TRANSPORT OF MATERIALSIN THE DYNAMIC EARTH
October 2 – 6, 2001
Kurayoshi, Japan

Program Committee
Eiji Ito (Chair)
Masaru Kono
Bjorn Mysen
Dave Rubie
Michael Walter

Institute for Study of the Earth’s Interior, Misasa
Institute for Study of the Earth’s Interior, Misasa
Geophysical Laboratory, Washington D.C.
Bayerisches Geoinstitut, Beyreuth
Institute for Study of the Earth’s Interior, Misasa

Local Organizing Committee
Masaru Kono (Chair), Hitoshi Chiba, Eiji Ito, Masami Kanzaki, Tomoo Katsura,
Masako Kawata, Noriaki Kishima, Toshiro Kobayashi, Minoru Kusakabe, Yukari
Shimizu, Michael Walter, Shigeru Yamashita, Akira Yoneda

Sponsors
Ministry of Education, Culture, Sports, Science, and Technology
Tottori Prefecture
Kurayoshi City
Misasa Town
Okayama University
Institute for Study of the Earth’s InteriorKurayoshi City Map

to Yonago
Kurayoshi Sta.
to Tottori, Osaka

N
Tenjin-Gawa
Riv.

Hotel St. Palace

KurayoshiOhashi Br.

1km
TakedaBashi Br.

Old Central City

KurayoshiPark SquareCity Office

to ISEI, Misasa

Kurayoshi Mirai
Chushin(Park Square)
1F

Atrium(Reception, Posters)

Tottori Nijisseiki
Nashi Kinenkan(Pear Museum)ii

Small Hall
(Session Room)Bus Time Table
Shuttle Bus

Hotel −→ Meeting Site
Oct. 3 (Wed) – 5 (Fri): Dept. at 08:30
Meeting Site −→ Hotel
Oct. 2 (Tue): Dept. at 21:30
Oct. 3 (Wed) – 5 (Fri): Dept. at 18:30

Public Bus
Carrier: Hinomaru Bus (Kurayoshi Eki (Station) ←→ Ikuta-Shako)
Nikko Bus (Kurayoshi Eki (Station) ←→ Nishi-Kurayoshi)
Nearest Bus Stop: “Kurayoshi Park Square (倉吉パークスクエア)”
Time Required: about 12 minutes.

Dept.
Arri.
Dept.
Arri.

Kurayoshi Sta. → Park Square
07:43 08:00 08:45 09:40 10:55 12:00
07:55 08:12 08:57 09:52 11:07 12:12
14:10 15:15 16:50 18:10 19:05
14:22 15:23 17:02 18:22 19:13

12:55
13:07

Dept.
Arri.
Dept.
Arri.

Park Square → Kurayoshi Sta.
07:15 09:08 09:38 10:48 11:58 13:18
07:27 09:20 09:50 11:00 12:10 13:30
15:48 16:48 17:38 17:58 18:53 19:08
16:00 17:00 17:50 18:10 19:05 19:20

14:38
14:50

iii

Scientific Program
October 2

Tuesday PM

14:00–19:30

Registration

19:00–21:00

Reception

Kurayoshi Mirai Chushin(Posters can be displayed in this period)
Kurayoshi Mirai Chushin

October 3

Wednesday AM

Session 1: Imaging the Mantle

09:00
09:00–10:00

Convenor
Yoshio Fukao

10:00–10:30
10:30–10:50
10:50–12:00
12:00–13:30

Guy Masters
Coffee break
Discussion
Lunch

Welcome amd introduction
The upper and lower mantle transition region as temporal reservoirs of subducted slabs and rising plumes
Seismological constraints on geochemical reservoirs?

October 3

Wednesday PM

Session 2: Formation and evolution of geochemical reservoirs

13:30–14:30
14:30–15:00

Stan Hart
Marc Hirschmann

Formation and evolution of geochemical mantle domains
A petrologic view of the characteristics of mantle geochemicalreservoirs

15:00–15:20
15:20–16:30
16:30–18:00

Coffee break
Discussion
Poster Session

October 4

Thursday AM

Session 3: High pressure mineral physics

09:00–10:00

Takehiko Yagi

10:00–10:30

Eiji Ito

Developments and problems in laboratory high pressure experiments
Recent developments in high pressure mineral physics experiments using sintered diamond anvils

10:30–10:50
10:50–12:00
12:00–13:30

Coffee break
Discussion
Lunch

October 4

Thurdsay PM

Session 4: Dehydration of subducting slabs and water storage in the mantle

13:30–14:30

Max Schmidt

14:30–15:00
15:00–15:20
15:20–16:30
16:30–18:00

Eiji Ohtani
Coffee break
Discussion
Poster Session

Dehydration vs. melting vs. dissolution: Where and how to
generate a mobile component during subduction
Circulation and Redisribution of Water in the Earth’s Interior(Poster preparation)iv

October 2 – 5, 2001v

October 5

Friday AM

Session 5: Magmatic differentiation and the evolution of the
mantle and core

09:00–10:00
10:00–10:30
10:30–10:50
10:50–12:00
12:00–13:30

Eiichi Takahashi
Bernie Wood
Coffee break
Discussion
Lunch

Melt Generation in the Hawaiian Plume
Mantle composition and conditions of core formation

October 5

Friday PM

Session 6: Future Perspectives

13:30–14:30

Ho-kwang Mao

14:30–15:00
15:00–15:20
15:20–16:30
16:30–18:00
19:00–21:00

Dave Rubie
Coffee break
Discussion
Poster Session
BanquetFuture perspectives on high-pressure experimental geophysics
and geochemistry
Title to be announced

October 6

SaturdayPost-conference field trip

08:00

Departurefrom Kurayoshi St. Palace Hotel

Kurayoshi St. Palace Hotel

Field trip to Daisen Volcano (by chartered bus)
20:00

Returnto the Hotel

Front cover: The National Treasure “Nageire-do” in Misasa Town built in 8th century on a precipitousflank of Mt. Mitoku-san.

Kurayoshi Symposiumvi

Table of Contents

1

2

Map of Kurayoshi Cityi

Bus time tableiii

Scientific Programiv

Abstracts of Papersv

Imaging the mantle
1.1 Yoshio Fukao: The upper and lower mantle transition region as temporal reservoirsof subducted slabs and rising plumes . . . . . . . . . . . . . . . . . . . . . . . . . .
1.2 Guy Masters, Gabi Laske: Seismological constraints on geochemical reservoirs? . .
1.3 Mamoru Kato1, and Hitoshi Kawakatsu2 : Seismological in situ estimation of density
jumps across the transition zone discontinuities beneath Japan . . . . . . . . . . . .
1.4 Hitoshi Kawakatsu: Seismic discontinuities/scatterers in the lower mantle . . . . . .
1.5 Frank D. Stacey: Thermal vs compositional interpretations of lower mantle tomography
1.6 Dapeng Zhao: A new global tomography of the earth: structure of mantle plumes . .

1
3
4
5
6
7
8

Formation and evolution of geochemical reservoirs
9
2.1 Stanley R. Hart: Formation and evolution of geochemical mantle domains . . . . . . 11
2.2 M. Hirschmann1 , T. Kogiso1 , and Daniel J. Frost2 : A petrologic view of the characteristics of mantle geochemical reservoirs . . . . . . . . . . . . . . . . . . . . . . . 13
2.3 Ian H. Campbell: The relationship between continental and oceanic crust formationand the development of the depleted mantle . . . . . . . . . . . . . . . . . . . . . . 15
2.4 P.R. Castillo1∗ , R.U. Solidum1,2, and R.S. Punongbayan2: Origin of high-Nb basalts
from the Sulu Arc, southern Philippines . . . . . . . . . . . . . . . . . . . . . . . . 16
2.5 Takeshi Hanyu∗, Tibor J. Dunai, Gareth R. Davies, Ichiro Kaneoka, Susumu Nohda,
and Kozo Uto: Evidence for distinct primitive mantle sources; a noble gas study of
the Reunion hotspot . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17
2.6 Manuel Moreira: Noble gases and the primitive reservoir . . . . . . . . . . . . . . . 18
2.7 Katsura Kobayashi and Eizo Nakamura: Geochemical evolution of Akagi volcano,northeast Japan: implications for interaction between island arc magma and lower
crust, and generation of isotopically variable magmas. . . . . . . . . . . . . . . . . . 21
2.8 Katsura Kobayashi1 , Takuya Moriguti1, Ryoji Tanaka1 , Kenji Shimizu2, Eizo Nakamura1:
Analytical procedure for the measurement of lithium and boron isotope composition
by High Resolution Secondary Ion Mass Spectrometer and its Application . . . . . . 22
2.9 Takeshi Kuritani and Eizo Nakamura: A simple method for precise isotope analysis
of ng-level Pb for natural rock samples . . . . . . . . . . . . . . . . . . . . . . . . . 23
2.10 Naoko Matsumoto∗ and Masao Kitamura: Chemical compositions of a binary dilutesolid solution during growth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24
2.11 Takuya Moriguti and Eizo Nakamura: Precise Li isotope analysis by thermal ionization mass spectrometry (TIMS) using lithium phosphate as an ion source material andhighly-yield lithium separation from natural samples . . . . . . . . . . . . . . . . . 25
2.12 Takuya Moriguti and Eizo Nakamura: Li-B-Pb isotope systematic study of Izu arc
volcanic rocks: implications for crust-mantle recycling at subduction zone . . . . . . 26

