Geochemical diversity in submarine HIMU basalts ...

Contrib Mineral Petrol DOI 10.1007/s00410-013-0926-x

ORIGINAL PAPER

Geochemical diversity in submarine HIMU basalts from Austral Islands, French Polynesia Takeshi Hanyu • Laure Dosso • Osamu Ishizuka • Kenichiro Tani • Barry B. Hanan • Claudia Adam • Shun’ichi Nakai • Ryoko Senda • Qing Chang • Yoshiyuki Tatsumi

Received: 30 November 2012 / Accepted: 10 July 2013 Ó Springer-Verlag Berlin Heidelberg 2013

Abstract We present the first report of geochemical data for submarine basalts collected by a manned submersible from Rurutu, Tubuai, and Raivavae in the Austral Islands in the South Pacific, where subaerial basalts exhibit HIMU isotopic signatures with highly radiogenic Pb isotopic compositions. With the exception of one sample from Tubuai, the 40Ar/39Ar ages of the submarine basalts show no significant age gaps between the submarine and subaerial basalts, and the major element compositions are indistinguishable at each island. However, the variations in Pb, Sr, Nd, and Hf isotopic compositions in the submarine basalts are much larger than those previously reported in subaerial basalts. The submarine basalts with less-radiogenic Pb and radiogenic Nd and Hf isotopic compositions show systematically lower concentrations in highly Communicated by T L. Grove.

Electronic supplementary material The online version of this article (doi:10.1007/s00410-013-0926-x) contains supplementary material, which is available to authorized users. T. Hanyu (&)  K. Tani  R. Senda  Q. Chang  Y. Tatsumi Institute for Research on Earth Evolution, Japan Agency for Marine-Earth Science and Technology, Yokosuka 237-0061, Japan e-mail: [email protected] L. Dosso Centre National de la Recherche Scientifique, UMR 6538, IFREMER, BP70, 29280 Plouzane´, France O. Ishizuka Institute of Geology and Geoinformation, Geological Survey of Japan/AIST, Tsukuba 305-8567, Japan

incompatible elements than the typical HIMU basalts. These geochemical variations are best explained by a twocomponent mixing process in which the depleted asthenospheric mantle was entrained by the mantle plume from the HIMU reservoir during its upwelling, and the melts from the HIMU reservoir and depleted asthenospheric mantle were then mixed in various proportions. The present and compiled data demonstrate that the HIMU reservoir has a uniquely low 176Hf/177Hf decoupled from 143 Nd/144Nd, suggesting that it was derived from an ancient subducted slab. Moreover, the Nd/Hf ratios of the HIMU basalts and curvilinear Nd–Hf isotopic mixing trend require higher Nd/Hf ratios for the melt from the HIMU reservoir than that from the depleted mantle component. Such elevated Nd/Hf ratios could reflect source enrichment by a subducted slab during reservoir formation. Keywords HIMU  Ocean islands  Submarine basalts  Isotopes  Recycled crust  Mantle reservoir

C. Adam Centro de Geofı´sica, Universidade de E´vora, 7002-554 E´vora, Portugal S. Nakai Earthquake Research Institute, The University of Tokyo, Tokyo 113-0032, Japan Y. Tatsumi Earth and Planetary Sciences, Kobe University, Kobe 657-8501, Japan

B. B. Hanan Department of Geological Sciences, San Diego State University, San Diego, CA 92182-1020, USA

123

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Introduction The South Pacific region is unique from a geodynamical perspective because it has been recognized as the site of a superplume, which is a large-scale mantle upwelling from the base of the Earth’s mantle (McNutt and Judge 1990; Larson 1991). Secondary small-scale plumes generated from the superplume formed several volcanic chains on the moving Pacific plate (Davaille et al. 2002; Courtillot et al. 2003; Tanaka et al. 2009). Previous studies have documented elemental and isotopic variations between hotspots, which suggest the presence of multiple mantle reservoirs in the deep mantle (Vidal et al. 1984; Zindler and Hart 1986; Nakamura and Tatsumoto 1988; Dupuy et al. 1988; Weaver 1991; Chauvel et al. 1992; Hofmann 1997). However, how mantle reservoirs were formed and preserved in the convecting mantle and how partial melts for ocean island basalts (OIB) are fed by mantle plumes derived from these reservoirs are questions that remain a subject of debate. In this study, we focus on one such mantle reservoir, the high-l (HIMU) reservoir, which is characterized by highly radiogenic Pb isotopes. The occurrence of oceanic basalts with pure HIMU signatures is limited to St. Helena in the mid-Atlantic and several volcanoes in the Austral–Cook Islands in the South Pacific (Nakamura and Tatsumoto 1988; Chaffey et al. 1989; Chauvel et al. 1992; Woodhead 1996; Kogiso et al. 1997; Lassiter et al. 2003; Hanyu et al. 2011a; Kawabata et al. 2011). Geochemical data available

thus far have been confined to subaerial basalts; however, the majority of these volcanic edifices are located below sea level. Therefore, studies on submarine rocks from ocean islands are important for gaining a better understanding of the geochemical characteristics of the HIMU reservoir. During the 2006 expedition with R/V Yokosuka (YK0614), we performed bathymetrical mapping, underway geophysical measurements, single-channel seismic profiling, and dives with the Shinkai 6500 manned submersible around Rurutu, Tubuai, and Raivavae in the Austral Islands. In this paper, major element, trace element, and Pb–Sr–Nd–Hf isotopic compositions together with 40 Ar/39Ar ages and petrographical descriptions of the submarine rocks are presented. Direct observation with the manned submersible enabled us to explore the submarine structure of the volcanic edifices. Rock specimens collected from the submarine volcanism provide robust constraints on the source compositions and melting processes in the mantle plume.

Geological settings and geochemical background Rurutu, Tubuai, and Raivavae in the Austral Islands are aligned parallel to the present direction of Pacific plate motion (Fig. 1). The K–Ar ages of the subaerial Tubuai and Raivavae basalts are 7–13 and 5–8 Ma, respectively

Fig. 1 Map of the Austral–Cook chain and bathymetry around Rurutu, Tubuai, and Raivavae. Dive tracks are shown by bold lines annotated with dive numbers

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(Duncan and McDougall 1976; Bellon et al. 1980; Maury et al. 1994). Two distinct stages of volcanic activity have been recognized at Rurutu, including old- and young-Rurutu activity at 11–13 and 1–2 Ma, respectively (Dalrymple et al. 1975; Duncan and McDougall 1976; Turner and Jarrard 1982). Old-Rurutu, Tubuai, and Raivavae show systematic age progression along the island chain, which is consistent with island generation by a fixed mantle plume. The young-Rurutu basalts may be related to a separate mantle plume presently located near the Arago Seamount (Bonneville et al. 2002). The major rock types of the subaerial basalts range from transitional and alkali basalts to basanites (Kogiso et al. 1997). More evolved rocks such as analcites and nephelinites also appear in Tubuai (Caroff et al. 1997). Mineral modes in the basalts vary from aphyric to highly porphyritic (up to 50 %) rocks with coarse-grained olivines and clinopyroxenes. The origin of these minerals is likely cumulative phenocrysts rather than xenocrysts derived from the mantle or oceanic crust because their trace elements are in equilibrium with host lavas together with Sr and Nd isotopes for clinopyroxenes (Hanyu and Nakamura 2000; Jackson et al. 2009; Hanyu et al. 2011a). The chemical zoning in the olivines is weak. In contrast, most of the clinopyroxenes exhibit robust chemical zoning; they have Ti-rich rims and some show oscillatory zoning with Ti-rich lamellae in Ti-poor cores. Because the Ti-rich and Ti-poor domains of the clinopyroxenes exhibit indistinguishable Sr and Nd isotopic compositions, crustal assimilation is unlikely during clinopyroxene crystallization in the magmas (Hanyu and Nakamura 2000; Jackson et al. 2009). The subaerial basalts from old-Rurutu and Tubuai have been recognized to exhibit HIMU signatures with highly radiogenic Pb isotopic compositions such as 206 Pb/204Pb [ 20.5 (Vidal et al. 1984; Nakamura and Tatsumoto 1988; Chauvel et al. 1992; Kogiso et al. 1997; Hanyu et al. 2011a). They define an isotopic trend oblique to the MORB–OIB array in the Sr–Nd and Nd–Hf isotope spaces (Chauvel et al. 1992; Woodhead 1996; Salters and White 1998; Hanyu et al. 2011a). The subaerial youngRurutu basalts show coherently less-radiogenic Pb and radiogenic Nd and Hf isotopic compositions relative to the subaerial old-Rurutu and Tubuai basalts, whereas radiogenic Sr isotopic signature of the young-Rurutu basalts contrasts to the unradiogenic Sr isotopes of the old-Rurutu and Tubuai basalts (Chauvel et al. 1997; Hanyu et al. 2011a). The subaerial Raivavae basalts can be divided into two groups in terms of isotopic compositions. One group exhibits Pb–Sr–Nd–Hf isotope ratios that partly overlap with the old-Rurutu and Tubuai basalts (Lassiter et al. 2003). In contrast, the second group exhibits less-radiogenic Pb isotopes and radiogenic Nd and Hf isotopes.

