainly on the nature of the plume source. On the one hand, some argue that Marquesan volcanoes are sampling a heterogeneous plume source containing a depleted component, which is either an intrinsic constituent of the plume or a part of the asthenospheric mantle assimilated by the plume during its ascent. During the shield-building phase of volcanism, large degree of melting or sampling at the center of the plume produces more homogeneous tholeiitic to mildly alkalic lavas whereas during the post-shield phase, small degree of melting or sampling at the periphery of the plume produces alkalic lavas (Duncan et al., 1986; Liotard et al., 1986; Dupuy et al., 1987; Vidal et al., 1987; Desonie et al., 1993; Caroff et al., 1999; Legendre et al., 2005a,b). Others, argue that the plume source is compositionally homogeneous, but the lower oceanic lithosphere is involved during the shield-building phase, and that the relatively ‘uncontaminated’ plume source is sampled during the post-shield phase (Duncan et al., 1986; Woodhead, 1992; Desonie et al., 1993). Interestingly, the few previously reported 3He/4He isotopic ratios of Marquesan lavas span only a narrow range that overlaps with the normal mid-ocean ridge basalt (MORB) value of 8 ± 1 relative to atmospheric 3He/4He ratio (RA). However, 3 He/4He ratios still show ∼inverse correlation with 87 Sr/86Sr (Desonie et al., 1991). 3. Samples and analytical methods Samples analyzed in this study were collected by the late H. Craig for the specific purpose of investigating the

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Table 1 Major and trace element and He, Sr, Nd, and Pb isotopic composition of Marquesan lavas Sample

FH-OM-1

FH-OM-3

FH-OM-4

FH-TOP

FH-HV-01

FH-HV-2

FH-HV-3

HO-PUA-2

Location

Fatu Hiva

Fatu Hiva

Fatu Hiva

Fatu Hiva

Fatu Hiva

Fatu Hiva

Fatu Hiva

Hiva Oa

Rock type a

m-alk bas

m-alk bas

thol bas

m-alk bas

mug

m-alk bas

thol bas

haw

SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 Total

46.06 3.58 11.86 12.37 0.15 10.85 10.01 2.12 0.75 0.20 97.95

47.03 3.32 11.88 12.75 0.14 11.92 8.55 2.24 0.89 0.11 98.84

42.51 1.77 5.84 14.72 0.16 25.46 5.85 0.83 0.37 0.13 97.62

46.60 3.56 13.04 12.29 0.15 9.00 8.17 2.48 1.22 0.32 96.83

52.28 2.45 15.00 8.79 0.12 3.76 6.30 3.34 2.72 0.43 95.20

46.06 3.27 11.16 12.35 0.15 11.47 8.87 2.22 1.11 0.35 97.01

43.85 3.33 10.59 13.59 0.15 14.26 9.13 1.67 0.27 0.21 97.04

47.56 3.37 14.26 11.49 0.15 6.90 8.21 2.96 2.34 0.58 97.82

Rb Ba Nb Sr Zr Y La Ce Nd Sm Eu Gd Dy Er Yb Lu Ba/Nb Ba/Zr Nb/Zr Rb/Sr

9 340 43 578 289 19 47.6 100.0 47.3 8.97 1.59 9.52 6.78 3.72 2.56 0.33 7.95 1.18 0.15 0.02

11 264 42 535 269 15 34.8 74.8 39.9 7.95 1.44 7.90 5.77 2.90 2.04 0.28 6.35 0.98 0.16 0.02

4 280 24 269 120 9 8.06 18.1 9.68 2.32 1.04 2.21 1.54 0.66 0.47 11.7 2.32 0.20 0.02

32 378 56 571 354 28 30.0 64.7 36.2 7.90 3.11 7.42 5.72 2.73 2.16 0.32 6.77 1.07 0.16 0.06

21 300 34 427 217 20 8.46 19.8 13.0 2.75 2.18 4.10 3.07 1.94 1.35 0.21 8.77 1.38 0.16 0.05

26 314 40 582 286 26 27.3 61.1 34.0 7.32 2.72 6.91 4.99 2.32 1.65 0.19 7.75 1.10 0.14 0.04

2 209 35 412 239 16 10.1 24.7 14.4 3.70 0.80 3.57 2.81 1.23 0.97 0.15 5.96 0.87 0.15 0.01

69 681 55 724 386 33 45.5 95.6 46.5 9.33 3.11 8.90 6.58 3.29 2.49 0.35 12.4 1.77 0.14 0.10

3

He/4He RA [4He] c

8.96 66.16

6.93/6.86 b 56.09/5.98

7.50 11.06

6.79 2.06

7.73 2.12

7.66 7.66

8.31 4.80

9.54 16.97

87

0.703613 ± 9 0.512921 ± 7 19.578 ± 3 15.581 ± 2 39.210 ± 6

0.703721 ± 11

0.704027 ± 9

0.703682 ± 9 0.512865 ± 8 19.664 ± 3 15.621 ± 2 39.371 ± 7

0.703794 ± 9

0.703747 ± 10 0.512922 ± 8 19.646 ± 7 15.627 ± 6 39.351 ± 15

0.703444 ± 10 0.512929 ± 8 19.675 ± 5 15.595 ± 4 39.322 ± 10

0.705376 ± 14 0.512778 ± 7 19.099 ± 2 15.602 ± 2 39.018 ± 4

Sr/86Sr Nd/144Nd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb 143

19.609 ± 3 15.603 ± 2 39.238 ± 6

a m-alk bas = mildly alkalic basalt; thol bas = tholeiitic basalt; haw = hawaiite; mug = mugearite; picr bas = picritic basalt. HO-PUA-3 HO-HI-2 HO-HI-3 HO-HI-05 HO-AT-1 HO-AT-3 HO-AT-4 HO-AT-Pcr b Values for 2 consecutive steps of 5 min each. Hiva Hiva Oa Hiva Oa Hiva Oa Hiva Oa Hiva Oa Hiva Oa Hiva Oa c [He]Oa concentrations in 1 × 10− 9 cc(STP)/g. d Values m-alk basfor duplicate m-alkruns. bas m-alk bas m-alk bas haw m-alk bas picr bas

He isotope systematics. Thus only lavas that contain large (∼few mm to N1 cm) and fresh olivine crystals from readily accessible outcrops and cobbles in streambeds were collected. The OM and TOP samples were collected from the village of Omoa and HV samples from the village of Hanavave in the southern island of Fatu Hiva (FH-samples); PUA samples from Puamau Bay, HI from

NH-HA-1 Nuku Hiva m-alk bas

Hanaiapa Bay, AT-1 to -3 samples from the village of Atuona, and AT-4 and AT-Picr from Taahanka valley in the ∼central island of Hiva Oa (HO-samples); and HA samples from the village of Hatiheu in the ∼northern island of Nuku Hiva (NH-sample; Table 1 and Fig. 1). The exterior surfaces of the samples collected were trimmed and the interior portions were broken into

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Table 1 (continued) HO-PUA-3