October 2 – 5, 2001vii

2.13 Takuya Moriguti and Eizo Nakamura: Li isotope variation in quaternary basaltic rocks
in northeastern Japan arc . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27
2.14 Chie Sakaguchi and Eizo Nakamura: Chemical and isotopic evolution of the Allendemeteorite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29
2.15 K. Shimizu1∗, E. Nakamura1, and S. Maruyama2 : Petrology, geochemistry and garnet
and pyroxene xenocrysts of Belingwe volcanics, Zimbabwe - indisputable evidence
for crustal contamination . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30
2.16 Hiroyuki Takei, Tetsuya Yokoyama, Akio Makishima and Eizo Nakamura: Formationand suppression of AlF3 during HF digestion of rock samples in Teflon bomb for
precise trace element analyses by ICP-MS and ID-TIMS . . . . . . . . . . . . . . . 32
2.17 Hiroyuki Takei and Eizo Nakamura: Origin of the ore-forming fluids in the Hishikari
gold deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33
2.18 Tetsuya Yokoyama, Akio Makishima, Eizo Nakamura: Evaluation of coprecipitationof incompatible trace elements with fluoride during silicate rock dissolution by aciddigestion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35
2.19 Tetsuya Yokoyama, Akio Makishima, Eizo Nakamura: Separation of thorium anduranium from silicate rock samples using two commercial extraction chromatographic
resins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37
2.20 Tetsuya Yokoyama, Akio Makishima, Eizo Nakamura: Precise analysis of 234 U/238 U
39
ratio using UO+
2 ion with TIMS . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.21 Tetsuya Yokoyama, Eizo Nakamura: Precise determination of ferrous iron in silicate
rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41
2.22 Tetsuya Yokoyama, Katsura Kobayashi, Takeshi Kuritani, Eizo Nakamura: Mantle
metasomatism and rapid melt ascent beneath Miyakejima volcano: evidence from238
U-230 Th disequilibrium . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 43
2.23 Tomohiro Usui1 , Eizo Nakamura1, Katsura Kobayashi1 , Shigenori Maruyama2 andHerwart Helmstaedt3 : Petrology and trace element geochemistry of eclogite xenoliths
from the Colorado Plateau . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 46
2.24 Tomohiro Usui, Katsura Kobayashi and Eizo Nakamura: Preparation of reliable zircon standard for U-Pb dating method by HR-SIMS . . . . . . . . . . . . . . . . . . 47
2.25 Eizo Nakamura and Toshio Nakano: Trace element migration associated with evolution of metamorphic fluid during subduction zone-related metamorphism . . . . . . . 49
2.26 Eizo Nakamura, Masaaki Tanimoto, Takeshi Kuritani and Masako Yoshikawa: Geochemical evolution of the Horoman peridotite complex revised by incorporation oflead isotope systematics: Implications for melt extraction, metasomatism, and compositional layering in the mantle . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50
2.27 Gray E. Bebout1 , and Eizo Nakamura: Record in metamorphic tourmalines of subductionzone devolatilization and boron cycling . . . . . . . . . . . . . . . . . . . . . . . . 51
2.28 Toshio Nakano and Eizo Nakamura: Boron isotope geochemistry of metasedimentaryrocks and tourmalines in a subduction-zone metamorphic suite . . . . . . . . . . . . 53
2.29 Akio Makishima, Mayumi Nakanishi and Eizo Nakamura: Simultaneous determination of Ru, Pd, Re, Os, Ir and Pt contents and 187 Os/188 Os ratios in geological samples
using a new anion exchange chromatographic procedure . . . . . . . . . . . . . . . 54
2.30 Estelle F. Rose, A. Makishima, E. Nakamura: Major and rare earth elements geochemistry of the Icelandic Plume: an island-wide approach for source determination . 55
2.31 Sophie Alves1 , Pierre Schiano2, Bernard Bourdon3, Claude J. Allègre3: Magma formation at Saba Island, Lesser Antilles: a numerical model for mantle wedge-slab
mixing according to osmium data . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57
3 High pressure mineral physics

59Kurayoshi Symposium

viii
3.1
3.2

Takehiko Yagi: Developments and problems in laboratory high pressure experiments
61
Eiji Ito: Recent developments in high pressure mineral physics experiments using
sintered diamond anvils . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61
3.3 Hitoshi Yamada1, Tomoo Katsura1 , Toru Shinmei1, Atsushi Kubo1, Osamu Nishikawa1 ,Maoshuang Song1 , Eiji Ito1, Kenichi Funakoshi2: In situ determination of the olivinemodified spinel phase boundary in the system Mg2 SiO4 - Fe2SiO4 . . . . . . . . . . 61
3.4 Masaki Akaogi1∗ , Akira Tanaka1 , and Eiji Ito2: Garnet-ilmenite-perovskite transitions in the system Mg4 Si4O12 -Mg3Al2 Si3O12 : phase equilibria, calorimetry and application to the mantle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63
3.5 Masanori Matsui: MD simulation of the density and P- and S-wave velocity jumps
across the 410km and 660km seismic discontinuities . . . . . . . . . . . . . . . . . 66
3.6 Kiyoshi Fujino1∗, Nobuyoshi Miyajima1 , Yohei Sasaki1 , Reki Odawara1 , Hisayuki
Ogawa1 , Nagayoshi Sata2 , Takehiko Yagi2 : Mineralogy of the lower mantle by
the combined method of a laser-heated diamond anvil cell experiment and analytical transmission electron microscopy . . . . . . . . . . . . . . . . . . . . . . . . . . 67
3.7 S. Ono1 , E. Ito2: Mineralogy of subducted basaltic crust (MORB) in the lower mantle 69
3.8 Tetsuo Irifune∗, Yuichiro Sueda, Takeshi Sanehira, Hisanobu Naka, Toru Inoue, Yuji
Higo, Takayuki Ueda, Ken-ichi Funakoshi, and Wataru Utsumi: In situ X-ray diffraction measurements for some mantle minerals to 40 GPa using a combination of synchrotron radiation and multianvil apparatus with sintered diamond anvils . . . . . . . 71
3.9 Atsushi Kubo, Eiji Ito, Tomoo Katsura, Toru Shinmei, Hitoshi Yamada, Osamu Nishikawa,
Maoshuang Song, Kenichi Funakoshi: Exploration of β-Fe using sintered diamondanvils . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72
3.10 Satoru Urakawa∗, Masayuki Hasegawa, Junji Yamakawa, Ken-ichi Funakoshi, Wataru
Utsumi: High-pressure phase relationships of FeS . . . . . . . . . . . . . . . . . . . 75
3.11 Yoshitaka Aizawa: Experimental determination of compressional wave velocity ofolivine aggregate at 1 GPa up to 1273 K . . . . . . . . . . . . . . . . . . . . . . . . 76
3.12 N. Mayama∗ , I. Suzuki, T. Katsura, A. Yoneda, I. Ohno: Temperature dependence ofelastic moduli of polycrystalline wadsleyite and ringwoodite . . . . . . . . . . . . . 77
3.13 Takaya Nagai, Koichi Kittaka, Takamitsu Yamanaka: X-ray diffraction study of CaO
at high pressure and temperature: equation of state of CaO . . . . . . . . . . . . . . 78
3.14 Yu Nishihara, Eiichi Takahashi, Tomohiro Iguchi, Keisuke Nakayama, Kyoko Matsukage, Miki Shirasaka and Takumi Kikegawa: Thermal equation of state of an omphacite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 79
3.15 Naohisa Hirao, Tadashi Kondo, Eiji Ohtani, and Takumi Kikegawa: Static compression of Fe-Si alloy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80
3.16 Taku Tsuchiya and Katsuyuki Kawamura: Theoretical estimations of contribution ofelectrons to the thermal expansion of Au and Pt metals . . . . . . . . . . . . . . . . 81
3.17 Masahiro Osako1 , Eiji Ito2 and Akira Yoneda2: Thermal diffusivity and thermal conductivity of olivine and garnet under pressures up to 8 GPa and at temperatures up to
1000 K . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83
3.18 Akira Yoneda: Non-hydrostatic thermodynamics of anisotropic materials: geologicalimplications for pressure solution . . . . . . . . . . . . . . . . . . . . . . . . . . . . 86
3.19 Christian Holzapfel1, D.C. Rubie1 , S. Chakraborty2, F. Langenhorst1 , D.J. Frost1:
Rates of cation diffusion in olivine and silicate perovskite under mantle conditions . 88
3.20 KT Koga, MJ Walter, E Nakamura, and K Kobayashi: Carbon self-diffusion in diamond 90
3.21 Osamu Nishikawa: Stored strains and boundary characters of dynamically recrystallized quartz aggregates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 92
3.22 Masami Kanzaki: In-situ observation of plastic deformation at high pressure and temperature: A preliminary study . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 93