Results Submarine rocks Submarine rocks were collected by the Shinkai 6500 manned submersible from Rurutu, Tubuai, and Raivavae. The dive tracks were designed to traverse and climb along the submarine ridges where rocky surfaces were likely exposed (Fig. 1). The first dive (6K988) was conducted on the eastern submarine ridge of Rurutu. Two additional dives at Rurutu were undertaken to examine the lower (6K992) and upper (6K993) sections of the submarine ridge extending westward from the island. At Tubuai, two dives (6K989 and 6K990) were conducted to inspect two different submarine ridges north of the island. Dive 6K991 was dedicated to the observation of the northern submarine ridge of Raivavae. We collected lava fragments from sheet flows, pillow lavas, and brecciated lavas. On the basis of outcrop observations, most of these samples were interpreted to be in situ erupted lavas. The basalts range from aphyric basalts to sparsely olivine (±clinopyroxene, plagioclase)– phyric basalts with large phenocrysts with up to 3 mm in size. Relatively fresh rock chips were handpicked, avoiding altered pieces, and used to prepare rock powders for geochemical analyses. Detailed descriptions of the submarine survey, collected rocks, and the geochemical analytical procedures are presented in electronic supplementary material. Major and trace element compositions Major element compositions of the whole-rock samples with a total value higher than 97 wt% are presented in Table S1 in electronic supplementary material and the trace element compositions in Table 1. Samples with low major element totals \97 wt% (16 samples out of 56) tend to have low SiO2 and elevated P2O5 and CaO for a given MgO content relative to the submarine and subaerial samples with total values [97 wt% from the same island, presumably owing to post-eruption alteration. Therefore, they were excluded from further discussion. The major and trace element compositions of the submarine basalts broadly overlap with the olivine and clinopyroxene fractionation trends defined by the subaerial basalts from the same islands (Fig. 2). Rurutu samples show a large variation in alkalinity. 6K988 samples are transitional basalts with relatively high MgO contents (10.0–10.7 wt%). 6K992 and 6K993 samples are alkali basalts with the most evolved major element compositions (2.8–5.4 wt% MgO) in Rurutu. These samples show higher TiO2, Al2O3, and Sr than the basalts with MgO [ 5 wt%, confirming that olivines and clinopyroxenes are major

123

123

39.6

Ce

1.74

5.41

0.834

4.65

0.865

2.29

0.292

1.77

0.251

3.41

1.25

1.43

1.83

0.480

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

Hf

Ta

Pb

Th

U

143

(duplicate)

0.512869

0.703024

Nd/144Nd

0.702820

(duplicate)

87

Sr/86Sr

Initial isotope ratiosb

5.10

Eu

21.6

Sm

Nd

5.00

17.4

La

Pr

87.7

0.095

19.9

126

21.3

285

5.90

6K988R02A Rurutu 22° 25.350 S 151° 17.240 W 1,554 Old stage

Ba

Cs

Nb

Zr

Y

Sr

Rb

Trace elementsa

Depth (m) Remarks

Location Latitude Longitude

Sample ID

Whole rocks

0.512879

0.512862

0.702986

0.702879

0.498

1.79

1.42

1.21

3.33

0.242

1.73

0.282

2.21

0.843

4.52

0.810

5.22

1.70

5.00

21.1

4.88

38.3

17.3

101

0.093

19.6

123

20.4

326

7.07

Rurutu 22° 25.630 S 151° 17.260 W 1,319 Old stage

6K988-R04

0.512886

0.702924

1.51

5.77

3.74

3.74

5.94

0.209

1.52

0.261

2.21

0.902

5.33

1.09

7.88

2.75

9.05

47.2

11.9

100

50.2

318

0.229

60.8

219

22.6

585

19.1

Tubuai 23° 17.150 S 149° 24.640 W 1,764

6K989-R03

1.53

5.85

3.82

4.04

5.97

0.211

1.54

0.261

2.25

0.907

5.39

1.10

7.89

2.77

9.10

47.4

11.9

100

49.9

317

0.231

62.7

217

22.3

583

19.0

(duplicate)

0.512897

0.703104

1.72

6.48

3.61

4.53

6.01

0.205

1.53

0.265

2.29

0.938

5.62

1.15

10.1

3.06

9.63

51.0

12.9

110

54.8

354

0.240

73.4

230

22.3

653

21.6

6K989R06A Tubuai 23° 17.420 S 149° 24.630 W 1,594

Table 1 Trace element and isotope compositions of submarine samples

0.512888

0.702939

1.82

6.86

3.87

4.80

6.30

0.213

1.57

0.273

2.36

0.974

5.83

1.21

8.72

3.08

10.1

53.1

13.6

116

57.8

365

0.261

76.9

242

23.5

651

22.7

Tubuai 23° 17.760 S 149° 24.690 W 1,301

6K989-R10

1.82

6.77

3.94

4.75

6.27

0.210

1.59

0.270

2.39

0.976

5.83

1.21

8.95

3.06

10.1

53.1

13.6

115

57.2

362

0.265

76.3

240

23.0

667

22.9

(duplicate)

0.512976

0.702944

1.51

3.10

2.07

1.55

4.84

0.281

2.03

0.345

2.74

1.07

6.05

1.15

7.79

2.61

7.97

34.8

8.06

62.5

29.5

178

0.030

23.6

189

27.7

597

7.19

6K990R03B Tubuai 23° 17.050 S 149° 29.160 W 1,435

0.512971

0.703066

0.703001

0.778

2.71

1.75

2.08

4.98

0.240

1.73

0.293

2.36

0.932

5.26

0.993

6.77

2.26

6.94

29.9

6.97

54.7

24.7

146

0.159

32.8

198

22.9

466

15.7

Tubuai 23° 17.150 S 149° 29.050 W 1,327

6K990-R05

0.512960

0.702983

2.23

3.56

2.44

2.23

4.86

0.404

2.55

0.399

3.01

1.07

5.71

1.07

7.37

2.55

7.81

36.4

8.66

67.6

33.4

214

0.029

36.1

181

36.0

683

9.10

6K991R01B Raivavae 23° 47.480 S 147° 41.350 W 1,310

0.512972

0.702972

1.83

3.38

2.18

2.97

4.96

0.256

1.74

0.287

2.32

0.895

5.02

0.974

6.77

2.38

7.23

34.0

8.09

64.2

30.4

193

0.060

43.4

181

24.5

637

5.93

6K991R04A Raivavae 23° 47.560 S 147° 41.400 W 1,221

0.512813

0.512812

0.702793

1.44

5.40

3.77

3.53

7.83

0.629

4.11

0.646

4.93

1.76

9.02

1.63

13.0

3.88

12.0

61.4

15.0

111

63.8

300

0.300

55.6

328

58.9

809

14.2

Rurutu 22° 25.140 S 151° 26.940 W 2,053 Old stage

6K992-R01

0.512861

0.702796

0.954

5.71

3.87

2.47

8.31

0.436

3.01

0.489

3.86

1.46

7.92

1.51

10.4

3.65

11.5

57.4

14.2

116

54.8

323

0.117

43.8

321

37.3

872

8.97

6K992R02A Rurutu 22° 25.010 S 151° 26.860 W 1,862 Old stage

Contrib Mineral Petrol

39.850

15.716

20.525 39.968

15.728

20.672

0.282954

Tubuai 23° 17.150 S 149° 24.640 W 1,764

6K989-R03

61.2

7.72

32.8

7.49

2.49

7.64

1.16

6.41

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

0.396

28.1

La

Tm

164

Ba

3.08

0.057

Cs

Er

35.6

Nb

1.17

194

Zr

Ho

28.8

Y

0.376

2.95

1.12

6.03

1.10

7.27

2.36

7.09

31.3

7.37

58.6

27.1

156

0.290

34.7

194

27.7

0.374

2.92

1.11

6.02

1.10

7.22

2.35

7.07

31.3

7.38

58.4

27.1

157

0.294

34.5

193

27.6

446

461

Sr

444

(duplicate) 15.8

15.9

11.8

Rb

6K992-R04B Rurutu 22° 25.050 S 151° 26.570 W 1,813 Young stage

0.282916

0.282923

Rurutu 22° 25.630 S 151° 17.260 W 1,319 Old stage

6K988-R04

6K992-R03B Rurutu 22° 25.010 S 151° 26.600 W 1,882 Young stage

39.799

15.716

20.468

0.282924

6K988R02A Rurutu 22° 25.350 S 151° 17.240 W 1,554 Old stage

Trace elementsa

Sample ID Location Latitude Longitude Depth (m) Remarks

Pb/204Pb

208

Pb/204Pb

Pb/204Pb

207

206

(duplicate)