HO-HI-2

HO-HI-3

HO-HI-05

HO-AT-1

HO-AT-3

HO-AT-4

HO-AT-Pcr

NH-HA-1

Hiva Oa

Hiva Oa

Hiva Oa

Hiva Oa

Hiva Oa

Hiva Oa

Hiva Oa

Hiva Oa

Nuku Hiva

m-alk bas

m-alk bas

m-alk bas

m-alk bas

haw

m-alk bas

picr bas

m-alk bas

45.16 2.94 10.04 12.63 0.15 14.31 8.33 1.93 0.99 0.33 96.81

45.42 3.90 11.70 12.71 0.15 9.67 9.53 2.13 1.41 0.42 97.04

45.29 3.94 11.55 12.35 0.14 8.90 9.42 2.10 1.43 0.38 95.50

44.95 2.41 7.41 14.33 0.16 17.84 7.98 1.42 0.87 0.22 97.60

47.55 3.38 14.22 11.34 0.15 6.75 8.25 2.90 2.32 0.57 97.43

45.42 3.90 11.70 12.71 0.15 9.67 9.53 2.13 1.41 0.42 97.04

39.50 0.27 0.99 17.26 0.18 37.85 1.04 0.12 0.08 0.06 97.33

48.34 3.26 13.27 12.47 0.15 7.19 10.48 2.41 0.99 0.28 98.86

68 677 50 738 391 35 12.9 30.9 20.6 4.98 1.14 5.13 3.88 1.78 1.29 0.19 13.6 1.73 0.13 0.09

30 471 49 669 351 27 36.5 79.2 41.5 8.62 3.06 8.02 5.57 2.50 1.75 0.23 9.58 1.34 0.14 0.04

31 498 51 678 353 25 38.8 85.6 46.5 9.78 3.49 8.71 6.33 2.73 2.02 0.31 9.70 1.41 0.15 0.05

26 325 25 292 187 19 15.8 36.9 21.5 4.88 0.87 4.92 3.73 1.78 1.31 0.16 13.1 1.73 0.13 0.09

74 803 56 712 496 37 38.3 76.6 39.6 8.21 3.11 7.89 5.94 3.01 2.40 0.32 14.3 1.62 0.11 0.10

57 715 47 719 336 30 26.5 59.3 32.1 6.73 1.18 6.36 4.55 2.13 1.47 0.21 15.3 2.13 0.14 0.08

4 16 8 27 28 3 0.76 2.44 1.88 0.26 0.09 0.22 0.23 0.12 0.15 1.95 0.55 0.28 0.15

19 573 29 446 257 35 23.6 51.0 31.7 7.13 2.79 7.46 5.92 3.10 2.35 0.33 19.9 2.22 0.11 0.04

9.03 4.60

8.21 22.34

8.34 26.15

8.69 30.92

1437/14.43 d 7.43/4.55

11.79 80.34

10.43 27.32

10.93 182.02

10.71 0.32

0.705338 ± 10 0.512765 ± 11 19.086 ± 2 15.578 ± 2 38.946 ± 5

0.704137 ± 9

0.704143 ± 10 0.512827 ± 6 19.336 ± 3 15.595 ± 2 39.224 ± 6

0.704174 ± 12

0.704711 ± 7 0.512834 ± 7 19.017 ± 3 15.570 ± 3 38.823 ± 9

0.705306 ± 6 0.512788 ± 8 19.003 ± 2 15.601 ± 2 38.901 ± 5

0.705126 ± 9 0.512781 ± 8 19.029 ± 2 15.586 ± 1 38.860 ± 4

0.704844 ± 16

0.704466 ± 9 0.512833 ± 9 19.142 ± 4 15.559 ± 4 39.050 ± 10

∼ 0.5 cm3 chips. Some of the chips were leached for ∼ 45 min in 2 N nitric acid and then powdered in an alumina ceramic container with an automated grinder. Major and trace element and Sr, Nd and Pb isotopic analyses were done on sample powders using procedures similar to those described in Janney and Castillo (1996, 1997). For major element analysis, aliquots of 3-gram mixture of sample powder and LiBO3 flux (5:1 mix) were melted to form glass discs. Major element

oxides were determined from the discs using a fully automated, wavelength-dispersive ARL 8410 X-ray fluorescence (XRF) instrument at the Scripps Institution of Oceanography (SIO). Calibration lines were constructed using up to 15 international rock standards that were prepared the same way as the unknown samples. For trace element analysis, about 50-mg powder of each sample was digested with a double-distilled, 2:1 mixture of concentrated HF:HNO3 acid in a clean Teflon

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beaker. The digested sample was diluted 1000-fold with a 1% HNO3 solution containing 100 ppb 115In as internal standard. Trace element concentrations were obtained using a VG PlasmaQuad 2+ inductively-coupled mass spectrometer (ICP-MS) at SIO. Calibration was performed using a series of 4 synthetic multi-element standard solutions. Time- and mass-dependent instrumental drift was corrected for by applying mass interpolated internal standard correction and by correcting measured sample concentrations with a well-analyzed in-house basalt standard analyzed between every 7 samples. Replicate analyses of international rock standards treated as unknown samples show major element precision to be between 0.3% and 2.5% and are considered accurate to about 1% for Si, Ti, Al, Fe, Mg, and Ca, and to within 3–5% for Mn, Na, K and P; trace element precision is within 3%, except those for Nb and Zr, which are precise to within 6%. More information on the estimated precisions for the major and trace element analyses are reported in Janney and Castillo (1996, 1997). Strontium, Nd and Pb isotope ratios were determined using a 9-collector, Micromass Sector 54 themal ionization mass spectrometer (TIMS) at SIO. Rock powders were digested using the same procedure for trace element analysis described above. Lead was first separated by re-dissolving the dried samples in 1 N HBr and then passing the solutions in a small ion exchange column in an HBr medium. Strontium and rare earth elements (REE) were separated from the residual solutions in an ion exchange column using HCl as eluent. Finally, Nd was separated from the rest of the REE in an ion exchange column using alpha hydroxyisobutyric acid as eluent. Total procedural blanks are b35 pg for Sr, b10 pg for Nd and b30 pg for Pb. Strontium isotopic ratios were fractionation-corrected to 86Sr/88Sr = 0.1194 and are reported relative to 87 Sr/86Sr = 0.710257 for NBS 987. Neodymium isotopic ratios were measured in oxide form, fractionation corrected to 146 NdO/144NdO = 0.72225 (146 Nd/144Nd = 0.7219) and are normalized to 143Nd/144Nd = 0.511850 for the La Jolla Nd Standard. Lead isotopic ratios were corrected for mass fractionation based on average measured values of NBS 981 using the values of Todt et al. (1996). Analytical uncertainty based on repeated measurements of standards is ± 0.000018 for 87Sr/86Sr, ± 0.000016 for 143Nd/144 Nd, ± 0.008 for 206 Pb/204 Pb and 207Pb/204Pb, and ± 0.030 208Pb/204Pb. 2σ precisions listed in Table 1 refer to within-run statistics. Chips of the remaining interior portions of the samples were ground and fresh olivine grains were picked under a binocular microscope. Helium gas trapped in fluid inclusions in the grains was extracted using the