October 2 – 5, 2001
3.23 D. Sakamoto∗, A. Yoshiasa, O. Ohtaka, T. Yamanaka and K. Ota: Conduction mechanism of fayalite and gamma-Fe2SiO4: electrical conductivity jump by olivine-spineltransition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.24 Nobuyoshi Miyajima1, Falko Langenhorst1 , Daniel Frost1, Dave Rubie1 and TakehikoYagi2 : Electron energy-loss spectroscopy of garnet-perovskite high-pressure assemblages . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.25 X. Xue∗, T. Kawamoto, M. Kanzaki: NMR evidence for a new water dissolution
mechanism in depolymerized silicate melts: results for hydrous diopside composition
3.26 X. Xue and M. Kanzaki: Ab initio calculation of the 17 O and 1 H NMR parameters forvarious OH groups: Implications to the speciation and dynamics of dissolved waterin silicate glasses . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.27 T. Tsujimura(1) , X. Xue(2) , M. Kanzaki(2) , A. Kitakaze(1) and Y. Kudoh(1) : The effectof sulfur on silicate network structure in sodium silicate glasses . . . . . . . . . . . .
3.28 Shigeru Yamashita: Water speciation in basaltic melts: constraints from water solubility measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.29 Tomohiro Iguchi∗, Eiichi Takahashi, Yu Nishihara: Pressure dependence of the emf ofthermocouples: Comparison of W-Re and Pt-Rh thermocouples with P-measurements
by in situ X-ray . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.30 M. Song1, K. Aoki2, H. Yamawaki2, H. Fujihisa2 , and M. Sakashita2 : Infrared Observation on Low Temperature Phase Transformations of Ice . . . . . . . . . . . . . . .

ix

95

96
99

100
101
102

103
104

4 Dehydration of subducting slabs and water storage in the mantle
107
4.1 Max W. Schmidt: Dehydration vs. melting vs. dissolution: Where and how to generate a mobile component during subduction . . . . . . . . . . . . . . . . . . . . . . . 109
4.2 Eiji Ohtani: Circulation and Redisribution of Water in the Earth’s Interior . . . . . . 111
4.3 Ganesha Atni Venkatramiah, J.A.K. Tareen: An experimental study on the Mg-Al
granulites and production of granitic melts in K2O-MgO-Al2O3 -SiO2-H2O system . . 112
4.4 Harry W. Green, II: Earthquakes at depth, their enabling mineral reactions, and transport of H2O into the deep mantle . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114
4.5 Dapeng Zhao: Slab dehydration and earthquakes: Seismological evidence . . . . . . 116
4.6 Hikaru Iwamori: Transportation of H2 O and melting in subduction zones . . . . . . 117
4.7 Tomoaki Morishita∗ , Shoji Arai, Akihiro Tamura: Apatite concentration in the Finero
phlogopite-peridotite, Italian Western Alps, as the late product of subduction-related
mantle metasomatism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 118
4.8 Hiroyuki Kagi: High internal pressure retained in fluid inclusions of mantle-derived
materials . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 120
4.9 Tatsuhiko Kawamoto and Kyoko Matsukage: In situ observation of supercritical behaviors between H2O fluids and subduction zone magmas. . . . . . . . . . . . . . . 121
4.10 Robert P. Rapp1 and N. Shimizu: Geochemical Consequences of Deep Subduction of
Oceanic Crust: Insights from Melting Experiments on Hydrous Metabasalt at 3-16 GPa122
4.11 Miki Shirasaka and Eiichi Takahashi: The fate of subducted carbonate within oceanic
crusts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124
4.12 Joseph R. Smyth1∗ , Daniel J. Frost2: An Experimental Study of the Effect of Wateron the 410-km Discontinuity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126
4.13 Toru Inoue§∗ , Yuji Higo†, Takayuki Ueda† , Yasutika Tanimoto† and Tetsuo Irifune§:
The effect of water on the high-pressure phase boundaries in the system Mg2 SiO4Fe2 SiO4 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 128
4.14 G.C.M Richard∗, M. Monnereau: Is the transition zone a water reservoir ? . . . . . . 130
4.15 Motohiko Murakami∗ , Kei Hirose, Hisayoshi Yurimoto, Satoru Nakashima: Solubility of water in lower mantle minerals . . . . . . . . . . . . . . . . . . . . . . . . . . 132

xKurayoshi Symposium
4.16 Tomoo Katsura1 , Kiyoshi Fujita2, and Keishi Shinoda3: Electrical conductivity ofhydrous materials at high pressures and high temperatures. . . . . . . . . . . . . . . 133

5

Magmatic differentiation and the evolution of the mantle and core
135
1
1
1
1
1
5.1 E. Takahashi , K. Nakajima , R. Takeguchi , K. Shinozaki , Ren Zhang-Young , R.
Tanaka2 and E. Nakamura2: Melt Generation in the Hawaiian Plume . . . . . . . . . 137
5.2 B.J. Wood: Mantle composition and conditions of core formation . . . . . . . . . . . 140
5.3 Takeshi Kuritani: Boundary layer fractionation constrained by differential information from the Kutsugata lava flow, Rishiri Volcano, Japan . . . . . . . . . . . . . . . 141
5.4 Ulrich Faul: Constraints on Porosity in Partially Molten Regions in the Upper Mantle
from Permeability Measurements and U-series Modeling . . . . . . . . . . . . . . . 142
5.5 Ryoji Tanaka and Eizo Nakamura: Evidence for slab tear and the related magmatism
beneath the northwestern El Salvador, Central America . . . . . . . . . . . . . . . . 144
5.6 Ryoji Tanaka1 , Eizo Nakamura1, Eiichi Takahashi2 : Geochemical evolution of Koolau
Volcano, Hawaii . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 145
5.7 Masaaki Obata∗, Eiichi Takazawa: A possible discontinuous behavior of melt percolation in partially molten mantle and its implication to compositional layering in
orogenic lherzolites. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146
5.8 Yuki Asahara, Eiji Ohtani: The role of water in generation of komatiite magma . . . 147
5.9 Eizo Nakamura1, Toshihiro Suzuki2, Katsura Kobayashi1 , Akio Makishima1 and Masaki
Akaogi2 : Determination of trace element partitioning between majorite and silicate
melt in the primitive mantle composition at pressure of 15 GPa . . . . . . . . . . . . 148
5.10 E. Ito, A. Kubo, T. Katsura, and M. J. Walter: Melting of peridotite under lower
mantle conditions and its implication to differentiation of a deep magma ocean . . . 150
5.11 M. Walter: Core formation in a reduced magma ocean: new constraints from W, P, Ni
and Co . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 152
5.12 D.C. Rubie1 , C. Holzapfel1, J. Reid1 , S. Fortenfant1, H.J. Melosh2 , B.T. Poe1, D.J.
Frost1 , K. Righter2: Mechanisms of metal transport during formation of the earth’s
core: constraints from the kinetics of metal-silicate reactions . . . . . . . . . . . . . 154
5.13 H. Terasaki1 , T. Kato1, S. Urakawa2, K. Funakoshi3, A. Suzuki4,
K. Sato5 , M. Hasegawa2 : Effect of Light Elements on Viscosity of Liquid Iron-Alloy 156
5.14 Takashi Yoshino, Michael J. Walter and Tomoo Katsura: Critical melt fraction for
percolation of core-forming materials . . . . . . . . . . . . . . . . . . . . . . . . . 158
5.15 Tadahiro Hatakeyama and Masaru Kono: Geomagnetic field model for the last 5 myand conditions at core-mantle boundary . . . . . . . . . . . . . . . . . . . . . . . . 160