Hf/177Hf

176

Depth (m) Remarks

Location Latitude Longitude

Sample ID

Whole rocks

Table 1 continued

0.564

4.16

1.46

7.44

1.36

9.29

3.17

9.89

48.7

12.0

97.1

48.9

324

0.417

57.2

290

47.9

810

24.0

0.493

3.79

1.40

7.45

1.40

9.51

3.28

10.3

50.2

12.4

101

48.9

327

0.338

47.7

312

38.6

806

25.5

6K993-R06 Rurutu 22° 25.410 S 151° 25.870 W 1,343 Old stage

40.103

15.736

20.872

0.283001

Tubuai 23° 17.760 S 149° 24.690 W 1,301

6K989-R10

6K992-R06B Rurutu 22° 25.240 S 151° 26.210 W 1,642 Old stage

40.208

15.750

20.969

0.282946

6K989R06A Tubuai 23° 17.420 S 149° 24.630 W 1,594

0.499

3.84

1.40

7.53

1.41

9.39

3.26

10.3

50.2

12.4

100

48.4

324

0.335

54.0

310

38.3

806

25.3

(duplicate)

39.192

15.591

19.504

0.283044

6K990R03B Tubuai 23° 17.050 S 149° 29.160 W 1,435

0.495

3.79

1.41

7.59

1.44

9.75

3.41

10.6

51.9

12.8

105

50.2

334

0.184

57.9

284

37.5

882

13.4

6K993-R11B Rurutu 22° 25.680 S 151° 25.260 W 1,175 Old stage

39.267

15.599

19.589

0.283053

Tubuai 23° 17.150 S 149° 29.050 W 1,327

6K990-R05

39.075

15.577

19.360

0.283055

6K991R04A Raivavae 23° 47.560 S 147° 41.400 W 1,221

0.148

1.36

0.601

3.76

0.765

5.55

1.84

5.89

23.0

4.56

27.2

8.26

2.04

2.02

138

12.9

135

0.112

6K989-R06A-cpx Tubuai 23° 17.420 S 149° 24.630 W 1,594

Clinopyroxenes

39.091

15.599

19.301

0.283062

6K991R01B Raivavae 23° 47.480 S 147° 41.350 W 1,310

0.164

1.51

0.664

4.16

0.851

6.15

2.04

6.48

25.6

5.08

30.5

9.34

1.78

2.24

166

14.8

144

0.105

6K989-R10-cpx Tubuai 23° 17.760 S 149° 24.690 W 1,301

39.749

15.718

20.350

0.282920

Rurutu 22° 25.140 S 151° 26.940 W 2,053 Old stage

6K992-R01

0.317

2.29

0.804

4.13

0.728

4.69

1.46

5.00

25.8

6.92

64.7

36.7

493

1.23

25.3

134

20.6

437

37.0

JB-1a

Standardc

39.963

15.736

20.686

0.282919

6K992R02A Rurutu 22° 25.010 S 151° 26.860 W 1,862 Old stage

Contrib Mineral Petrol

123

123

0.337

5.28

2.34

1.85

2.84

0.597

Lu

Hf

Ta

Pb

Th

U

Sr/86Sr

39.624

15.672

20.105

0.283010

0.512909

Trace element concentrations are in ppm

39.621

15.666

20.130

0.283017

0.512916

0.703284

0.730

2.71

1.68

2.24

5.05

0.324

2.27

6K992-R04B Rurutu 22° 25.050 S 151° 26.570 W 1,813 Young stage

0.729

2.73

1.69

2.24

5.04

0.321

2.28

40.044

15.740

20.793

0.282923

0.512859

0.702896

1.25

4.87

3.50

3.56

7.21

0.557

3.61

6K992-R06B Rurutu 22° 25.240 S 151° 26.210 W 1,642 Old stage

39.989

15.738

20.707

0.282916

0.512859

0.702794

1.40

5.07

3.78

2.67

7.31

0.453

3.09

6K993-R06 Rurutu 22° 25.410 S 151° 25.870 W 1,343 Old stage

1.41

5.11

3.74

3.64

7.48

0.464

3.11

40.017

15.745

20.726

0.282917

0.282917

0.512853

0.702810

1.38

5.33

3.73

3.85

7.66

0.455

3.07

6K993-R11B Rurutu 22° 25.680 S 151° 25.260 W 1,175 Old stage

0.282948

0.512895

0.702829

0.035

0.163

0.069

0.350

5.88

0.116

0.834

6K989-R06A-cpx Tubuai 23° 17.420 S 149° 24.630 W 1,594

Clinopyroxenes

0.282991

0.512914

0.702811

0.033

0.202

0.068

0.416

7.24

0.129

0.933

6K989-R10-cpx Tubuai 23° 17.760 S 149° 24.690 W 1,301

1.65

8.98

6.38

1.63

3.59

0.304

2.06

JB-1a

Standardc

c

Isotope data for standard materials are presented in Table S2 in Electronic Supplementary Material

Initial isotope ratios are calculated with the age of 14.5, 1.2, 7.0, and 5.5 Ma for old-Rurutu, young-Rurutu, Tubuai, and Raivavae samples, respectively. Raw data are presented in Table S2 in Electronic Supplementary Material

b

a

Pb/204Pb

208

Pb/204Pb

Pb/204Pb

207

206

(duplicate)

Hf/177Hf

176

(duplicate)

Nd/144Nd

143

(duplicate)

87

0.703276

2.39

Yb

Initial isotope ratiosb

6K992-R03B Rurutu 22° 25.010 S 151° 26.600 W 1,882 Young stage

Sample ID Location Latitude Longitude Depth (m) Remarks

Table 1 continued

Contrib Mineral Petrol

Contrib Mineral Petrol

(a)

(b)

(c)

(d)

(e)

(f)

(g)

(h)

(i)

Fig. 2 Plots of major element compositions for the submarine basalts. a SiO2 versus total alkali content. b–i MgO versus other major oxides. The red, blue, and green symbols indicate samples from Rurutu (6K988, 6K992, and 6K993), Tubuai (6K989 and 6K990), and

Raivavae (6K991), respectively. The colored fields indicate compositional ranges of the subaerial basalts. The submarine Rurutu basalts with young ages (6K992-R03 and -R04) are highlighted by the oval

fractionation phases, whereas plagioclase and titanomagnetite fractionation is minor (Kogiso et al. 1997). 6K992R03 and -R04 from the western submarine ridge are distinguished from the other Rurutu samples by high CaO and low Al2O3, P2O5, and total alkali contents (Fig. 2), together with low Ba, Sr, Th, Zr, Nb, and Y (not shown), at a given MgO. The submarine samples collected from Tubuai are alkali basalt (6K990) and basanite (6K989), and their major element compositions overlap with those of the subaerial basalts. Rocks collected by a single dive at Raivavae (6K991) are typical alkali basalts (Fig. 2). The major and trace element compositions of these rocks overlap with those of subaerial Raivavae basalts. Primitive mantle-normalized trace element abundance patterns of submarine rocks are broadly similar to those of the subaerial HIMU basalts in terms of a steep slope from

heavy rare earth elements (HREE) to light REE (LREE), high concentrations of incompatible elements, and depletion in Rb, Ba, and Pb relative to the elements with similar incompatibility (Fig. 3) (Dupuy et al. 1988; Weaver 1991; Kogiso et al. 1997; Willbold and Stracke 2006). Note that samples 6K990-R03B, 6K991-R01B, and 6K991-R04A exhibit positive spikes at U and P with small positive anomalies for Y and HREE, which may be attributed to phosphate precipitation during post-eruption alteration (Fodor et al. 1992). The concentrations of large-ion lithophile elements (LILE) and LREE generally decrease in the order, basanites–alkali basalts–transitional basalts (Fig. 3). Although trace element concentrations may have been modified by fractional crystallization, the submarine alkali basalts generally exhibit trace element abundance patterns subparallel to those of the subaerial ones from the same

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Contrib Mineral Petrol

(a)

(b)

(c)

(d)