‘crushing under vacuum’ procedure described in Scarsi (2000). A single crushing step in 8 min was adopted and the gas component was transferred into the inlet system of the mass spectrometer. After the Ne gas was cryogenically removed using a charcoal trap at 34 °K, 3 He/4He ratio and He concentration were measured statically on the double-collecting split-tube mass spectrometer GAD at the SIO Isotope Lab using the procedure described in Rison and Craig (1983). 3 Helium was measured by ion counting with a Johnston electron multiplier and 4He by collection in a Faraday cup. The measured isotope ratio was normalized to aliquots of a high 3He/4He (R/RA = 16.45) standard gas (MM). Blank was measured before each experiment and the system is pumped out for a significant amount of time to minimize possible memory effects. The error for 3 He/4He ratios is normally b 10% with respect to the natural dispersion for the measured ratios and was obtained by adding instrumental errors on the 3He and 4 He measurements on samples with negligible blank contribution (generally b 10% for samples with [He] N 1 × 10− 9 cc(STP)/g). The possible contribution of cosmogenic 3He component to the He gas trapped in fluid inclusions as well as accuracy and precision of the analytical method were also monitored through a crushing experiment performed on sample FH-OM-3. This sample was crushed in 2 steps of 5 min each and although the concentrations of He extracted in the 2 steps are very different (56.085 and 5.984 × 10− 9 cc(STP)/g), their 3He/4He ratios (6.93 and 6.86 RA, respectively) are identical within errors. Moreover, sample HO-AT-01 was analyzed twice and the 3He/4He analyses (14.37 and 14.43 RA) are also identical (Table 1). 4. Analytical results The major and trace element results are presented in Table 1. Megascopic examination reveals that the lava samples generally have fine-grained matrix and majority of them have accumulated olivine phenocrysts during crystal fractionation. Thus the major and some of the trace element analyses may have been compromised. This is particularly true for HO-AT-Picr, which is a picritic basalt and has 37.9 wt.% MgO. The proportions of olivine in the other samples are variable and their MgO contents range from 3.5 to 25.5 wt.%. At any rate, based on total alkalis (wt.% Na2O + wt.% K2O) versus silica (wt.% SiO2; Le Bas et al., 1986), two of the Fatu Hiva samples are classified as tholeiitic basalts, one as a mugearite and the rest as mildly alkalic basalts. Two of the Hiva Oa samples are hawaiites, one is the picritic

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basalt, which is similar to the tholeiitic picrobasalt collected from Eiao Island (Caroff et al., 1999), and the rest are mildly alkalic basalts. The lone sample from Nuku Hiva is also a mildly alkalic basalt. For comparison, all volcanic rocks previously collected from Fatu Hiva are tholeiitic to mildly alkalic and belong to the shield-building phase only whereas volcanics from Hiva Oa and Nuku Hiva belong to both shield-building and post-shield phases of volcanism (e.g., Duncan and McDougall, 1974; Duncan et al., 1986; Woodhead, 1992; Desonie et al., 1993). Despite the possible effects of olivine accumulation, clearly our samples are mildly to highly enriched in incompatible trace elements, similar to previous results. In general, Fatu Hiva basalts are only mildly enriched in incompatible trace elements whereas Hiva Oa and Nuku Hiva basalts are highly enriched. The trace element variations of our samples overlap with those of the previously analyzed mafic (i.e., basalt to mugearite) volcanics (Table 1; Fig. 2A to D). The large ranges of their highly to moderately incompatible trace elements ratios such as Ba/Zr (0.6–2.2), Ba/Nb (2–20), Nb/Zr (0.1–1.5), and Rb/Sr (.01–0.1) strongly indicate a compositionally variable mantle source. The isotope analyses are also presented in Table 1. Olivine does not contain appreciable amounts of Rb, Sr, Sm, Nd, U, Th, and Pb, and thus we believe that our isotopic analyses represent bulk rock compositions. Indeed, the range of our new Sr, Nd and Pb isotopes overlaps with previous Marquesan isotope values (Fig. 3A to C). Consistent with incompatible trace element ratios, 87Sr/86Sr ratios span a relatively large range, strongly indicating a compositionally heterogeneous mantle source. Isotopes of Nd and Pb also show distinct, though more limited variations. A closer inspection of the new data reveals that the Pb isotopes define two distinct groups. Fatu Hiva samples have higher 206Pb/204Pb but ∼similar 207 Pb/204Pb and 208Pb/204Pb values as Hiva Oa and Nuku Hiva samples. The new Fatu Hiva Pb isotopes overlap with previous analyses for tholeiitic lavas such as those from Ua Pou, which plot below the Northern Hemisphere Reference Line (NHRL — Hart, 1984) in Pb isotope diagrams and thus were previously doubted as they may have been affected by some form of Pb contamination (Woodhead, 1992). Nevertheless, similar to our 87 Sr/86Sr and 143Nd/144 Nd, 206Pb/204Pb ratios generally correlate with some of the highly to moderately incompatible element ratios (Fig. 4A to C). Majority of the samples analyzed have 3He/4He ratios around 8 RA similar to previous results (Desonie et al., 1991), but there are some samples with 3He/4He

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ratios that range up to ∼ 14 RA (Table 1). Fatu Hiva samples have a narrower and generally lower range (6.8 to 9.0 RA) than Hiva Oa and Nuku Hiva samples (8.2 to 14.4 RA). There is no correlation between He isotopic ratios and concentrations (Table 1), but 3He/4He ratios correlate with some of the incompatible trace element ratios (Fig. 4D). Moreover, 3 He/ 4 He ratios show systematic relationships with Sr, Nd and Pb isotopic ratios, but these are more complex and will be discussed in a later section. Thus data strongly suggest that the 3 He/4He ratios are magmatic in origin despite the fact that He gas was extracted from olivine phenocrysts (see also, Kurz et al., 1996; Kurz, 2005). 5. Discussion 5.1. Volcanologic and geochemical evolution of the Marquesas Archipelago It is pleasantly surprising that despite the limited number, reconnaissance nature of sampling and variable effects of olivine accumulation on major and trace element chemistry, our samples fall within the established evolutionary scheme of Marquesan volcanism. Except for the picritic basalt from Hiva Oa, Fatu Hiva samples have the only tholeiitic basalts, are only mildly enriched in incompatible trace elements and the average values of their Ba/Zr, Ba/Nb and Rb/Sr ratios are lower than those of Hiva Oa and Nuku Hiva samples (Table 1 and Fig. 2). More important, Fatu Hiva samples have low 87Sr/86Sr (0.7034 to 0.7040) and high 143Nd/144 Nd (0.51287 to 0.51293) ratios that are respectively within the values (0.7030 to 0.7040 and 0.51274 to 0.51294, respectively) for the shield-building phase lavas (Woodhead, 1992). Thus our Fatu Hiva samples clearly were erupted during the shield-building phase of volcanism. On the other hand, our Hiva Oa and Nuku Hiva samples have high 87 Sr/ 86 Sr (0.7041 to 0.7054) and low 143 Nd/144Nd (0.51278 to 0.51283) ratios that also fall respectively within the values (N 0.7041 and b 0.51285) of the post-shield phase lavas. Thus our Hiva Oa and Nuku Hiva samples including the picritic basalt, by virtue of its high 87Sr/86Sr ratio (0.7048), represent postshield volcanic phase only. Despite their limited range, the Pb isotopes also provide an important criterion to differentiate the tholeiitic, shield-building phase from the alkalic, postshield phase of volcanism. Shield-building phase lavas have slightly higher 206Pb/204 Pb ratios than post-shield phase lavas and the 206Pb/204Pb ratios as a whole correlate with some geochemical tracers (Fig. 4B; see also Desonie et al., 1993). Moreover, we speculate that