6

Future Perspectives
6.1 Ho-kwang Mao: Future perspectives on high-pressure experimental geophysics andgeochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.2 Dave Rubie: Title to be announced . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.3 S. Honda1 and M. Yoshida2: Mantle convection as a tool for understanding the links
among the various geological phenomena . . . . . . . . . . . . . . . . . . . . . . .
6.4 Takamitsu Yamanaka∗, A. Yoshiasa, O. Otaka and N. Nagai: High-pressure crystallography in mineral physics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
6.5 Masaru Kono: Dynamo process in the core under varying energy flow conditions . .
6.6 Eizo Nakamura, Akio Makishima, Takuya Moriguti, Katsura Kobayashi, Chie Sakaguchi, Tetsuya Yokoyama, Ryoji Tanaka, Takeshi Kuritani & Hiroyuki Takei: Comprehensive geochemical analyses of extremely small amounts of terrestrial and extraterrestrial materials for sample-return missions . . . . . . . . . . . . . . . . . . .

161
163
163
163
165
167

169

October 2 – 5, 2001
6.7

xi

Katsura Kobayashi, Tomohiro Usui and Eizo Nakamura: Zircon age dating by dynamic multi-collection method using high resolution secondary ion mass spectrometer 170

List of Participants

172

Author Index

177

October 3 (Wednesday) 08:50 — 12:00

Session 1

Imaging the mantle

Keynote: Yoshio Fukao, Earthquake Research Institute, U. Tokyo
Chair: Guy Masters, Scripps Institution of Oceanography, UCSD

1

Imaging the mantle
1.1

3

The upper and lower mantle transition region as temporal reservoirs of subducted slabs
and rising plumes
Yoshio Fukao
Earthquake Research Institute, Univ. Tokyo

We examined cross-sectional views across typical island arcs for several recent P and S wave
tomographic models. We found that the slabs that started or restarted subduction after the Eocene
tend to be now subhorizontally deflected in the upper and lower mantle transition region at depths
from 400 to 1000km, and that the slabs that started subduction before the Eocene are now descending
through the deep lower mantle without upward continuation to surface plates. These findings suggestthat the pre-Eocene slabs were trapped in the transition region just as in the present day and that theresultant gravitational instability in the Eocene epoch induced a global fall of stagnant slabs into thelower mantle and consequently a global change in plate motion (Eocene plate reorganization).
In order to obtain sharper images of deeply subducted slabs and rising plumes, we added independent datasets in our whole mantle P-wave tomography. Addition of the first arrival time datafrom many Russian observatories has enabled us to obtain sharper images of deeply subducted slabs
beneath the northern Kurile-Kamchatka arc and the Aleutian arc. Beneath the northern KurileKamchatka arc, the subducted slab shows a complex contortion at its leading edge in a depth range
below the 660-km discontinuty but above about 900-km. Beneath the Aleutian arc, the horizontal
slab detached from the presently subducting slab lies just above the 660-km discontinuity. Thus, theaddition of a new dataset has revealed stagnant slabs for the hitherto poorly resolved region. The
upper and lower mantle transition region appears to be a temporal reservoir for subducted slabs.
Addition of the PP-P differential travel times obtained by a cross-correlation technique frombroadband seismograms has improved the resolution significantly for the Hawaiian hot plume, which
has been imaged as an extensive low-velocity zone confined in depth from the uppermost mantle
down to about 900-km depth. The Hawaiian plume appears to have its origin in the upper and lower
mantle transition region. Although the improvement of resolution by the addition of PP-P data is limited for the Pacific and African superplumes, the associated low-velocity anomalies rise all the way
through the mantle to remarkably increase their intensities in the upper mantle, beneath French Polynesia and the Afar Triangle, respectively. There may be some mechanism for these rising plumes tohave their pools in the upper mantle. For the African superplume, the addition of arrival time data of
core-diffracted P-waves has emphasized greatly the low-velocity anomaly just above the core-mantle
boundary, a feature consistent with the recent seismological studies.

Session 1

4
1.2

Seismological constraints on geochemical reservoirs?
Guy Masters, Gabi Laske
Scripps Institute of Oceanography, UCSD

During the past 15 years, our view of the seismic structure of the interior of the Earth has improved
immensely. There is now general consensus on the gross 3D structure of the mantle: the upper mantle
and lowermost mantle are characterized by large amplitude, long-wavelength structure while the midmantle contains low-amplitude slab-like fast features which surround broad regions of slow velocities.
Almost all the global models show some level of decorrelation of structure at or near the 660km
discontinuity consistent with the presence of a weak impediment to flow (such as a phase transitionor viscosity jump) but strict stratification of the mantle at this depth can be ruled out. This is also
consistent with the evidence for slab “ponding” at certain locations in the transition zone.
Evidence for layering deeper in the mantle is extremely tenuous, however, several joint inversionsof global seismic datasets for shear velocity and bulk sound speed suggest that there is a negative
correlation between bulk and shear velocity in the bottom few hundred kilometers of the mantle. This
negative correlation is difficult to explain and is unlikely to be due to subsolidus thermal effects oreven partial melt. One current hypothesis is that the regions of negative correlation (under the midPacific and Africa) correspond to a geochemical reservoir which is complementary to the depleted
mantle above. Iron enrichment in these regions might give rise to the observed anticorrelation. On
the other hand, it is unlikely that such a reservoir would be large enough to satisfy the geochemical
constraints.

Imaging the mantle
1.3

5

Seismological in situ estimation of density jumps across the transition zone
discontinuities beneath Japan
Mamoru Kato1 , and Hitoshi Kawakatsu2

1 School of Earth Sciences, Faculty of Integrated Human Studies, Kyoto University, Kyoto 606-8501 Japan;
2

mkato@gaia.h.kyoto-u.ac.jp,
Earthquake Research Institute, University of Tokyo, Bunkyo, Tokyo 113-0032 Japan;
hitosi@eri.u-tokyo.ac.jp

Density jumps at the transition zone discontinuities in the Earth’s mantle are among the key parameters that control the style of mantle convection. How large they are is particularly of geodynamicalinterest, especially in modeling efficiency of mixing of the mantle. We can, on one hand, measure densities of mantle minerals at high pressure and temperature in laboratory, and seismologicalobservations of density, especially at the transition zone discontinuities, would be valuable assets inmodeling mantle mineralogy. It is, however, difficult to observe in situ density jumps seismologically,
and most previous attempts rely on empirical scaling laws between velocities and density.
We take advantage of recent development of high-density broadband seismic networks in Japan,
and estimate shear velocity and density jumps at the 410-km and 660-km discontinuities in this region. Our estimation is based on purely seismological observations such as seismic reflection andtransmission coefficients at the discontinuities, and is free from any mineralogical models. We firstestimate near-vertical S-to-S reflection coefficients by modeling ScS reverberation waveforms for the
Izu-Japan corridor [Kato et al., 2001]. We then estimate conversion-transmission coefficients for Pto-S conversion using receiver-side converted phases that are observed as later phases to the direct P
phase for teleseismic events. The former are controlled by both the shear velocity and density jumps
at the discontinuities, and the latter are mainly controlled by the shear velocity jump, and by combining these observations we are able to overcome the trade-off among model parameters that is inherent
in such seismological studies. We use AK135 as a reference model, and correct effects of anelastic attenuation with our own estimates of seismic Q. We take into account the associated uncertainties, andestimate acceptable ranges of shear velocity and density jumps at 410-km and 660-km discontinuities[Kato and Kawakatsu, 2001].
Our best estimates for the 660-km discontinuity (∆Vs, ∆ρ : 8.5%, 5.4%) lie slightly more than2σ away from AK135. The density jump at the 670-km discontinuity in PREM appears to be toolarge, and is less favorable to explain our observations. For the 410-km discontinuity (∆Vs, ∆ρ :4.9%, 1.9%), the estimated density jump is smaller than both in PREM and AK135, but is more orless consistent with predictions of high-pressure mineralogical models. Our results do not agree withestimates with PP and SS precursors by Shearer and Flanagan [1999] at either discontinuity, whosemodel has apparently a larger impedance contrast at the 410-km than the 660-km discontinuities.