Fig. 3 Primitive mantle-normalized trace element abundance patterns for the submarine samples from a old-Rurutu b young-Rurutu, c Tubuai, and d Raivavae. The abundance patterns for transitional basalts, alkalic lavas (alkali basalts and hawaiites), and basanites are shown by the lines with triangles, solid circles, and open circles, respectively. The shaded areas indicate the compositional range of

the subaerial alkalic lavas on each island. The ranges of MgO for the compiled subaerial basalts are 4.8–7.6 wt% for old-Rurutu, 4.4–7.1 wt% for young-Rurutu, 6.2–15.9 wt% for Tubuai, and 4.2–13.0 wt% for Raivavae. Note that some submarine basalts have MgO outside the range of compiled subaerial basalts. The primitive mantle values are adopted from McDonough and Sun (1995)

island (Fig. 3). However, it must be stressed that the submarine alkali basalts from Tubuai (6K990) have lower concentrations of highly incompatible elements (e.g., Ba, Th, and Nb) than the subaerial ones, despite similar abundance patterns for moderately and less incompatible elements. This finding is further discussed below.

olivines (\0.1 wt%), indicating that these olivines are phenocrysts. An exception is a single olivine crystal in 6K989-R10 with a Fo content of 88–90 and low CaO (\0.1 wt%); thus, this olivine may be a xenocryst. While chemical zoning is weak in olivines in the subaerial basalts, olivines in the submarine basalts exhibit wider compositional ranges with intricate chemical zoning. Olivines in 6K988 samples show normal zoning, while those in 6K989 samples show reverse zoning in terms of Fo contents, CaO, NiO, and MnO. Fo contents in olivine rims are close to those expected from FeO/MgO ratios in host lavas, assuming Fe/Mg equilibrium between olivines and melts (Fig. S4 in electronic supplementary material). Although

Mineral compositions Compositions of the olivines and clinopyroxenes are shown in Figs. S2 and S3 with some detailed description in electronic supplementary material. Most of the olivines have a CaO content higher than that in typical mantle

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Contrib Mineral Petrol

olivine cores show various Fo contents in each sample, the range is similar among the samples, irrespective of the FeO/MgO ratios in the host lavas. Such variation, which is also observed in St. Helena HIMU basalts, indicates that some olivines are cumulates crystallized from the earlier lavas (Kawabata et al. 2011). Most of the clinopyroxenes in the submarine basalts have Ti-rich rims and some show oscillatory zoning with Ti-rich lamellae in Ti-poor cores. The clinopyroxenes in the submarine basalts exhibit major element compositions overlapping with those in the subaerial basalts. Low Na2O/ TiO2 ratios (\0.5) preclude a lithospheric origin for the clinopyroxenes (Jackson et al. 2009). This is further supported by the clinopyroxene/whole-rock concentration ratios of trace elements with patterns aligned subparallel to the trends predicted by the partition coefficients between the clinopyroxene and melt (Fig. S5 in electronic supplementary material). The Ti-rich and Ti-poor domains in clinopyroxenes exhibit continuous major element trends, reflecting magmatic fractionation during crystallization. Although the clinopyroxene Mg numbers (Mg#) vary within each sample, the range is similar among the samples, irrespective of the FeO/MgO ratios of the host lavas. This again suggests the involvement of cumulative clinopyroxenes in the magmas (Fig. S4 in electronic supplementary material). Because the Sr and Nd isotopic compositions in the clinopyroxenes are indistinguishable from those in the whole rocks, they are likely phenocrysts crystallized from earlier magmas with compositions similar to the host lavas (see below).

whole rocks and negligible in the clinopyroxenes; therefore, the 87Sr/86Sr disequilibrium between the whole rocks and clinopyroxenes cannot be accounted for by radiogenic ingrowth (Table 1 and Table S2 in electronic supplementary material). If crustal contamination modified 87Sr/86Sr in the host lavas after clinopyroxene crystallization, the Nd and Hf isotope ratios would also differ between whole rocks and clinopyroxenes. However, the whole-rock samples and clinopyroxenes do not show systematic differences in Nd and Hf isotopes. A likely explanation for the 87 Sr/86Sr disequilibrium is that seawater-derived radiogenic Sr was partly implanted in the whole rocks of 6K989-R06A and 6K989-R10, while Nd and Hf isotopic compositions were not disturbed because these elements are less mobile during alteration and they are much less concentrated in seawater than Sr (Millet et al. 2008; Hanyu et al. 2011a). A similar test could not be performed for Pb isotopes. However, 206Pb/204Pb in OIB is modified only for 0.2 by seawater alteration (Verma 1992; Nobre Silva et al. 2009; Hanyu et al. 2011a). This modification is significantly smaller than the entire range of 206Pb/204Pb measured in the studied samples. Moreover, the Pb isotopes correlate strongly with the Nd and Hf isotopes, which cannot be explained simply by seawater alteration. Thus, we believe that if a modification of Nd, Hf, and Pb isotope ratios occurred by seawater alteration, such changes were minor compared with the overall isotopic variation observed among the studied samples. The significance of isotopic variation in the submarine basalts is discussed later. Ages

Isotopic compositions The measured and age-corrected isotope ratios are presented in Table S2 (electronic supplementary material) and Table 1, respectively. Post-eruption radiogenic ingrowth is small, but cannot be disregarded when discussing internal isotopic covariations (Jackson et al. 2009; Hanyu et al. 2011a). Therefore, we use age-corrected isotope ratios in the following discussion. The submarine basalts show much wider isotopic ranges than the subaerial basalts in the Pb–Sr–Nd–Hf isotope spaces (Table 1; Fig. 4). The Sr isotopic compositions may have been partly affected by secondary processes such as crustal contamination, seawater alteration, and post-eruption radiogenic ingrowth. These possibilities were assessed by comparing the isotopic compositions of the whole rocks and clinopyroxene separated from the sparsely clinopyroxene–phyric basalts, 6K989-R06A and 6K989-R10 (Fig. 4). Although the whole-rock samples were acid leached prior to analysis, the measured 87Sr/86Sr of the whole rocks was higher than that in the clinopyroxenes by 0.000138–0.000285. The maximum radiogenic ingrowth for 87Sr/86Sr is 0.000014 in the

The 40Ar/39Ar plateau ages were determined for seven samples (Table 2 and Table S3 in electronic supplementary material). With the exception of 6K989-R03, all samples showed flat age spectra in the contiguous gas fractions, yielding plateau ages with small uncertainties (Fig. S1 in electronic supplementary material). Sample 6K989-R03 exhibited a declining age spectrum with progressive Ar release. Because the total age is in good agreement with the plateau age of this sample, the shape of the age spectrum was likely caused by internal redistribution of 39Ar by the recoil effect during irradiation. Therefore, the plateau age of 6K989-R03 is associated with relatively large uncertainty. The submarine Rurutu basalts show a binary distribution of ages at 13.7–14.3 Ma (6K988-R02A, 6K992-R01) and 1.2–1.4 Ma (6K992-R03B, 6K992-R04B), which broadly correspond to the respective old and young activities defined by the subaerial basalts (Fig. 5). However, the submarine basalts of the old activity slightly predate subaerial equivalents. At Tubuai, 6K989-R03 also exceeds the age of the oldest subaerial basalt by more than the two-

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Contrib Mineral Petrol

(a)

(b)

(c)

(d)

(e)

(f)

sigma uncertainty, whereas 6K990-R05 is younger and overlaps the age range of the subaerial basalts. The 40 Ar/39Ar age of 6K991-R04A from Raivavae is within the age range of the subaerial basalts. As a result, the 40Ar/39Ar ages of the submarine basalts document no significant age

123

gaps between the submarine and subaerial basalts, except for 6K989-R03 from Tubuai. The 40Ar/39Ar age of 6K989R03 may reflect the longer period of monotonous lava emplacement at Tubuai than that estimated by the subaerial basalts. However, Adam and Bonneville (2008) suggested