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Fig. 2. Incompatible trace element diagrams for basaltic, hawaiitic and mugearitic volcanic rocks from the Marquesas Archipelago. (A.) Ba vs. Nb, (B.) Ba vs. Zr, (C.) Nb vs. Zr, and (D.) Rb vs. Sr. Shown for reference are Marquesan lavas from the GEOROC (2006) electronic database and typical values for normal mid-ocean ridge basalts (solid line) and ocean island basalts (dash line) from Sun and McDonough (1989).

the ‘unusually’ ∼low 207 Pb/204Pb and 208Pb/204Pb values of the shield-building phase lavas actually provide an additional insight into the nature of their source. Recall that the previous Pb isotopes were doubted because these may have been affected by some form of contamination (Woodhead, 1992). The suspicion originates from a study by McDonough and Chauvel (1991) suggesting that a similar 206Pb/204Pb enrichment but 207 Pb/204Pb and 208Pb/204Pb depletion in some Rurutu Island basalts results from Pb contamination ‘prior’ to isotopic analysis. Despite the severe acid-leaching procedure that produced a more coherent and tighter Pb isotopic composition for all Rurutu basalts in that study, however, some Rurutu basalts still plot below NHRL, especially in the 208Pb/204 Pb versus 206 Pb/204Pb diagram. Although Chauvel et al. (1997) later showed that the contaminant in the basaltic rocks produced during the later Rurutu volcanic phase is a residual carbonititic liquid in the lithosphere, here we suggest that the contaminant in the initial Marquesan shieldbuilding volcanic phase is the 50–65 m.y. old, variably

altered oceanic crust that currently underlies the islands and has high present day 206 Pb/204 Pb and ∼low 207 Pb/204Pb and 208 Pb/204Pb ratios (e.g., Janney and Castillo, 1997; Hauff et al., 2003). This old crust has a high 238U/204 Pb (μ) ratio most likely due to addition of seawater derived U and thus has high 206Pb/204 Pb, but ∼low 207 Pb/ 204 Pb ratio because most 235 U have decayed in early part of the Earth's history. Thorium, unlike U, is not appreciably added into the oceanic crust and hence 208Pb/204 Pb is also not enriched. Although this topic is important, it is beyond the scope of this paper and thus will not be discussed further here. It is important to note, however, that the hint of crustal assimilation provided by the Pb isotopes is consistent with some other geochemical data suggesting that the shield-building phase volcanism derives some of its lavas from the lithosphere (e.g., Duncan et al., 1986; Woodhead, 1992; Desonie et al., 1993). An indirect evidence for crustal assimilation of the shield-building phase volcanics is provided by the presence of xenoliths and xenocrysts plus high oxidation

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Fig. 3. (A.) 87Sr/86Sr vs. 143Nd/144Nd, (B.) 206Pb/204Pb vs. 207Pb/204Pb and (C.) 206Pb/204Pb vs. 208Pb/204Pb diagrams for Marquesan lavas. Symbols and reference for Marquesan lavas as in Fig. 2. Fields for lavas from other oceanic linear volcanic chains (Zindler and Hart, 1986; Dupuy et al., 1987; Vidal et al., 1987; Hart, 1988; Farley et al., 1992, 1993; Caroff et al., 1999; Hofmann, 2003; Legendre et al., 2005a) and the proposed DMM, HIMU, EM1, and EM2 mantle source components (Zindler and Hart, 1986; Hart, 1988) are shown for reference. Note that the new analyses for Marquesan lavas are clearly divided into two groups in the 206Pb/204Pb vs. 208Pb/204Pb diagram.

states of some Marquesan lavas (e.g., Caroff et al., 1999; Legendre et al., 2005a,b). Although these studies focus on alkalic differentiates of the post-shield volcanic phase, they clearly show the importance of crustal assimilation in the petrogenesis of linear volcanic chain lavas. In the Marquesas, we speculate that the vigorous shield-building phase volcanism causes large scale and efficient assimilation of the pre-existing oceanic crust whereas the smaller volume post-shield phase causes localized

assimilation and/or melting of the shield-building volcanics (e.g., Caroff et al., 1999; Legendre et al., 2005a,b). Our results also show that He isotopes provide another criterion for distinguishing the two phases of volcanism because the post-shield lavas have higher 3He/4He ratios than the shield-building lavas. This is consistent with the proposal that the former may be coming directly from the mantle (Woodhead, 1992). On the other hand, the mantlederived parental magmas of the shield-building lavas have

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Fig. 4. Rb/Sr vs. (A.) 87Sr/86Sr, (B.) 206Pb/204Pb, (C.) 143Nd/144Nd, and (D.) 3He/ 4He ratios for Marquesan lavas. Symbols and reference for Marquesan lavas as in Fig. 2. Shield-building Fatu Hiva volcanic rocks have lower Rb/Sr, 87Sr/86Sr and 3He/ 4He, but higher 206Pb/204Pb and 143 Nd/144Nd ratios than Hiva Oa and Nuku Hiva volcanic rocks.

been severely affected by assimilation of the oceanic crust that is enriched in U, a major source of 4He (e.g., Hilton et al., 1995). However, the bulk of the radiogenic 4He produced in the crust is most probably not retained because the U decay process happens in the submarine environment in the presence of circulating fluids, as evidently shown by the occurrence of many calcite veins and zeolite amygdules in old oceanic crust (e.g., Lancelot et al., 1990). More important, the glassy rind, olivine and clinopyroxene that trap He gas have already solidified and in fact variably altered in the oceanic crust. Hence assimilation of crustal materials was able to lower the 3 He/4He signature of the initial shield-building phase parental magmas only to the nominal upper mantle value of ∼8 RA. A similar temporal He isotopic evolution so far is observed only in Hawaii, although the high 3He/4He ratios there are a characteristic feature of the shieldbuilding phase (e.g., Kurz, 2005). However, this may not come as a surprise as Woodhead (1992) has previously shown based on other geochemical grounds that the Marquesas and Hawaii respectively represent the weak and strong end-members in terms of plume– lithosphere interaction. At any rate, we concur with

Kurz (2005) that such a temporal isotopic evolution provides important constraints on ocean island volcanism, but has not been fully established in the global geochemical compilations. 5.2. The mantle source of Marquesan lavas Previous isotopic analyses of Marquesan volcanic rocks define a wide, but roughly elongate field with two ends clearly pointing toward EM2 and HIMU mantle components (Fig. 5A). Most of the HIMU lavas belong to the shield-building phase and define a pseudo linear array that trends toward the DMM component whereas EM2 lavas generally belong to the post-shield phase and define another pseudo linear array that trends toward the HIMU–DMM join (e.g., Liotard et al., 1986; Dupuy et al., 1987; Vidal et al., 1987; Desonie et al., 1993). The apparent mixing among EM2, HIMU and DMM mantle components forms the basis for the heterogeneous mantle plume source of Marquesan lavas (represented by dash lines in Fig. 5A). Our new data define a narrower, elongate array that overlaps with previous analyses. Although the ends of the new array also point toward EM2 and HIMU, there is a small bend near the