Session 1

6
1.4

Seismic discontinuities/scatterers in the lower mantle
Hitoshi Kawakatsu

Earthquake Research Institute, University of Tokyo, 1-1-1 Yayoi, Bunkyo-ku, Tokyo 113-0032, Japan; e-mail:hitosi@eri.u-tokyo.ac.jp

Indications of lower mantle discontinuities are known for decades, but still little is known about
their properties, and their origins are enigmatic. For example, in our recent studies, broad-bandseismograms of deep events have been explored with the aid of a novel technique (Vinnik et al.,1998, 2001). In our data there are indications of discontinuities near 860-880 km, 1010-1120 km,
1170-1250 km, 1340-1430 and 1670 - 1800 km depths, and the clearest signals are obtained fromthe discontinuity at a depth of 1200 km in several subduction zones. On the other hand, from theanalyses of short-period array data, the presence of “mid-mantle discontinuities” at depths of 9001100km has been shown (Kawakatsu and Niu, 1994; Niu and Kawakatsu, 1997). The presence ofdipping discontinuities/scatterers in the lower mantle has been also confirmed from short-period arraydata (Kaneshima and Helffrich, 1998, 1999; Castle and Creager, 1999). The emerging results fromthe broad-band and short-period array analyses do not necessarily give a consistent picture, and thesituation seems rather confusing.
In hope for clarifying the situation, results of a systematic analysis of data from the high density (& quality) broadband and short-period arrays recently deployed in Japan will be shown. We
especially focus on the mid- and lower mantle structure beneath the Tonga arc. The questions to beaddressed also include (1) should subducted oceanic crusts in the lower mantle all explain these discontinuities/scatterers, (2) is there any evidence for the postulated deep chemical boundary (Kellogg
et al., 1999)?

References
Castle, J. C and K. C. Creager, 1999. A steeply dipping discontinuity in the lower mantle beneath
Izu-Bonin. J. Geoph. Res., 104, 7279-7292.
Kaneshima, S. and G. Helffrich, 1998. Detection of lower mantle scatterers northeast of the Mariana
subduction zone using short-period array data, J. Geoph. Res., 103, 4825-4838.
Kaneshima, S and G. Helffrich, 1999. Dipping low-velocity layer in the mid-lower mantle: evidence
for geochemical heterogeneity, Science, 283, 1888-1891.
Kawakatsu, H., and F. Niu, 1994. Seismic evidence for a 920-km discontinuity, Nature, 371, 301-305.
Kellogg, L. H., B. H. Hager, and R. D. van der Hilst, 1999, Compositional stratification in the deep
mantle, Science, 283, 1881-1884.
Niu, F., and H. Kawakatsu, 1997, Depth variation of the mid-mantle seismic discontinuity, Geophys.
Res. Lett., 24, 429-432.
Vinnik, L., F. Niu, and H. Kawakatsu, 1998, Broad-band converted phases from midmantle discontinuities, Earth, Planets, and Space (OHP special issue), 50, 987-997.
Vinnik, L., M. Kato, and H. Kawakatsu, 2001, Search for seismic discontinuities in the lower mantle,
Geophys. J. International, in press.

Imaging the mantle
1.5

7

Thermal vs compositional interpretations of lower mantle tomography
Frank D. Stacey

CSIRO Exploration and Mining, PO Box 883, Kenmore, Qld. 4069, Australia f.stacey@cat.csiro.au

Reported variations in lower mantle seismic velocities, Vp and Vs , appear too dependent on geography to be explicable solely as thermal anomalies. Two questions arise: Are the temperature sensitivities of elastic moduli strong enough to admit temperature as the dominant effect? and, if so, canwe distinguish regions where velocity variations may be entirely of thermal origin from those wherecompositional effects are required? These questions are addressed by a new analysis of thermoelastic properties, based on an equation of state constrained to give reliable derivative properties (Stacey
and Isaak, 2001). A linear relationship between elasticities and pressure, which is well observed forthe lower mantle (Stacey, 1992,1995), permits calculation not only of the Grüneisen parameter but
its volume derivative as functions of depth. Then thermodynamic identities suffice to calculate thethermal expansion coefficient and the temperature dependences of elasticities.
The adiabatic Anderson-Grüneisen parameter (deltaS) varies with depth from 2.28+/-0.32 at the
top of the lower mantle to 1.24+/-0.13 at the bottom. The corresponding non-dimensional temperaturedependence of rigidity, epsilon, varies from 6.1+/-0.3 to 5.7+/-0.5. As other authors have noted (e.g.
Isaak et al, 1992), this is constant within the uncertainties of calculation. Volume expansion coefficient decreases from 23.3+/-1.2×10−6 /K to 11.3+/-0.1×10−6 /K. The anelastic contribution to epsilon(noted by Karato,1993), which may be up to 15% of the total epsilon, is included in this calculation and there is no significant anelastic contribution to deltaS. Thus these numbers convert directly
to temperature dependences of seismic velocities. In the depth range 670km to 3290km the total
temperature-dependences vary with depth as
dlnVs /dT=5.9to-2.7 × 10−5 /K
dlnVp/dT=-3.4to-1.1 × 10−5 /K
dlnVs /dlnVp increases with depth from 1.75 to 2.46 with much less uncertainty because the uncertainties in the Vs and Vp variations are correlated.
There is a general similarity to tomographic observations in the calculated dlnVs /dlnVp and its
depth variation, inviting the conclusion that thermal effects dominate, even if some compositional
effect is required, most obviously where there are reports of negatively correlated Vs and “bulk soundvelocity” (for which deltaS<1). However, the gross magnitude of the thermal effect gives trouble. If
we postulate an anomalous spherical volume, of radius 500 km, in the middle of the lower mantle,
with a 0.5% contrast in Vp relative to its surroundings, this requires a 250K temperature contrast andconsequent density contrast of 19kg/m3. Then, equating the buoyancy force to the Stokes flow viscousdrag on its motion (Stacey, 1992, Eq.5.97) and imposing plausible limits on the heat flux implied by
the convective speed, we face three alternative conclusions: (i) mid lower mantle viscosity exceeds1023 Pas, (ii) anomalous volumes with velocity contrast of order 0.5% are very much smaller than a
500km radius sphere, or (iii) compositional effects are more important than thermal effects.
References
D G Isaak et al, 1992. Geophys. Res. Lett. 19:741-744.
S Karato, 1993. Geophys. Res. Lett. 20:1623-1626.
F D Stacey, 1992. Physics of the Earth, 3rd ed., Brookfield Press, Brisbane
F D Stacey, 1995. Phys. Earth Planet. Inter. 89:219-245.
F D Stacey and D G Isaak, 2001. Geophys. J. Int. 146:143-154.