Contrib Mineral Petrol b Fig. 4 Pb–Sr–Nd–Hf isotope diagrams for the submarine and sub-

aerial basalts. The submarine samples from Rurutu, Tubuai, and Raivavae are shown by large circles. Small open circles and fields denote age-corrected isotopic compositions of the subaerial basalts determined by clinopyroxene analysis and conventional whole-rock analysis in previous studies, respectively (Hanyu et al. 2011a and references therein). The color codes are old-Rurutu: red; youngRurutu: orange; Tubuai: blue; Raivavae: green, and Mangaia: black. Isotopic ranges for basalts without HIMU signatures in the Austral– Cook chain are shown by open fields annotated with ‘‘non-HIMU’’ (compilation after Nakamura and Tatsumoto 1988; He´mond et al. 1994; Schiano et al. 2001; Lassiter et al. 2003; Bonneville et al. 2006; Hanyu et al. 2011a; Salters et al. 2011). Isotopic ranges for young basalts from the Arago Seamount (0.23 Ma; Bonneville et al. 2006) are shown for comparison with young-Rurutu lavas. Isotopic ranges for the EPR–PAR basalts and the MORB–OIB array are indicated by brown and yellow fields, respectively (compilation after White et al. 1987; Kingsley et al. 2007; Hamelin et al. 2011; Salters et al. 2011). The HIMU reservoir and the depleted component (local shallow mantle or C) are schematically shown by arrows in isotope spaces. The isotopic range of old seamounts (40–56 Ma) discovered in the Austral Seamount chain is from Bonneville et al. (2006). The isotopic compositions of old seamounts with ages 20–33 Ma overlap those of EPR and PAR basalts; thus, they are not shown in the diagrams. Dashed lines in (a) and (b) define the Northern Hemisphere Reference Line (NHRL; Hart 1984). Blue diamonds in (e) and (f) denote isotopic compositions of clinopyroxene separated from the submarine Tubuai samples (6K989-R06A and 6K989-R10). Elevated 87Sr/86Sr in whole rocks compared with those in clinopyroxenes are attributed to seawater alteration and are schematically shown by blue dashed arrows in (e). The Indian–Pacific MORB dividing line in (f) is adopted from Pearce et al. (2007)

the possibility of two-stage volcanism at Tubuai. The elastic thickness of the lithosphere determined by bathymetry and gravity data is much less than that expected in this region. Such a discrepancy can be resolved by assuming an event of volcanic loading on the lithosphere that occurred prior to the subaerial basalt eruption. The oldest submarine basalt from Tubuai may be a manifestation of such unrecognized early plume activity overlapped by the mantle plume that later formed the subaerial basalts.

Discussion Occurrence of submarine young-Rurutu basalts Previous K–Ar dating has determined that Rurutu Island is composed of basal old-Rurutu and surface-covering youngRurutu basalts (Dalrymple et al. 1975; Duncan and McDougall 1976; Turner and Jarrard 1982). The difference in isotopic compositions between old- and young-Rurutu basalts suggests that the two-stage volcanic episode at Rurutu was caused by distinct mantle plumes (Chauvel et al. 1997; Bonneville et al. 2006). We discovered submarine equivalents of the old- and young-Rurutu basalts defined by subaerial volcanism. The two submarine basalts of 6K992-R03B and 6K992-04B can be recognized as

products of young-Rurutu activity because of their 1.2–1.4 Ma 40Ar/39Ar ages. This theory is supported by the fact that their Sr–Nd–Hf–Pb isotopic compositions are similar to those of the subaerial young-Rurutu and Arago Seamount basalts (Fig. 4; Chauvel et al. 1997; Bonneville et al. 2002). In contrast, other basalts from the western and eastern submarine ridges must be related to old-Rurutu activity on the basis of their ages and isotopic compositions. Samples 6K992-R03B and 6K992-04B were recovered from a mound-like structure, interpreted to be the locus of eruption, on the western submarine ridge that was previously built by old-Rurutu activity. This discovery suggests that the eruption of young-Rurutu basalts was not confined to the center of volcanic edifices. Peripheral eruptions may have occurred when the magmas ascending through the central conduit encountered the level of neutral buoyancy, and subsequently moved laterally (Tilling and Dvorak 1993). Such magmas could have been preferably transferred along an elastically weak path caused by previous magma intrusion during old-Rurutu activity; therefore, young-Rurutu basalts likely migrated through the same path as old-Rurutu basalts, resulting in eruption on old submarine ridges. Despite the similar Pb–Sr–Nd–Hf isotope ratios between submarine and subaerial young-Rurutu rocks, similar major and trace element compositions are not always observed. Subaerial young-Rurutu basalts are highly alkalic and are typically hawaiite and basanite, whereas submarine youngRurutu rocks are alkali basalts with moderate alkalinity (Fig. 2). Moreover, the subaerial young-Rurutu basalts are characterized by higher concentrations of LILE, high-field strength elements (HFSE), and REE than those in the submarine basalts (Fig. 3). These results may reflect various degrees of partial melting from an isotopically homogeneous source (Caroff et al. 1997). Semi-quantitative testing indicates that the trace element compositions of the moderately alkalic submarine basalts can be reproduced by higher degrees of partial melting than the highly alkalic subaerial basalts (Fig. S6 in electronic supplementary material). What is the relationship between the old- and youngRurutu sources? Young-Rurutu basalts show less-radiogenic Pb isotopes and more radiogenic Nd and Hf isotopes than old-Rurutu basalts. However, they plot on the trends defined by the submarine and subaerial basalts from Mangaia, old-Rurutu, Tubuai, and Raivavae in the Pb–Nd– Hf isotope spaces, suggesting that the young-Rurutu basalts involve melts derived from the HIMU reservoir (Chauvel et al. 1997). On the contrary, the young-Rurutu basalts exhibit higher 87Sr/86Sr than old-Rurutu basalts. It is unlikely that seawater alteration is the main cause of elevated 87 Sr/86Sr in the young-Rurutu basalts because both

123

Contrib Mineral Petrol Table 2 Summary of stepwise-heating analysis for Sample ID

40

Ar/39Ar dating

Total age (±1r)a Integrated age (Ma)b

Inv. isochron age (Ma)c

Plateau age (±1r)a 40

36

Ar/ Ar intercept

MSWD

d

Weighted average (Ma)

Inv. isochron age (Ma)c

40

36

Ar/ Ar intercept

MSWD

d

Fraction of 39 Ar (%)

Analysis No.

6K988-R02A

14.5 ± 0.4

14.0 ± 0.6

340 ± 57

0.39

14.3 – 0.3

14.0 ± 0.5

340 ± 57

0.39

6K989-R03

16.4 ± 0.5

13.5 ± 1.3

341 ± 17

1.61

16.7 – 0.6

13.6 ± 1.5

372 ± 34

0.97

69.6

U10015

6K990-R05

7.72 ± 0.08

7.45 ± 0.19

309 ± 5

1.52

7.02 – 0.14

6.7 ± 0.4

538 ± 290

0.54

61.9

U10178

6K991-R04A

5.70 ± 0.15

-30 ± 90

10500 ± 94100

0.62

5.47 – 0.18

6.9 ± 0.3

-120 ± 80

0.22

90.3

U10180

12.97 ± 0.14

12.8 ± 0.4

305 ± 11

2.51

13.66 – 0.21

13.8 ± 0.4

272 ± 77

0.60

52.1

U08369

6K992-R03B

1.43 ± 0.08

1.36 ± 0.14

323 ± 27

1.20

1.44 – 0.10

1.36 ± 0.14

323 ± 27

1.20

100

U08370

6K992-R04B

1.21 ± 0.03

1.17 ± 0.06

302 ± 14

0.89

1.20 – 0.03

1.17 ± 0.06

302 ± 14

0.89

100

U09438

6K992-R01

a

Ages were calculated with kb = 4.962 9 10-10 year-1, ke = 0.581 9 10-10 year-1 and

b

Integrated ages were calculated using sum of the total gas released

c

Inv. isochron age: inverse isochron age

d

MSWD: mean square of weighted deviates

40

100

U10182

K/K = 0.01167 %

lithosphere fertilized by the old plume had a very low concentration of Sr relative to Pb, Nd, and Hf. Alternatively, the isotopic compositions of young-Rurutu may simply reflect those of the recent plume. In this model, both old and young plumes were derived from the HIMU reservoir with small compositional heterogeneity at least in Sr isotopes. Such heterogeneity may be related to the degree of differentiation (i.e., time-integrated Rb/Sr) and the formation ages of the mantle reservoir.