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Fig. 5. (A.) 87Sr/86Sr vs. 206Pb/204Pb diagram for Marquesan volcanic rocks (slightly modified after Desonie et al., 1993). Symbols and reference for Marquesan lavas as in Fig. 2. Dash lines with arrows are the previously proposed binary mixing relationships between HIMU, EM2 and DMM mantle components (Dupuy et al., 1987; Vidal et al., 1987; Desonie et al., 1993). Also shown for reference is the field for Iceland lavas (White, 1985; Zindler and Hart, 1986; Hart, 1988; Taylor et al., 1997; Hilton et al., 1999; Hofmann, 2003); open circles represent some Iceland samples with He isotope data (Taylor et al., 1997; Hilton et al., 1999). (B.) 3He/ 4He vs. 87Sr/86Sr vs. 206Pb/204Pb diagram for the same samples shown in (A.). Symbols with lines represent 3He/ 4He values N8 RA; short solid lines with arrows are our proposed binary mixing relationships between a mantle source with a high 3 He/ 4He signature and EM2 or HIMU components. There is no direct mixing between EM2 and HIMU mantle components; there may be an additional DMM signature coming from the oceanic lithosphere and this mixes with both HIMU and EM2 components.

EM2 end of the new data array that appears to point toward the DMM–HIMU join, similar to previous results. The new He isotopic data, however, define a slightly different set of mantle components. In the 3He/4He vs. 87 Sr/86Sr diagram, Marquesan samples define a rough, inverted ‘V’ pattern whose corners point toward EM2, primitive helium mantle (PHEM) and HIMU or DMM, which occupy the same corner of the diagram (Fig. 6A). The 3He/4He ratios of the post-shield Hiva Oa and Nuku Hiva samples with high 87Sr/86Sr (N 0.7041) generally

point to these three end-members whereas those of the shield-building Fatu Hiva samples with low 87Sr/86Sr (b0.7040) remain at 7 to 9 RA. In the 3He/4He vs. 143 Nd/144Nd diagram, however, Marquesan samples show a roughly triangular array, the corners of which are hinting toward HIMU, PHEM and EM2, but not toward DMM (Fig. 6B). Finally, in the 3He/4He vs. 206 Pb/204Pb diagram, Marquesan samples define a curvilinear array, the two ends of which appear to trend to HIMU and PHEM and the middle curving toward EM2; again, there is no trend toward DMM (Fig. 6C). Thus the

Fig. 6. 3He/ 4He versus (A.) 87Sr/86Sr, (B.) 143Nd/144Nd and (C.) 206Pb/204Pb diagrams for Marquesan lavas. Same symbols as in Fig. 2. In the diagrams, linear fields or arrays of volcanic chain lavas are interpreted to be the results of mixing among proposed mantle source components (Zindler and Hart, 1986; Hart, 1988; Hart et al., 1992; Hofmann, 2003).

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presence of DMM in the source of Marquesan volcanics is only a plotting artifact in two-dimensional Sr, Nd and Pb isotope diagrams as the combined Pb, Nd and He isotopic data do not indicate the presence of DMM. Our combined Sr, Nd, Pb, and He isotopes suggest that some if not all Marquesan lavas that trend toward the low 87 Sr/86Sr and 206Pb/204Pb direction of DMM have higher 3 He/4He ratios than typical MORB (Fig. 5B). Thus we believe that some Marquesan lavas, particularly those belonging to the post-shield volcanic phase, have a distinct mantle source with high 3He/4He ratios and DMM-like characteristics, represented by some lavas from Iceland. We also believe that additional DMM-like characteristics of the shield-building volcanic phase lavas may be coming from the assimilated oceanic crust. The rationale for proposing that DMM is in the mantle source of Marquesan lavas was partly influenced by the landmark study by Hart et al. (1992), which states that the majority of volcanic chain lavas define binary mixing arrays in the Sr–Nd–Pb isotopic space. One end of the mixing arrays trends to either one of the recycled mantle components (HIMU, EM1 or EM2), but the other end converges to a focal zone, which was originally thought to have a mixed HIMU–DMM composition. As the Marquesan volcanics also define a linear isotopic array in the Sr–Nd–Pb isotopic space, by analogy it was concluded that this array represents binary mixing between EM2 and the mixed HIMU–DMM source (Desonie et al., 1993). The focal zone, however, has since then been relocated and most mantle geochemists now agree that the isotopic convergence lies internal to the Sr–Nd–Pb isotopic space (e.g., Hanan and Graham, 1996; Hilton et al., 1999; Hofmann, 2003). Most important, a number of studies have suggested that the focal zone has high 3He/4He ratios (e.g., Farley et al., 1992, 1993; Hart et al., 1992; Hilton et al., 1999). We therefore conclude that the mantle source of the Marquesan volcanics consists of HIMU, EM2 and a high 3He/4He mantle component that lies relatively close to DMM in terms of Sr, Nd and Pb isotopes, which we designate here as PHEM (Fig. 6A to C; see also, recent review of Class and Goldstein, 2005). 5.3. Implications for the origin of linear volcanic chains Both the hotspot and alternate plate stress hypotheses for the origin of linear volcanic chains envision that crustal materials recycled in the mantle comprise the bulk of the source of their lavas, either as upwelling plumes from the deep mantle or as heterogeneities in the upper mantle, respectively. Thus to a first order, the geochemistry part of the debate on the origin of linear

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volcanic chains can be narrowed down to the nature and origin of high 3He/4He lavas. The contemporary hotspot hypothesis claims that the most important terrestrial source of 3He is degassing from the Earth's interior and thus, 3He/4He is a unique tracer of mantle sources involved in volcanism (e.g., Kurz et al., 1982, 1983; Lupton, 1983; Staudacher and Allegre, 1989; Graham et al., 1993; Hilton et al., 1999; Graham, 2002; Porcelli and Ballentine, 2002). The difference in the 3He/4He ratio observed between MORB and ocean island basalts (OIB), which are predominantly erupted along linear volcanic chains, is cited as evidence for two distinct mantle sources of oceanic lavas. MORB have a narrow range of 3He/4He (average ∼ 8 RA) whereas OIB have a wider, but higher range (highest values to date are up to ∼ 50 RA from olivines from picritic lavas in Baffin Island — Stuart et al., 2003). This is taken as evidence for the existence of an upper, uniformly degassed DMM source and a separate, lower OIB source containing different plumes, at least one of which taps noble gases from primitive, undegassed regions of the deeper mantle that may have remained isolated over geologic time. Meibom et al. (2003; Meibom and Anderson, 2004; Anderson, 2001), on the other hand, argue that the use of He isotopic signatures as a fingerprint of plume components in oceanic lavas is not justified. They argue that the average He isotopic signatures of OIB and unfiltered MORB with 95% confidence are drawn from the same statistical population. The data display a wide, nearly Gaussian distribution produced by shallow mantle processes involving mixing between different proportions of recycled, high and low 3He/4He end-members produced by the variation of (U + Th)/He concentration ratios. For example, the high 3He/4He end-member can result from a decrease of 4He-producing (U + Th) relative to He concentration in olivines whereas the higher than chondritic (U + Th)/He elemental ratio in the evolving and partially degassed MORB melt can provide the radiogenic (i.e., low 3 He/4 He) end-member during partial melting events in the upper mantle (e.g., Brooker et al., 2003a,b). High 3He/4He ratios can also result from the capture of noble gas-rich CO2 bubbles, but not of Th and U, by olivine crystals during their growth in shallow mantle conditions (e.g., Natland, 2003). Finally, 3 He/4He ratios in linear volcanic chain lavas can also be increased by fine extraterrestrial materials, which have much higher 3He/4He than terrestrial materials (N1 × 10− 4 versus b 1 × 10− 7, respectively). These materials accumulate on the seafloor, get subducted and together with the subducted crustal materials form the high 3He/4He source in the upper mantle (Anderson, 1993). 3Helium flux (∼ 1 × 10− 15 cc/(cm2 yr)) and 3He