Session 18
1.6

A new global tomography of the earth: structure of mantle plumes
Dapeng Zhao

Geodynamics Research Center, Ehime University, Matsuyama 790-8577, Japan,<zhao@sci.ehime-u.ac.jp>

During the last two decades, seismologists have made continuous efforts to determine the threedimensional (3-D) seismic velocity structure of the mantle. In order to better understand the Earth’s
deep structure and dynamics, I have attempted to develop a new model of whole mantle seismic tomography with a novel approach (Zhao, 2001). I adopted a grid parameterization instead of blocks andspherical harmonic expansions which were used in most of the global tomographic studies. Ray paths
and travel times are computed with an efficient 3-D ray tracing scheme (Zhao et al., 1992). Moreover, the topography of mantle discontinuities at 410 and 660 km depths and the Moho discontinuity(Flanagan and Shearer, 1998; Mooney et al., 1998) are taken into account in the tomographic inversions. The three discontinuities exhibit lateral depth variations of tens of kilometers, which greatly
affects the ray path and travel times, hence their depth changes should be taken into account in the
inversions.
This new approach was applied to a large data set of travel times (P, PP, PcP, pP) which werecompiled by the International Seismological Center (ISC) and were reprocessed by Engdahl et al.
(1998), resulting in a new model of whole mantle P-wave tomography. For the shallow mantle, our
new model contains the general features observed in the previous models: a low-velocity ring aroundthe Pacific Ocean basins and high-velocity anomalies under the old and stable continents in the depth
range of 0-400 km. One significant difference from the previous models is that stronger and wider
high-velocity anomalies are visible in the transition zone depths under the subduction zone regions,which suggests that most of the slab materials are stagnant for a long time in the transition zone before
finally dropping down to the lower mantle. Plume-like slow anomalies are visible under the hotspot
regions in most parts of the mantle. The slow anomalies under hotspots usually do not show a straightpillar shape, but exhibit winding images, which suggests that plumes are not fixed in the mantle but
can be deflected by the mantle flow. As a consequence, hotspots are not really fixed but can wander
on the Earth’s surface, as evidenced by the recent geomagnetic and numeric modeling studies. Widerand more prominent slow anomalies are visible at the core-mantle boundary (CMB) than most of thelower mantle, and there is a good correlation between the distribution of slow anomalies at the CMB
and that of hotspots on the surface, which suggest that most of the mantle plumes under the hotspots
may originate from the CMB. However, there may be some small-scaled, weak plumes originatingfrom the transition zone depths.
References
Engdahl, E.R. et al. (1998) Global teleseismic earthquake relocation with improved travel times andprocedures for depth determination. Bull. Seismol. Soc. Am., 88, 722-743.
Flanagan, M. and P. Shearer (1998) Global mapping of topography on transition zone velocity discontinuities by stacking SS precursors. J. Geophys. Res., 103, 2673-2692.
Mooney, W., G. Laske and G. Master (1998) CRUST 5.1: A global crustal model at 5 X 5. J. Geophys.
Res., 103, 727-747.
Zhao, D., A. Hasegawa and S. Horiuchi (1992) Tomographic Imaging of P and S wave velocity
structure beneath northeastern Japan. J. Geophys. Res., 97, 19909-19928.
Zhao, D. (2001) Seismic structure and origin of hotspots and mantle plumes, Earth Planet. Sci. Lett.,
in press.

October 3 (Wednesday) 13:30 – 16:30

Session 2

Formation and evolution of geochemical reservoirs

Keynote: Stan Hart, Woods Hole Oceanographic Institution
Chair: Marc Hirschmann, University of Minnesota

9Formation and evolution of geochemical reservoirs
2.1

11

Formation and evolution of geochemical mantle domains
Stanley R. Hart
Woods Hole Oceanographic Institution, Woods Hole, MA 02543

I will review our current understanding of the placement, genesis, evolution and lithologic natureof the several known geochemical mantle domains, and the dynamics of melting and melt extraction from these mantle domains. This arena has been recently subjected to several paradigm “warps”:
one is the increasing evidence that mesoscale mantle heterogeneities (veins, pods, layers) may be
important in creating some of the isotopic heterogeneity observed in mantle-derived melts. Thisin turn demands a better understanding of melting and melt transport processes. The second is thegeophysically-based concept that the deepest mantle (> 1700km) may be convectively isolated fromthe overlying mantle by virtue of a compositional restraint on deep mantle buoyancy. This in turndemands a better understanding of the chemical and lithologic make-up of mantle plumes, as they
may provide our most direct sampling of the deep mantle.
300
ENRICHED COMP.
MALUMALU
TA'U

BASALT/PUM

100

30

10

3

1
Cs Rb Ba Th U Nb Ta K La Ce Pb Nd Sr Zr Hf Sm Eu Ti Gd Tb Dy Ho Y Er Yb LuELEMENT

Figure 1. Spidergrams (primitive earth normalized) for basalts from Ta’u (average of 18), Malumalu (averageof the two highest 87 Sr/86 Sr basalts) and a model “Enriched Component.” The Malumalu samples arecorrected for olivine control, so as to have the same MgO as the Ta’u average. The mixing is calculated as
75% Ta’u plus 25% EC equals Malumalu. The enriched component bears no resemblance to any typicaloceanic sediment.

Several of the isotopically-defined domains in the mantle zoo are thought to be a result of lithospheric recycling, and yet we have little understanding of the parentage or lithologic expression ofthis material, or of its melting dynamics. We are pursuing several avenues as a means to confront thisproblem. One is an in-depth study of the Samoa hotspot, which is the EM2 mantle holotype. Samoa
is the most enriched of the EM hotspots (87 Sr/86 Sr ∼ 0.7089), and this distinctive chemistry is conventionally ascribed to recycling of ancient lithosphere containing a terrigenous sediment component.
However, no model has as yet satisfactorily accounted for many of the puzzling geochemical characteristics of EM2 hotspots (e.g. the coexistence of enriched 87 Sr/86Sr, heavy δ 18 O, and high 3 He/4 He).
By comparing the geochemical character of basalts from Ta’u Island, Samoa, which are the leastenriched of Samoan basalts (87 Sr/86 Sr = 0.70464) with the most-enriched material, from Malumalu
seamount (87 Sr/86Sr = 0.70889), we are able to characterize the trace-element patterns of the “enrichment”; this is shown in the spidergram Figure 1 below. What is noteworthy is the very large negative

12

Session 2

Ti anomaly in the enriched component, indeed suggestive of a terrigenous sediment component. However, the lack of Pb, Nb, and Ta anomalies, and the negative Ba anomaly, are difficult to reconcile witha terrigenous pedigree for the enriched component. One can simply assert that chemical processing
during subduction has erased the continental fingerprint. We are exploring magic-free explanations as
well, such as autometasomatism of a mixed lithology of peridotite with mafic enclaves.
The second avenue is a chemical and isotopic study of pre-aggregated melts (as captured in melt
inclusions in phenocrysts) in the expectation that these will shed significant light on the melting
lithologies and melt transport networks involved in hotspot volcanism. Pb isotope studies of melt
inclusions from end-member EM1, EM2 and HIMU hotspots show remarkable diversity on a handspecimen scale. Typically, the isotopic arrays are anchored at one end by compositions comparable
to the bulk basalts; the other ends extend into the interior regions of the mantle isotope tetrahedron,
perhaps to a FOZO-like composition. This behavior is what one might expect of melts derived from amantle with mesoscale mafic enclaves, and is thus supportive of a lithologically diverse mantle containing recycled lithospheric domains. At the same time, it is consistent with a monolithologic peridotitic mantle, with isotopically diverse domains being serviced by a fractal melt-extraction network.
In this case, disparate and remote “drainage basins” can end up with a common area of confluence,
mixing and trapping, in shallow melt plumbing systems.
The debate over whole-mantle versus layered-mantle convection is nearing its 30th birthday.
While the geophysical evidence for whole-mantle mass exchange appears incontrovertible, full accommodation with the geochemical evidence has been a challenging goal.

Formation and evolution of geochemical reservoirs
2.2

13

A petrologic view of the characteristics of mantle geochemical reservoirs
M. Hirschmann1, T. Kogiso1 , and Daniel J. Frost2
1 Univ. Minnesota, Dept. Geol. Geophys. Minneapolis, MN 55455 USA
2 Bayerisches Geoinstitut, Universitaet Bayreuth.