Fig. 5 Ages of islands and seamounts versus distances from Macdonald Seamount in the Austral–Cook chain. Gray bars exhibit the range of K–Ar ages of the subaerial basalts (Dalrymple et al. 1975; Duncan and McDougall 1976; Bellon et al. 1980; Turner and Jarrard 1982; Maury et al. 1994). Circles indicate the 40Ar/39Ar ages of the submarine samples with 2r error bars. The basalts with HIMU isotopic signatures are restricted to the island chain from Raivavae to Mangaia, whereas those with EM isotopic signatures appear along the chains including the Macdonald Seamount, Atiu, and Rarotonga. Note that the submarine basalts with old (filled red circles) and young (open orange circles) ages with different isotopic compositions were discovered at Rurutu

submarine and subaerial basalts show similar 87Sr/86Sr. This fact suggests the presence of a material with radiogenic Sr isotopes in the source of young-Rurutu basalts. Chauvel et al. (1997) suggested that the lithosphere beneath Rurutu was once metasomatized by the mantle plume with HIMU geochemical signatures during the old-Rurutu activity. The recent plume with isotopic compositions equivalent to enriched mantle (EM) basalts (e.g., Atiu; Fig. 5) then ascended to produce the young-Rurutu basalts through interaction with such a lithosphere. However, it is not clear why the isotopic characteristics of EM are observed only in Sr but not in Pb, Nd, or Hf isotopes (Lassiter et al. 2003). A possible explanation is that the recent plume had a very high concentration of Sr or that the

123

Geochemical variations between submarine and subaerial basalts Figure 4 shows the isotopic compositions of several submarine samples from Tubuai and old-Rurutu plotted outside of the isotopic range defined by the subaerial basalts. Notably, the two submarine samples from Tubuai show extremely unradiogenic Pb isotopic compositions close to those of the mid-ocean ridge basalts (MORB) occurring along the East Pacific Rise (EPR) and the Pacific Antarctic Ridge (PAR). This discovery suggests that the apparently uniform isotopic compositions observed in the subaerial old-Rurutu and Tubuai basalts simply reflect biased sampling because most volcanic edifices are under water. In addition, a group of subaerial basalts from Raivavae also exhibit less-radiogenic Pb isotopic compositions similar to the submarine Raivavae basalts collected in this study (Lassiter et al. 2003). Such large geochemical variations could have been a common feature in the magmas that formed these islands. In the Pb isotope diagrams, a linear isotopic trend is defined by both submarine and subaerial basalts of oldRurutu, Tubuai, and Raivavae, together with the subaerial basalts from Mangaia in the same island chain (Fig. 4a, b). This concept is best explained by two-component mixing in variable proportions of the HIMU reservoir contained in

Contrib Mineral Petrol

(a)

(b)

(c)

(d)

Fig. 6 Calculated trace element and isotopic compositions of the lavas with strong and weak HIMU signatures. (a; upper panel) Trace element concentrations of the subaerial Tubuai basalts with radiogenic Pb isotopes are divided by those of the submarine Tubuai basalts with less-radiogenic Pb isotopes, which highlights higher concentrations of Rb, Ba, Th, U, Nb, Ta, and LREE in the former than those in the latter, whereas the concentrations of moderately and less incompatible elements are similar between them. (a; lower panel) It is tested whether trace element compositions of the submarine Tubuai basalts with less-radiogenic Pb isotopes (6K990-R05; corrected for 10 % olivine fractionation; black line) are reproduced by the mixing of subaerial HIMU melt with radiogenic Pb isotopes (average subaerial Tubuai basalts; blue line) and melt from the depleted mantle (DM) component (green line). Element compositions of the melts and partition coefficients used in the calculation are shown in Table 3. Low-degree partial melting (2 % in this model) of DM is required to reproduce a steep trend from HREE to LREE in the mixed melt. The mixed melt (pink line), containing 40 % HIMU melt (blue line) and 60 % DM melt, exhibits incompatible element

concentrations comparable to those of sample 6K990-R05. b– d Results for the mixing calculation for Pb–Nd–Hf isotopes and the Nd/Hf ratio. The mixing components (average subaerial Tubuai lavas and melt from DM) are shown by solid squares. The mixing lines (solid black lines) are annotated with mixing proportions. The isotopic compositions and Nd/Hf of 6K990-R05 are reproduced by mixing 80 % DM melt and 20 % subaerial Tubuai basalts. The dashed black lines denote the extension of the mixing line toward the pure HIMU component. b Open symbols indicate data from previous studies (Dupuy et al. 1989; Woodhead 1996; Hanyu et al. 2011a). Nd/ Hf of the primitive and depleted mantles (dashed red lines) and the oceanic crust (gray field) are after Stracke et al. (2003). The melt from DM (solid square at 206Pb/204Pb = 19) is assumed to be produced by 2 % partial melting of garnet peridotite with Nd/Hf = 3.6 (solid circle; Table 3). c–d The symbols and colored fields are the same as those presented in Fig. 4. The gray straight line in (d) indicates linear regression for age-corrected Nd and Hf isotope data of the submarine and subaerial samples with strong HIMU signatures (143Nd/144Nd \ 0.51295, 176Hf/177Hf \ 0.2830)

an upwelling mantle plume and a component with lessradiogenic Pb isotopes. The two-component mixing model is further supported by isotope systematics among Nd, Hf, Sr, and Pb. The HIMU reservoir should plot on the extension of the mixing trends toward the radiogenic Pb

and less-radiogenic Nd and Hf isotopic side. The component with less-radiogenic Pb isotopes, hereafter referred to as the depleted component, should plot on the other side and be constrained as having 206Pb/204Pb \ 19.3, 143 Nd/144Nd [ 0.5130, and 176Hf/177Hf [ 0.2831 (Fig. 4).

123

Contrib Mineral Petrol Table 3 Geochemical compositions and partition coefficients used for melt mixing model Element

Tubuai subaerial basaltsa

Depleted Mantleb

Bulk partition coefficientc

DM meltd

Mixed melt (X = 0.6)e

Mixed melt (X = 0.8)e

6K990R05f

Rb

20.8

0.088

0.00028

4.4

11.0

7.7

14.1

Ba

260

1.2

0.00005

60

140

100

132

Th

4.69

0.0137

0.0022

0.685

2.29

1.49

2.44

U

1.13

0.0047

0.0022

0.23

0.59

0.41

0.70

Nb

56.0

0.21

0.0015

10.5

28.7

19.6

29.5

La

40.1

0.234

0.0083

10.7

22.4

16.6

22.2

Ce Pb

80.3 1.72

0.772 0.0232

0.011 0.011

32.1 0.96

51.4 1.27

41.8 1.12

49.2 1.57

Sr

514

9.8

0.020

316

395

355

419

Nd

37.0

0.713

0.030

17.6

25.4

21.5

26.9

Zr

188

7.94

0.035

174

180

177

178

Hf

4.72

0.199

0.047

3.46

3.97

3.72

4.48

Sm

6.98

0.27

0.057

4.00

5.19

4.60

6.25

Eu

2.21

0.107

0.087

1.11

1.55

1.33

2.04

Gd

4.47

0.395

0.128

2.88

3.52

3.20

6.09

Dy

4.74

0.531

0.171

2.95

3.67

3.31

4.73

Y

18.9

4.07

0.225

17.5

18.1

17.8

20.6

Er

2.27

0.371

0.250

1.44

1.77

1.61

2.12

Yb

1.78

0.401

0.413

0.96

1.29

1.12

1.56

Lu

0.17

0.063

0.443

0.14

0.15

0.15

0.22

206

21.0

19.0

19.0

20.10

19.62

19.59

143

0.51288 0.28296

0.51302 0.28310

0.51302 0.28310

0.51294 0.28303

0.51297 0.28306

0.51297 0.28305

Pb/204Pb

Nd/144Nd 176 Hf/177Hf a

Average of subaerial basalts with MgO between 10 and 13 wt% as representative of HIMU melt with radiogenic

b

Depleted mantle (DM) compositions from Salters and Stracke (2004)

206

Pb/204Pb

c

Partition coefficients for garnet peridotite with 55 % olivine, 15 % clinopyroxene, 25 % orthopyroxene, and 5 % garnet after Stracke et al. (2003)

d

Assuming 2 % fractional melting of DM

e

Calculated melt composition. X denotes the mixing proportion of the DM melt against the HIMU melt (Tubuai subaerial basalts)

f

6K990-R05 corrected for 10 % olivine fractionation as an example for primary alkali basalt (MgO = 12 wt%) with less-radiogenic 206Pb/204Pb

The isotopic trend defined by Mangaia, old-Rurutu, Tubuai, and Raivavae differs from that defined by other basalts with strong EM1 signatures in the Austral–Cook chain such as Rarotonga, Rapa, and the Macdonald Seamount (Fig. 4). Consequently, the mantle plume that formed the island chain between Mangaia and Raivavae was free from the EM1 reservoir. The depleted component could be the Focal Zone (FOZO; Hart et al. 1992; Stracke et al. 2005) in the lower mantle, a common mantle source (C; Hanan and Graham 1996) in the mid-mantle or a local MORB source in the shallow mantle. Indeed, these materials are difficult to differentiate by radiogenic isotopic compositions. However, the Pb–He isotopic correlation defined by the subaerial basalts suggests that the depleted component has 3 He/4He as low as MORB values, precluding FOZO as a