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source must have changed over geologic time (Farley et al., 1997; Mukhopadhyay et al., 2003) and thus if these were greater in the past, then the extraterrestrial origin of the current mantle source for many volcanic chain lavas cannot be discounted. In summary, the plate stress hypothesis claims that there is no need for a distinct mantle source for the high 3He/4He lavas. Mantle melting processes, such as sampling upon melting and averaging (SUMA) process, can sample shallow mantle materials with different proportions of high and low 3 He/4He end-members through variations in the degrees of melting (Meibom and Anderson, 2004). Despite their limited range in values, our new He isotope data are similar to those for other volcanic chain lavas that show systematic correlations with other geochemical tracers (Class and Goldstein, 2005). Moreover, many of the lavas indicate mixing between EM2, EM1 and HIMU components on the one hand and a source with a high 3He/4He and common Sr, Nd and Pb isotopes on the other, but generally there is no mixing between these components nor between these components and MORB (Hart et al., 1992; see also, Saal et al., 1998). These systematic compositional features make it highly unlikely for linear volcanic chain lavas to originate from a multitude of dispersed, recycled materials in the upper mantle. Specifically, although the high 3He/4He signature can result from (U + Th)/He fractionation processes in the shallow mantle, the coupling of 3He/4He ratios to other geochemical tracers suggests that the signature is an inherent feature of an ∼old mantle source. Moreover, the SUMA concept should not prohibit mixing between different mantle components and especially between mantle components and DMM to form a statistical blend of contributions from the various mantle components, and yet such mixing relations are not seen in the majority of data. Finally, there are serious problems with the amount of the high 3He/4He extra terrestrial materials and their survival during the subduction process (e.g., Farley et al., 1997) and hence it is unlikely that they can be a viable source of the high 3He/4He signature of linear volcanic chain lavas. Even if we assume that extraterrestrial materials can be subducted into the mantle and form the high 3He/4He mantle source, such a source would definitely yield linear volcanic chain lavas with a high and variable 87Sr/86Sr isotopic signature because seawater alters the Sr isotopic composition of the uppermost oceanic crust and particularly its marine sediment cover that carries these fine materials prior to subduction. Our new data, similar to those for many linear volcanic islands, indicate that a high 3He/4He mantle component, the most controversial mantle source of

linear volcanic chain lavas, exists. The behavior of this source is far from random as it is a ubiquitous source for many of the volcanic chain lavas (Kurz et al., 1982, 1983; Hart et al., 1992; Graham et al., 1993; Hanan and Graham, 1996; Hilton et al., 1999; Graham, 2002; Class and Goldstein, 2005). This is a critical geochemical argument favoring the hotspot hypothesis. It is important to note, however, that available data cannot differentiate whether this source is a portion of the mantle that is still primitive and undegassed (e.g., Kurz et al., 1982, 1983; Farley et al., 1992, 1993; Graham, 2002) or an ancient residue of depletion in incompatible elements including He by the formation of continents early in the Earth history (Class and Goldstein, 2005) or an ancient residue of depletion in incompatible elements including Th and U but not He during mantle melting (Parman et al., 2005). 6. Conclusions 1. Our Fatu Hiva samples are tholeiitic to mildly alkalic in composition, mildly enriched in incompatible trace elements, have ∼low 87Sr/86Sr but ∼high 143Nd/144Nd ratios, and thus belong to the initial shield-building phase of volcanism in the Marquesas. Our Hiva Oa and Nuku Hiva samples are almost all alkalic in composition, more enriched in incompatible trace elements, have ∼high 87Sr/86Sr but ∼low 143Nd/144Nd ratios, and thus belong to the later post-shield phase. 2. Lead isotopic results reiterate the modestly higher 206 Pb/204Pb characteristic and also seem to provide evidence for an additional crustal source of the shield-building lavas. 3. The coherent relationship between He isotopes and other geochemical tracers indicates that the postshield have higher 3 He/4 He ratios than shieldbuilding lavas and that a DMM source is not directly involved in the production of Marquesan magmas. Instead, a distinct, high 3He/4He mantle source with a DMM-like geochemical signature, is indicated by the post-shield lavas. 4. Based on the combined new and literature data, we propose an alternative model to explain the evolution of Marquesan volcanism that draws heavily from previously proposed hotspot models, particularly that of Woodhead (1992). The Marquesan intraplate volcanism is weaker compared to that in Hawaii, and is thus greatly affected by the state of stress and thermal structure in the lithosphere. Specifically, its proximity to the Marquesan fracture zone and location above a thermally anomalous region (e.g., McNutt and Fischer, 1987; Janney and Castillo,

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1996, 1997) facilitate the assimilation of the oceanic crust during the initial shield-building volcanic phase. This is the most vigorous phase of volcanism as it forms the bulk of the islands, and thus a large volume of the oceanic lithosphere is involved, giving the impression that volcanism is mainly a crustal phenomenon (e.g., Gutscher et al., 1999; McNutt and Bonneville, 2000; Guille et al., 2002). Regardless of the exact origin of linear volcanic chain, there is definitely a melting anomaly beneath the Marquesas and its mantle source is compositionally heterogeneous. A HIMU component is a predominant source during the shield-building volcanic phase, but it is mixed heavily with assimilated, high υ crustal materials. The post-shield volcanic phase samples the waning stage of the melting anomaly. The mantle source during this phase is mainly a mixture of EM2 and high 3He/4He components. 5. The proposed DMM compositional signature in the Marquesan lavas is mainly coming from the high 3 He/4He mantle component and partly from assimilated oceanic lithosphere. The high 3He/4He signature in the Marquesan lavas cannot simply be a product of shallow mantle processes as predicted by the plate stress hypothesis. Instead, the high 3He/4He ratios appear to be a characteristic feature of the common mantle source of many linear volcanic chains as claimed by the hotspot hypothesis. We thus concur with the hotspot hypothesis that the most likely location of this mantle component is in the deeper part of the mantle, where subducted crustal materials also accumulate and age (Gurnis and Davies, 1986; Davies, 2002). Together, the high 3 He/4He portion of the mantle and recycled crustal materials form the heterogeneous plume source of many, but definitely not all linear volcanic chains. Acknowledgements This study was initiated and inspired by the late H. Craig. We acknowledge the laboratory assistance of C. MacIsaac, stimulating discussions with J. Hawkins and E. Winterer and careful and thorough reviews by R. Rudnick, M. Kurz and an anonymous reviewer, which greatly improved the quality of this manuscript. References Anderson, D.L., 1993. Helium-3 from the mantle: primordial signal or cosmic dust? Science 261, 170–176. Anderson, D.L., 1996. Enriched asthenosphere and depleted plumes. International Geology Review 38, 1–21.