The characteristics of mantle geochemical reservoirs are known chiefly from the geochemistryof oceanic basalts and particularly from the geochemistry of oceanic island basalt (OIB). OIB havereceived considerable attention from analytical geochemists, but relatively little from experiemental
petrologists. This is unfortunate, as some key problems regarding the character and origin of thesereservoirs are as much petrologic as they are geochemical. Some key petrologic questions regarding
the source regions of oceanic islands include:(1) Are the lithologic carriers of exotic mantle components easier to melt than those that carrytypical (depleted) mantle signatures? If so, than much of the variability we see among oceanic basalts
may be related to preferential fusion of “plums” in weak plumes and dilution of heterogeneities byvoluminous partial melting of depleted components in strong plumes and ridges. Additionally, thegeochemical mixing spectra in OIB may be refracted through the melting relations, making it difficult
to discern proportions of materials in OIB sources.
(2) Are there major element (phase equilibria) signatures associated with melting of exotic component source regions? Suggestions for such relationships have generally been attributed to the putative presence of eclogite or pyroxenite, but there are other possibilities, such as the role of enhanced
volatiles.
(3) Can geochemical signatures of small-scale heterogeneities survive the partial melting and melt
extraction process? This is an important question for understanding the Os isotopes of OIB source
regions, as the lithologies principally associated with radiogenic Os, pyroxenites, are low in Os. If
partial melts of pyroxenite mix intimately with surrounding peridotite, we would not expect to see aradiogenic Os signature in basalts derived from these sources. The question of reaction with ambient
peridotite is also important for the preservation of possible major element signatures.
Ongoing research at the University of Minnesota is addressing all 3 of these questions. Here wesummarize some of our results pertaining to question 2. In particular, we focus on the possible originof alkalic lavas similar to those common in the Cook-Austral chain, which produce OIB with a range
of geochemical signatures including extreme HIMU. These lavas are characteristically undersaturated
alkali olivine basalts, basanites, and nephelinites.
A survey of major element characteristics of OIB from the Cook-Austral chain suggests that theparental magma is likely a magnesian alkali olivine basalt or basanite with ∼15 wt.% MgO, 44-45
wt.% SiO2, ∼10 wt.% Al2 O3 , and CaO/Al2 O3 >1. Such liquids have not been produced from partial
melting of peridotite or carbonated peridotite at 3 GPa, though partial melts of garnet peridotite near
5 GPa (Walter, 1998) could plausibly be parental to such liquids. Partial melts of peridotite-basalt
mixtures (Kogiso et al., 1998) also do not produce appropriate liquids.
Our experiments partially melting garnet pyroxenite (Mix1G garnet Mg# = 79) at 1.5-2.5 GPa
produce liquids that are strongly alkalic with notable similarities to Cook-Austral lavas, but in detailthese lavas are too rich in Al2 O3 to be parental to the OIB. New experiments at 5.0 and 7.5 GPa usinga Walker-type multi-anvil apparatus at Bayerisches Geoinstitut indicate that expansion of the garnet
field with increasing pressure drives partial melt compositions to lower Al2 O3 contents, producing
liquids with 8% Al2O3 and quite similar to nominal parental HIMU lavas.
Thus, partial melting of garnet pyroxenite is capable of generating lavas similar to likely HIMU
parents, but the required temperatures are high (1650 ◦C at 5 GPa) and not much different fromthose needed to produce similar liquids from garnet peridotite and more importantly, may not bereconcilable with the potential temperature of the HIMU source. Further investigations, including
documentation of low-melt fractions of anhydrous peridotite and pyroxenite in this pressure range, are14

Session 2

needed. Finally, we note that the garnet pyroxenite (Mix1G) is unlike likely compositions for pristine
recycled oceanic crust. Recycled oceanic crust is thought by many to be the origin of the HIMU
source, but is incapable of producing partial melts similar to the alkalic lavas seen in HIMU localities.
It is unclear whether extended residence in the mantle and associated hybridization could produce
garnet-pyroxenite-like compositions that would have partial melting relations similar to Mix1G.

Formation and evolution of geochemical reservoirs
2.3

15

The relationship between continental and oceanic crust formation and the developmentof the depleted mantle
Ian H. Campbell
Research School of Earth Sciences, Australian National University, Canberra, ACT, 0200, Australia<Ian.Campbell@anu.edu.au>

The Nb/U and Th/U of the primitive mantle are 34 and 4.04 respectively, which compare with
9.7 and 3.96 for the continental crust. Extraction of continental crust from the mantle therefore has a
profound influence on its Nb/U but little influence on its Th/U. Conversely, extraction of MOR-typebasalts lowers the Th/U of the mantle residue but has little influence on its Nb/U. As a consequence,variations in Th/U and Nb/U with Sm/Nd can be used to evaluate the relative importance of continental
and basaltic crust extraction in the formation of the depleted (Sm/Nd enriched) mantle reservoir.
This study evaluates Nb/U, Th/U and Sm/Nd variations in suites of komatiites, picrites and theirassociated basalts, of various ages, in order to determine whether basalt and/or continental crust havebeen extracted from their source region. Emphasis is placed on komatiites and picrites because they
formed at high degrees of partial melting and are expected to have Nb/U, Th/U and Sm/Nd that are
essentially the same as the mantle that melted to produce them. The results show that all of the studied
suites, with the exception of the Barberton, have had both continental crust and basaltic crust extracted
from their mantle source region. The high Sm/Nd of the Gorgona and Munro komatiites require theelevated ratios seen in these suites to be due primarily to extraction of basaltic crust from their source
regions whereas basaltic and continental crust extraction are of sub-equal importance in the source
regions to the Yilgarn and Munro komatiites. The Sm/Nd of modern MORBs lies above the crustal
extraction curve on a plot of Sm/Nd against Nb/U, which requires the upper mantle to have had bothbasaltic and continental crust extracted from it.
It is suggested that the extraction of the basaltic reservoir from the mantle occurs at mid-ocean
ridges and that the basaltic crust, together with its complementary depleted mantle residue, is subducted to the core-mantle boundary. When the two components reach thermal equilibrium with theirsurroundings the lighter depleted component separates from the denser basaltic component. Both canreturn to the upper mantle in plumes, and as the counter flow to subducted slabs, but the lighter depleted component has a shorter residence time in the lower mantle than the denser basaltic component.
If the difference in the recycling times for the basaltic and depleted components is about 1.0 to 1.5
Ga, a basaltic reservoir is created in the lower mantle, equivalent to the amount of basalt that is subducted in 1.0 to 1.5 Ga, and that reservoir is isolated from the upper mantle. It is this reservoir that is
responsible for the Sm/Nd ratio of the upper mantle lying above the trend predicted by extraction ofcontinental crust on the plot of Sm/Nd against Nb/U.

Session 2

16
2.4

Origin of high-Nb basalts from the Sulu Arc, southern Philippines
P.R. Castillo1∗ , R.U. Solidum1,2, and R.S. Punongbayan2

1 Scripps Institution of Oceanography, UCSD, La Jolla, CA 92093-0212, USA

2

Philippine Institute of Volcanology and Seismology, Quezon City, Philippines∗<pcastillo@ucsd.edu>