123

candidate (Hanyu et al. 2011b). This theory is supported by the He isotope ratios of the submarine basalts that exhibit low 3He/4He relative to MORB values (unpublished data; Hanyu et al.). Isotopic compositions of the local shallow mantle may be estimated by old seamount groups occurring in the studied area with ages of 20–33 Ma and 40–56 Ma (Bonneville et al. 2006). However, these seamounts exhibit distinct isotopic compositions from younger (\20 Ma) basalts that plot away from the trend defined by the submarine and subaerial samples in the 206Pb/204 Pb–208Pb/204Pb and 206Pb/204Pb–143Nd/144Nd diagrams, and therefore are not proxies of the depleted component of the studied basalts (Fig. 4b, c). In contrast, the isotopic compositions of the local shallow mantle can be approximated by EPR and PAR basalts. In fact, such basalts exhibit large isotopic variations ranging from the depleted

Contrib Mineral Petrol

MORB mantle toward more enriched compositions. The isotopic variations in the submarine and subaerial basalts can be accounted for by assuming that either the depleted or enriched MORB mantle is a mixing end-member component in the Sr–Nd–Hf isotope spaces (Fig. 4c–f). However, the depleted MORB mantle has low 208Pb/204Pb for a given 206Pb/204Pb (e.g., 206Pb/204Pb \ 18.5) such that it plots below the extension of the linear Pb isotopic trend defined by the submarine and subaerial basalts in Fig. 4b. More likely, the depleted component involved in the HIMU basalts would have the most enriched isotopic compositions within the EPR and PAR fields (e.g., 206 Pb/204Pb [ 19) (Lassiter et al. 2003). Such enriched isotopic compositions are also consistent with C as the depleted component because C is considered to be the midmantle component refertilized by slab subduction. In such a case, C adapted to the South Pacific mantle has slightly more radiogenic Nd and Hf and less-radiogenic Sr isotopes than that originally proposed using global MORB and OIB data (Hanan and Graham 1996). The trace element compositions of basalts with radiogenic and less-radiogenic Pb isotopes likely differ according to the mixing ratios between the two components. To minimize the effect of partial melting, we compared the trace element characteristics of alkalic lavas (alkali basalts and hawaiites) with various Pb isotopic compositions. Indeed, the studied submarine basalts broadly exhibit trace element characteristics similar to those of the subaerial HIMU basalts. However, the submarine alkalic lavas with less-radiogenic Pb isotopes from Tubuai (6K990) and Raivavae (6K991) have lower concentrations of incompatible elements such as Rb, Ba, Th, U, Nb, Ta, and LREE than those in the subaerial alkalic lavas with radiogenic Pb isotopes, whereas the concentrations of moderately and less incompatible elements are similar between them (Fig. 6a). As a result, the basalts with more radiogenic Pb isotopes are characterized by higher concentrations in such incompatible elements, which demonstrates that this unique geochemical feature is derived from the HIMU reservoir (Dupuy et al. 1988; Weaver 1991; Kogiso et al. 1997; Willbold and Stracke 2006). As shown in Fig. 6a, we tested whether the trace element concentrations of the submarine Tubuai basalts (6K990-R05) are reproduced by the addition of melt from the depleted component to the melt from the HIMU reservoir represented by the subaerial Tubuai basalts (Table 3). The trace element concentrations of the depleted component were adopted from the depleted mantle values reported by Salters and Stracke (2004). The melt from the depleted component must be produced by low-degree (\5 %) partial melting with residual garnet; otherwise, the REE concentrations in the mixed melt would be too low.

The low-degree melt from the depleted component exhibits elevated concentrations of highly incompatible elements and LREE, although they are not as high as those of the melt from the HIMU reservoir. As such, the mixed melt, including the melt from the depleted component with a proportion of 60 %, would have trace element concentrations that are comparable to those of the submarine Tubuai basalts with less-radiogenic Pb (Fig. 6a). The isotopic compositions of such basalts can also be reproduced by the mixing model (Fig. 6c–d). The mixing proportion of the melt from the depleted component is estimated to be approximately 80 % by the Pb, Nd, and Hf isotope systematics. Although this estimate is higher than that determined by trace element compositions (60 %; Fig. 6a), such a small discrepancy may be attributed to uncertainties in isotope ratios and element concentrations in the assumed mixing components (Table 3). An important point from the mixing calculation is that the lowdegree melt from the depleted component is predominant in the submarine basalts with less-radiogenic Pb isotopes, whereas it exhibits a minor contribution in forming the basalts with radiogenic Pb isotopes. Involvement of the two components in the HIMU basalts has been previously proposed on the basis of diverse Pb isotopic compositions recorded in melt inclusions hosted by olivine phenocrysts from a single lava flow (Saal et al. 1998; 2005; Yurimoto et al. 2004). Many melt inclusions show Pb isotopic compositions overlapping with or even being more radiogenic than the host lavas with HIMU isotopic signatures. However, the presence of melt inclusions with distinctly less-radiogenic Pb isotopes documents the co-existence of the isotopically depleted melt with melt from the HIMU reservoir during olivine crystallization in the magmas. The quantity of melt from the depleted component is smaller than that from the HIMU reservoir because the former is preserved only in the melt inclusions, whereas all the host lavas display HIMU isotopic signatures. It was suggested that the depleted melt trapped in melt inclusions was derived from the oceanic lithosphere through which the melt from the HIMU reservoir percolated (Saal et al. 2005). Although the Pb isotopic variations observed in submarine basalts in this study are apparently similar to those recorded in melt inclusions, these observations may reflect different mixing phenomena. The aforementioned mixing model demonstrates substantial involvement (60–80 %) of the depleted melt in the submarine basalts with less-radiogenic Pb isotopes. It appears unlikely that such a large amount of lithospheric melt was produced solely by interaction with the percolating melt. As an alternative, the depleted component could be the asthenospheric mantle entrained by an upwelling plume from the HIMU reservoir, and both the melts from the HIMU reservoir and the entrained mantle

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may have been mixed in various proportions beneath the lithosphere. This theory could explain the isotopic variations observed between the whole-rock samples. Subsequently, the upwelling melt interacted with the lithospheric wall rock, the evidence of which is preserved in primary melt inclusions, but not in the host lavas. Enriched HIMU reservoir with subducted oceanic crust Nd and Hf isotopes are reliable geochemical tracers, particularly for the submarine samples, because of their low susceptibility to alteration. The subaerial and submarine samples with strong HIMU characteristics (206Pb/204Pb [ 20.5, 143Nd/144Nd \ 0.51295, 176Hf/177Hf \ 0.2830) from Mangaia, Rurutu, and Tubuai define a steep trend oblique to the MORB–OIB array (Salters and White 1998; Hanyu et al. 2011a) (Fig. 4f). This indicates that the HIMU reservoir has an indigenously low 176Hf/177Hf for a given 143 Nd/144Nd relative to other mantle reservoirs, which suggests ancient fractionation in terms of Lu/Hf and Sm/ Nd during the HIMU reservoir formation. MORB and gabbros in the oceanic crust show lower Lu/Hf and Sm/Nd than those in the MORB source, where Lu/Hf fractionation exerts greater effect on the isotopic evolution than Sm/Nd fractionation (Salters and White 1998). Consequently, 176 Hf/177Hf and 143Nd/144Nd of ancient subducted oceanic crust would evolve such that they plot below the MORB– OIB array in the Nd–Hf isotope space; therefore, this model accounts for the unique Nd and Hf isotopic characteristics of the HIMU reservoir (Salters and White 1998; Stracke et al. 2003; Chauvel et al. 2008; Hanyu et al. 2011a). On the other hand, the linear regression of the submarine and subaerial data with strong HIMU characteristics indicates that this line is too steep to fit all of the data points (gray line in Fig. 6d). Such a steep trend points to the isotopic field of the Indian mantle domain with relatively high 176Hf/177Hf for a given 143Nd/144Nd, whereas the basalts with more radiogenic Nd and Hf characteristics plot well within the field of the Pacific mantle domain (Pearce et al. 2007; Salters et al. 2011). These findings indicate that (1) the depleted component, which is either the local shallow mantle or C, has isotopic characteristics of the Pacific mantle and (2) the previously mentioned steep trend is part of a curvilinear mixing trend between the melts from the HIMU reservoir and the depleted component in the Nd– Hf isotope space. The source of the depleted component could be the mantle material that was entrained by the mantle plume from the HIMU reservoir during its upwelling. Recent studies have demonstrated that lithospheric and asthenospheric mantles contain regional geochemical heterogeneity in their depleted mantle matrices. In the South Pacific,