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Anderson, D.L., 2001. A statistical test of the two reservoir model for helium isotopes. Earth and Planetary Science Letters 193, 77–82. Brooker, R.A., Heber, V., Kelley, S.P., Wood, B.J., 2003a. Noble gas partitioning behavior during mantle melting: a possible explanation for “The He Paradox”? (abstract). EOS, Transactions of the American Geophysical Union 31, F03. Brooker, R.A., Du, Z., Blundy, J.D., Kelley, S.P., Allan, N.L., Wood, B.J., Chamorro, E.M., Wartho, J.-A., Purton, J.A., 2003b. The ‘ zero charge’ partitioning behavior of noble gases during mantle melting. Nature 423, 738–741. Caroff, M., Guillou, H., Lamiaux, M., Maury, R.C., Guille, G., Cotten, J., 1999. Assimilation of ocean crust by hawaiitic and mugearitic magmas: an example from Eiao, Marquesas. Lithos 46, 235–258. Chauvel, C., McDonough, W., Guille, G., Maury, R., Duncan, R., 1997. Contrasting old and young volcanism in Rurutu Island, Austral chain. Chemical Geology 139, 125–143. Class, C., Goldstein, S.L., 2005. Evolution of helium isotopes in the Earth's mantle. Nature 436, 1107–1112. Davies, G.F., 2002. Stirring geochemistry in mantle convection models with stiff plates and slabs. Geochemica et Cosmochimica Acta 66, 3125–3142. DePaolo, D.J., Manga, M., 2003. Deep origin of hotspots — the mantle plume model. Science 300, 920–921. Desonie, D.L., Duncan, R.A., Kurz, M.D., 1991. Helium isotopic composition of isotopically diverse basalts from hotspot volcanic lineaments in French Polynesia (abs). EOS, Transactions of the American Geophysical Union 72, 536. Desonie, D.L., Duncan, R.A., Natland, J.H., 1993. Temporal and geochemical variability of volcanic products of the Marquesas hotspot. Journal of Geophysical Research 98, 17649–17665. Duncan, R.A., McDougall, I., 1974. Migration of volcanism with time in the Marquesas Islands, French Polynesia. Earth and Planetary Science Letters 21, 414–420. Duncan, R.A., McCulloch, M.T., Barsczus, H.G., Nelson, D.R., 1986. Plume versus lithospheric sources for melts at Ua Pou, Marquesas Islands. Nature 322, 534–538. Dupuy, C., Vidal, P., Barsczus, H.G., Chauvel, C., 1987. Origin of basalts from the Marquesas Archipelago (south Central Pacific Ocean): isotope and trace element constraints. Earth and Planetary Science Letters 82, 145–152. Farley, K.A., Natland, J.H., Craig, H., 1992. Binary mixing of enriched and undegassed (primitive?) mantle components (He, Sr, Nd, Pb) in Samoan lavas. Earth and Planetary Science Letters 111, 183–199. Farley, K.A., Basu, A.R., Craig, H., 1993. He, Sr, and Nd isotopic variations in lavas from the Juan Fernandez Archipelago. Contributions to Mineralogy and Petrology 115, 75–87. Farley, K.A., Love, S.G., Patterson, D.B., 1997. Atmospheric entry heating and helium retentivity of interplanetary dust particles. Geochemica et Cosmochimica Acta 61, 2309–2331. Foulger, G.R., Anderson, D.L., 2005. A cool model for the Iceland hotspot. Journal of Volcanology and Geothermal Research 141, 1–22. Foulger, G.R., Natland, J.H., 2003. Is “Hotspot” volcanism a consequence of plate tectonics? Science 300, 921–922. Foulger, G.R., Sheth, H.C., et al., 2004. Discussing the Origin of Hotspot Magmatism. http://www.mantleplumes.org. GEOROC, 2006. Geochemistry of Rocks of the Oceans and Continents. Max Planck Institute für Chemie, http://georoc. mpch-mainz.gwdg.de/georoc/Entry.html.

220

P.R. Castillo et al. / Chemical Geology 240 (2007) 205–221

Graham, D.W., 2002. Noble gas isotope geochemistry of mid-ocean ridge and ocean island basalts: characterization of mantle source reservoirs. Reviews of Mineralogy and Geochemistry 47, 247–317. Graham, D.W., Christie, D.M., Harpp, K.S., Lupton, J.E., 1993. Mantle plume helium in submarine basalts from the Galapagos platform. Science 262, 2023–2026. Guille, G., Legendre, C., Maury, R.C., Caroff, M., Unschy, M., Balis, S., Chauvel, C., Cotten, J., Guillou, H., 2002. Les Marquises (Polynesie francaise): un archipel intraoceanique atypique. Geologie de la France 2, 5–36. Gurnis, M., Davies, G.F., 1986. The effect of depth-dependent viscosity on convective mixing in the mantle and the possible survival of primitive mantle. Geophysical Research Letters 13, 541–544. Gutscher, M.A., Olivet, J.L., Aslanian, D., Eissen, J.P., Maury, R.C., 1999. The lost Inca Plateau: cause of flat subduction beneath Peru? Earth and Planetary Science Letters 171, 335–341. Hanan, B.B., Graham, D.W., 1996. Lead and helium isotope evidence from oceanic basalts for a common deep source of mantle plumes. Science 272, 991–995. Hart, S.R., 1984. A large-scale isotope anomaly in the Southern Hemisphere mantle. Nature 309, 753–757. Hart, S.R., 1988. Heterogeneous mantle domains: signatures, genesis, and mixing chronologies. Earth and Planetary Science Letters 90, 273–296. Hart, S.R., Hauri, E.H., Oschmann, L.A., Whitehead, J.A., 1992. Mantle plumes and entrainment: isotopic evidence. Science 256, 517–520. Hauff, F., Hoernle, K., Schmidt, A., 2003. Sr–Nd–Pb composition of Mesozoic Pacific oceanic crust (Site 1149 and 801, ODP Leg 185): implications for alteration of oceanic crust and the input into the Iz–Bonin–Mariana subduction system. Geochemistry, Geophysics, Geosystems 4. doi:10.1029/2002GC000421. Hilton, D.R., Barling, J., Wheller, G.E., 1995. Effects of shallow-level contamination on the helium isotope systematics of ocean–island lavas. Nature 373, 330–333. Hilton, D.R., Gronvold, K., Macpherson, C.G., Castillo, P.R., 1999. Extreme 3He/ 4He ratios in northwest Iceland: constraining the common component in mantle plumes. Earth and Planetary Science Letters 173, 53–60. Hofmann, A.W., 2003. Sampling mantle heterogeneity through oceanic basalts: isotopes and trace elements. In: Carlson, R.W. (Ed.), The Mantle and Core. Treatise on Geochemistry, vol. 2, pp. 61–101. Janney, P.E., Castillo, P.R., 1996. Basalts from the Central Pacific Basin: evidence for the origin of Cretaceous igneous complexes in the Jurassic Western Pacific. Journal of Geophysical Research 101, 2875–2894. Janney, P.E., Castillo, P.R., 1997. Geochemistry of Mesozoic Pacific MORB: constraints on melt generation and the evolution of the Pacific upper mantle. Journal of Geophysical Research 102, 5207–5229. Kurz, M., 2005. Spatial and temporal isotopic variability in ocean island volcanism: the noble gas story. Geochimica et Cosmochimica Acta 69, A95. Kurz, M.D., Jenkins, W.J., Hart, S.R., 1982. Helium isotopic systematics of oceanic islands and mantle heterogeneity. Nature 297, 43–47. Kurz, M.D., Jenkins, W.J., Hart, S.R., Clague, D.A., 1983. Helium isotopic variations in volcanic rocks from Loihi Seamount and the Island of Hawaii. Earth and Planetary Science Letters 66, 388–406. Kurz, M.D., Kenna, T.C., Lassiter, J.C., DePaolo, D.J., 1996. Helium isotopic evolution of Mauna Kea Volcano: first results from the 1 km drill core. Journal of Geophysical Research 101, 11,781–11,791.

Lancelot, Y., Larson, R.L., et al., 1990. Proceedings of the Ocean Drilling Program Initial Reports, vol. 129. Ocean Drilling Program, College Station. 488 pp. Le Bas, M.J., Le Maitre, R.W., Streckeisen, A., Zanettin, B., 1986. A chemical classification of volcanic rocks based on total alkali– silica diagram. Journal of Petrology 27, 745–750. Legendre, C., Maury, R.C., Caroff, M., Guillou, H., Cotten, J., Chauvel, C., Bollinger, C., Hemond, C., Guille, G., Balis, S., Rossi, P., Savanier, D., 2005a. Origin of exceptionally abundant phonolites on Ua Pou Island (Marquesas, French Polynesia); partial melting of basanites followed by crustal contamination. Journal of Petrology 46, 1925–1962. Legendre, C., Maury, R.C., Savanier, D., Cotten, J., Chauvel, C., Hemond, C., Bollinger, C., Guille, G., Blais, S., Rossi, P., 2005b. The origin of intermediate and evolved lavas in the Marquesas archipelago: an example from Nuku Hiva island (French Polynesia). Journal of Volcanology and Geothermal Research 143, 293–317. Liotard, J.M., Barsczus, H.G., Dupuy, C., Dostal, J., 1986. Geochemistry and origin of basaltic lavas from Marquesas Archipelago, French Polynesia. Contributions to Mineralogy and Petrology 92, 260–268. Lupton, J.E., 1983. Terrestrial inert gases: isotope tracer studies and clues to primordial components in the mantle. Annual Reviews of Earth and Planetary Sciences 11, 371–414. McDonough, W.R., Chauvel, C., 1991. Sample contamination explains the Pb isotopic composition of some Rurutu island and Sasha seamount basalts. Earth and Planetary Science Letters 105, 397–404. McNutt, M., Fischer, K.M., 1987. The South Pacific superswell. In: Keating, B., Fryer, P., Batiza, R., Boehlert, G.W. (Eds.), American Geophysical Union Monograph, vol. 43, pp. 25–34. McNutt, M.K., Bonneville, A.A., 2000. A shallow, chemical origin for the Marquesas swell. Geochemistry, Geophysics and Geosystems 1 (1999GC00028). Meibom, A., Anderson, D.L., 2004. The statistical upper mantle assemblage. Earth and Planetary Science Letters 217, 123–139. Meibom, A., Anderson, D.L., Sleep, N.H., Frei, R., Page Chamberlain, C., Hren, M.T., Wooden, J.L., 2003. Are high 3He/4He ratios in oceanic basalts an indicator of deep-mantle plume components? Earth and Planetary Science Letters 208, 197–204. Mukhopadhyay, S., Lassiter, J.C., Farley, K.A., Bogue, S.W., 2003. Geochemistry of Kauai shield-stage lavas; implications for the chemical evolution of the Hawaiian plume. Geochemistry, Geophysics and Geosystems 4 (32 pp.). Natland, J.H., 2003. Capture of helium and other volatiles during the growth of olivine phenocrysts in picritic basalts from the Juan Fernandez Islands. Journal of Petrology 44, 421–456. Parman, S.W., Kurz, M.D., Hart, S.R., Grove, T.L., 2005. Helium solubility in olivine and implications for high 3He/ 4He in ocean island basalts. Nature 437, 1140–1143. Porcelli, D., Ballentine, C.J., 2002. Models for the distribution of terrestrial noble gases and the evolution of the atmosphere. Reviews of Mineralogy and Geochemistry 47, 412–480. Rison, W., Craig, H., 1983. Helium isotopes and mantle volatiles in Loihi Seamount and Hawaiian Islands basalts and xenoliths. Earth and Planetary Science Lettters 66, 407–426. Saal, A.E., Hart, S.R., Shimizu, N., Hauri, E.H., Layne, G.D., 1998. Pb isotopic variability in melt inclusions from oceanic island basalts, Polynesia. Science 282, 1481–1484. Scarsi, P., 2000. Fractional extraction of helium by crushing of olivine and clinopyroxene phenocrysts: effects on the 3He/ 4He measured ratio. Geochimica et Cosmochimica Acta 64, 3751–3762.

P.R. Castillo et al. / Chemical Geology 240 (2007) 205–221 Sheth, H.C., 2003. Beyond the plume hypothesis. Current Science 85, 1520–1581. Staudacher, T., Allegre, C.J., 1989. Noble gases in glass samples from Tahiti: Teahitia, Rocard and Mehetia. Earth and Planetary Science Letters 93, 210–222. Stuart, F.M., Lass-Evans, S., Fitton, J.G., Ellam, R.M., 2003. High 3 He/ 4He ratios in picritic basalts from Baffin Island and the role of a mixed reservoir in mantle plumes. Nature 424, 57–59. Sun, S.-S., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. Magmatism in the Ocean Basins, vol. 42. Geological Society Special Publication, pp. 313–345. Taylor, R.N., Thirlwall, M.F., Murton, B.J., Hilton, D.R., Gee, M.A.M., 1997. Isotopic constraints on the influence of the Icelandic plume. Earth and Planetary Science Letters 148, E1–E8. Todt, W., Cliff, R.A., Hanser, A., Hoffman, A.W., 1996. Evaluation of a 202Pb–205Pb double spike for high-precision lead isotope

221

analysis. In: Basu, A., Hart, S. (Eds.), Earth Processes: Reading the Isotopic Code. American Geophysical Union Geophysics Monograph, vol. 95, pp. 429–437. Vidal, Ph., Dupuy, C., Barsczus, H.G., Chauvel, C., 1987. Heterogeneite's du manteau et origine des basaltes des Marquises (Polyne'sie). Bulletin of the Geological Society of France 4, 633–642. White, W.M., 1985. Sources of oceanic basalts: radiogenic isotope evidence. Geology 13, 115–118. Woodhead, J.D., 1992. Temporal geochemical evolution in oceanic intra-plate volcanics: a case study from the Marquesas (French Polynesia) and comparison with other hotspots. Contributions to Mineralogy and Petrology 111, 458–467. Zindler, A., Hart, S.R., 1986. Chemical geodynamics. Annual Reviews of Earth and Planetary Sciences 14, 493–571.