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enriched material may have been embedded into the lithospheric mantle by ridge–hotspot interaction (Kingsley et al. 2007) or dispersed in the asthenospheric mantle by Pacific superplume activity (Janney et al. 2000). Moreover, the Nd and Hf isotopic variations observed in MORB demonstrate intrinsic isotopic heterogeneity caused by local ancient depletion events with varying amounts of melt extraction and various residual mineral phases (Salters et al. 2011). Whatever the cause of these Nd and Hf isotopic variations and whether the mantle beneath the South Pacific is depleted or enriched, it is characterized by a relatively low 176Hf/177Hf for a given 143Nd/144Nd. If C is the depleted component relevant to the studied basalts, its 176 Hf/177Hf has to be [ 0.2831 with 143Nd/144Nd[0.5130, which displaces the position of ‘‘South Pacific C’’ to more radiogenic Nd isotopic values compared with the range of values deduced from Atlantic MORB and OIB studies (Hanan et al. 2000; Blichert-Toft et al. 2005). Considering its lineage as a mid-mantle portion involving recently subducted oceanic crust (Hanan and Graham 1996), C beneath the South Pacific may have retained its Pacific mantle signature in terms of Nd and Hf isotopes. The curvilinear mixing trend inferred from the submarine and subaerial samples in the Nd–Hf space is consistent with the mixing model shown in Fig. 6d. In this model, Nd/Hf of Tubuai subaerial basalts (7.8) is higher than that of the melt from the depleted component (5.1; Table 3); therefore, mixing of these melts defines a concave-down mixing trajectory with a (Nd/Hf)HIMU melt/(Nd/ Hf)DM melt ratio of approximately 1.5. Extrapolating the hyperbolic mixing line reproduces the lowest 143Nd/144Nd and 176Hf/177Hf of the Mangaia basalts (Fig. 6d). These basalts also exhibit elevated Nd/Hf ranging from 6.8 to 10.8 with an average of 7.9. Despite a large scatter, the basalts with high 206Pb/204Pb tend to have higher Nd/Hf than those with low 206Pb/204Pb (Fig. 6b). Consequently, we suggest that the melt from the HIMU reservoir is characterized by elevated Nd/Hf relative to that of the melt from the depleted component. Such high Nd/Hf ratios have been recognized not only in the HIMU basalts but also in the EM basalts (Fig. 11 in Willbold and Stracke 2006). As noted by Willbold and Stracke (2006), the similarities between these basalts in terms of trace element characteristics including Nd/Hf imply that subducted oceanic crust was commonly involved in the genesis of the HIMU and EM reservoirs. Nd/Hf of the melt from the HIMU reservoir estimated above is significantly higher than that of the subducted oceanic crust. Any constituent of the oceanic crust, including N-MORB, altered MORB, gabbro, and bulk crust, shows a limited Nd/Hf range of 3.5–5.7, which overlaps with those of the depleted and primitive mantle (3.2–4.4; Stracke et al. 2003). Partial melting of the

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recycled eclogitic oceanic crust with residual garnet may facilitate an elevation in Nd/Hf in its melt. However, a 5 % partial melting of the eclogite consisting of 80 % clinopyroxene and 20 % garnet, for example, would increase Nd/Hf by only a factor of 1.3 (partition coefficients after Stracke et al. 2003), which appears insufficient to account for the elevated Nd/Hf in the melt from the HIMU reservoir. Recent petrological and geochemical studies suggest that OIB may have been derived from a fertile peridotite or pyroxenite source rather than eclogite (Sobolev et al. 2007; Day et al. 2009; Gurenko et al. 2009; Hanyu et al. 2011a; Kawabata et al. 2011). Such fertile sources with HIMU geochemical characteristics may have been formed by hybridization of the subducted oceanic crust with the surrounding peridotitic mantle. In such a scenario, elevated Nd/Hf in the HIMU basalts would be ascribed to individual processes, or a combination of them, as is subsequently described. First, the hybrid material could be converted to pyroxenites bearing a higher proportion of garnet than the subducted eclogite. Such pyroxenites can be formed by a reaction of subducted oceanic crust with mantle peridotite (Kogiso and Hirschmann 2006; Sobolev et al. 2007). Residual garnet would enhance Nd/Hf fractionation during the partial melting of the hybrid pyroxenites in an upwelling mantle plume. For example, assuming a source pyroxenite with 50 % clinopyroxene and 50 % garnet, a 5 % partial melt would have higher Nd/Hf than its source by a factor of 1.8. Second, elevated Nd/Hf could be the source characteristics implanted in the HIMU reservoir during hybridization. Hybridization would be most likely triggered by partial melting of the subducted oceanic crust and subsequent mixing with the surrounding mantle in a metasomatic manner (Day et al. 2009; Hanyu et al. 2011a). Although the dry solidus of the oceanic crust is higher than the geotherm at most mantle depths (Yasuda et al. 1994; Hirose and Fei 2002), partial melting of the subducted oceanic crust may occur in the presence of carbon dioxide to form carbonatite melt (Dasgupta et al. 2004; Ohtani et al. 2004; Walter et al. 2008). Natural carbonatite and experimentally produced carbonatite melt are enriched in REE with a remarkably high Nd/Hf (Hoernle et al. 2002; Bizimis et al. 2003; Dalou et al. 2009). Such carbonatite melt would imprint high Nd/Hf together with high Lu/Hf on the hybridized mantle; therefore, the HIMU reservoir, characterized by relatively low 176Hf/177Hf, must have been formed recently in this model (Bizimis et al. 2003). Alternatively, silicate melts might be produced from the subducted oceanic crust in the lowermost mantle, where a sharp thermal gradient crosses the solidus of silicate phases (Ono 2008; Fiquet et al. 2010). Silicate perovskite–melt partitioning (Hirose et al. 2004; Walter et al. 2004; Corgne

et al. 2005) indicates that the partial melt of the oceanic crust with residual Mg–perovskite would have elevated Nd/ Hf because Nd is significantly more incompatible in Mg– perovskite than Hf. Consequently, elevated Nd/Hf together with unique geochemical characteristics may have been imprinted in the mantle material by the metasomatic addition of carbonatite or silicate melts from the subducted oceanic crust during the formation of the HIMU reservoir.

Conclusions We reported the geochemical compositions of the submarine basalts from Rurutu, Tubuai, and Raivavae in the Austral Islands, where the subaerial basalts show robust HIMU signatures. These data in conjunction with previous geochemical data on subaerial basalts provide better constraints for melting processes and source materials in the mantle plume. The major findings of this study are summarized in the following points: 1.

2.

3.

4.

5.

6.

In situ submarine basalts collected by a manned submersible include transitional basalts, alkali basalts, hawaiites, and basanites, which are similar to the subaerial rock types reported from each island. The 40Ar/39Ar ages of the submarine basalts overlap with those of the subaerial basalts with the exception of one sample from Tubuai. At Rurutu, two-stage volcanism previously recognized for the subaerial basalts is also confirmed by the submarine basalts. The variation in Pb–Sr–Nd–Hf isotopic compositions in submarine basalts is significantly wider than that in the subaerial basalts. Such variations are best explained by the two-component mixing of the melts from the HIMU reservoir and the depleted mantle component. The submarine basalts with less-radiogenic Pb isotopes have different trace element compositions from the subaerial basalts with radiogenic Pb isotopes. Comparison of trace element compositions between basalts with radiogenic and less-radiogenic Pb isotopes confirms that high concentrations of Rb, Ba, Th, U, Nb, Ta, and LREE in the HIMU basalts is the unique geochemical feature derived from the HIMU reservoir. The submarine basalts in conjunction with the subaerial basalts confirm that the HIMU reservoir is characterized by low 176Hf/177Hf for a given 143Nd/144Nd, and hence, plots below the MORB–OIB array in the Nd–Hf isotope space. This result is consistent with the HIMU reservoir being formed from an ancient subducted slab. The submarine and subaerial HIMU basalts show elevated Nd/Hf relative to the depleted mantle and its

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melt. This is consistent with a hyperbolic mixing trend defined by these basalts in the Nd–Hf isotope space. Elevated Nd/Hf of the HIMU basalts may reflect the enrichment process of the mantle material by a subducted oceanic crust during formation of the HIMU reservoir. Acknowledgments We thank the crew and marine technicians on the R/V Yokosuka and the operation teams of the submersible Shinkai 6500 of the JAMSTEC Polynesian cruise in 2006. A. Bonneville and D. Suetsugu are acknowledged for their support and encouragement in conducting the research cruise. We thank M. Narui and M. Yamazaki at the International Research Center for Nuclear Materials Science, Institute for Materials Research, Tohoku University, for providing opportunities for neutron irradiation of samples at the JRR3 reactor. The SDSU geochemistry labs acknowledge support from the Keck Foundation and the National Science Foundation. We are grateful to J.-I. Kimura and A. R. L. Nichols for their constructive comments. We would also like to thank two anonymous reviewers and Editor T. L. Grove for their thoughtful comments that helped to improve the manuscript.

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