A hallmark geochemical feature of convergent zone lavas is their distinctive depletion in highfield strength elements (HFSE - e.g., Nb, Hf, Zr, Ta) relative to other trace elements. The origin of
this geochemical feature is of great importance to Earth Science as it bears directly on such global
problems as mass balance between material input and output along convergent margins and sourceof mantle plumes and/or mantle heterogeneities. Interestingly, there are rare high-Nb arc basalts withpronounced enrichment in Nb (i.e., basalts with >20 ppm Nb) and other HFSE. Understanding thepetrogenesis of such high-Nb basalts will lead to a better understanding of the behavior of HFSE inconvergent zone magmatism.
Lavas erupted along the Sulu Arc in southern Philippines range from typical, HFSE-depleted arclavas to HFSE-enriched, high-Nb basalts. Sajona and others (J. Petrol. 37: 1996), using the same
arguments of Defant and Drummond (Nature 347: 1990), propose that the enrichment of HFSE in
high-Nb basalts is connected to the presence of slab-derived melts in the eastern segment of Sulu Arc
and in the nearby island of Mindanao. Sajona and others envision that the initially subducted slab ismelted beneath the Sulu Arc. Slab melt rose through the mantle wedge beneath the arc, but only aportion reached the surface; the remainder of the melt enriched the mantle wedge in HFSE throughthe formation of metasomatic phases such as amphibole and rutile. Later, the metasomatized mantle
was dragged down to deeper portion of the wedge. The breakdown of metasomatic phases produced
the different types of Nb-enriched basalts in the Sulu Arc. Implicit in their model is that the presenceslab-derived melt in sub-arc mantle is one of the keys to a better understanding of the behavior of
HFSE in arc lavas - the production of slab-derived melt is the trigger for the production of high-Nb
basalts.
In this study, we analyze the major and trace element and Sr-Nd-Pb isotopic composition of basalts
from the Sulu Arc with variable HFSE contents, from slightly- to highly-enriched in Nb. Results showthat Nb and other HFSE are generally correlated with many other trace elements, producing binary
mixing arrays. Lavas from off-axis seamounts in the South China Sea basin overlap with or form theextension of the high-HFSE end of the binary array. On the other hand, typical mantle-derived andslab-derived arc lavas from eastern Sulu Arc and Mindanao Island occupy the other, HFSE-depletedend of the array. These observations clearly indicate that it is highly unlikely that the HFSE content of
the high-Nb and Nb-enriched basalts originated from the breakdown of HFSE-enriched metasomatic
phases generated by slab-derived melt. The high-Nb basalts and seamount lavas most likely share acommon mantle source which is distinct from the source of both the typical and slab-derived arc lavas.
Results of previous Sr, Nd and Pb isotope investigations suggest that South China Sea seamount lavas
come from an enriched-mantle source similar to that of ocean island basalts (OIB - e.g., Tu and others,
Chem. Geol. 97: 1992). The Sr, Nd, and Pb isotopic ratios of high-Nb basalts form broad, but distinct
correlations with their HFSE contents and these correlations also define a binary mixing relationshipbetween an enriched-mantle and source of typical arc lavas. Thus our isotope data strengthen ourconclusion that the HFSE content of the high-Nb basalts from the Sulu Arc originate from a distinct
mantle source rather from a slab-derived melt.

Formation and evolution of geochemical reservoirs
2.5

17

Evidence for distinct primitive mantle sources; a noble gas study of the Reunion hotspot

Takeshi Hanyu∗ , Tibor J. Dunai, Gareth R. Davies, Ichiro Kaneoka, Susumu Nohda, and Kozo Uto∗

Earthquake Research Institute, University of Tokyo

He, Ne and Ar isotopes were measured from two volcanoes on Reunion (Piton de la Fournaise,<0.53 Ma; Piton des Neiges, 2 Ma - 0.44 Ma) and three magmatic series from Mauritius (Older, 5-8
Ma; Intermediate, 2-3 Ma; Younger, < 1 Ma). These volcanics record 8 Myr activity of the Reunion
Hotspot. 3 He/4 He ratios of Piton de la Fournaise are between 12.5 and 13.5 Ra. 3 He/4He of Piton
des Neiges are constant regardless of 4 He concentration and sample age, and indistinguishable fromthose of Piton de la Fournaise. The Older Series of Mauritius, corresponding to the shield building
stage, has constant 3 He/4He around 11.5 Ra, slightly lower than the ratio of Reunion. Samples fromboth Reunion volcanoes have higher 20 Ne/22 Ne and 21 Ne/22 Ne than atmospheric value (up to 12.6
and 0.0395, respectively). These data define a linear trend on a Ne three-isotope diagram with a slope
between the Loihi and MORB correlation lines. There is a clear correlation between 20 Ne/22 Ne and40
Ar/36 Ar (up to 4600).
The similarity of 3 He/4 He and Sr-Nd isotope ratios between Piton des Neiges and Piton de laFournaise demonstrate that the volcanoes have had a common homogeneous source over a 2 Myr
period. The Older Series of Mauritius also have 3 He/4He and Sr-Nd ratios comparable to Reunion.
These data suggest that both islands were produced by the same hotspot activity with almost constant3
He/4 He over an 8 Myr period.
In contrast, Intermediate and Younger Series of Mauritius and Rodrigues have 3 He/4He ratios
similar to MORB and Sr and Nd isotope ratios closer to MORB than lavas from Reunion and Older
Series of Mauritius. These Intermediate and Younger Series lavas therefore record a late stage thermal
rejuvenation beneath Mauritius derived from a source that is unrelated to the mantle plume.
The isotopic characteristics of the source of the Reunion magmatism are relatively low 3 He/4 He(13 Ra), an intermediate slope in a Ne-three isotope diagram and relatively radiogenic Sr isotope
ratios. These source characteristics cannot be explained by either crustal contamination or MORBsource mixing with Loihi-type primitive mantle. Thus He-Ne-Ar-Sr-Nd isotopes demonstrate that thisplume source is distinct from the Loihi and Iceland plume sources, clearly showing that the mantle
contains several relatively less-degassed reservoirs and not a single primitive source.
Two possible models can account for the different isotopic signature of Reunion and Hawaii
hotspots; (1) the Reunion source contains more recycled materials than Loihi source and (2) the
Reunion source experienced stronger degassing/differentiation than Loihi source in the early stageof mantle evolution. In both cases a convection mode in the mantle is required that isolates several
less-degassed reservoirs to maintain their distinct isotopic signature in the convectively stirred lower
mantle.

Session 218
2.6

Noble gases and the primitive reservoir
Manuel Moreira

Laboratoire de géochimie et cosmochimie, Institut de Physique du Globe de Paris, 4 Place Jussieu, 75005
Paris cedex 05, France, Email: moreira@ipgp.jussieu.fr

Noble gas systematics in the mantle have been used over the past 20 years to constrain the origin of
the Earth’s atmosphere and the mantle dynamics. The simplest model of mantle structure obtained bynoble gas systematics is a depleted-degassed upper mantle overlying a less depleted/degassed lower
mantle. The origin of this model is the isotopic composition of He, Ne, Ar and Xe in MORB and OIB.
We present here an overview of the noble gas systematics in MORB and in OIB.
All the noble gas radiogenic isotopic ratios (e.g. 4 He/3He, 40 Ar/36 Ar, . . .) are higher in MORB
than in OIB. Three interpretations of such an observation are possible. -1- Higher atmospheric contamination of OIB; -2- The OIB source contains atmospheric-like noble gases ; -3- The OIB source
has lower parent/daughter ratios such as U/He, K/Ar.
One answer to these questions is the neon systematics. Neon has three isotopes: 20, 21 and 22.
In the mantle, only the 21 Ne changes significantively by nuclear reactions. The 20 Ne/22 Ne ratio istherefore constant in the mantle. Analysis of MORB and OIB for neon have shown than the 20 Ne/22 Ne
is different in the mantle and in the atmosphere (Figure 1). The origin of this difference will not be
discussed here but we use this property to constrain the other rare gas isotopic ratios by removing theatmospheric contamination using this neon property. The main difference between MORB and OIB
is the 21 Ne/22 Ne ratio (Figure 1). For similar 20 Ne/22 Ne, the 21 Ne/22 Ne is lower in OIB sources thanin MORB source. A low 21 Ne/22 Ne reflects a low (U+Th)/22 Ne ratio.

14.0
solar
Iceland
Pitcairn
Hawaii
Galapagos

20Ne/22Ne

13.0

MORB

12.0

11.0

10.0
AIR
0.030

0.040

0.050

0.060

0.070

21Ne/22Ne

The mantle has a 20 Ne/22 Ne ratio ranging between the highest measurement (=12.6) and the solar
ratio (=13.8). We will use in the following the solar value as the mantle value. It is possible to correctFormation and evolution of geochemical reservoirs

19

all the isotopic ratios ( 21 Ne/22 Ne, 40 Ar/36 Ar, 129 Xe/130 Xe) from air contamination by extrapolating
the observed mixing trend to this solar ratio (Figure 2). The results of such an approach is that the
OIB source does have lower radiogenic isotopic ratios (40 Ar/36 Ar <10,000;129 Xe/130 Xe<7.2) . This
excludes the higher contamination of OIB samples. Moreover, because the 20 Ne/22 Ne ratio is identical in both MORB and OIB sources, this implies also there is negligeable atmospheric noble gases
subducted in the mantle. Therefore, the OIB sources have lower chemical ratios U/He, K/Ar,. . . We
will call ”primitive reservoir”, this reservoir with low 4 He/3He, 21 Ne/22 Ne and 40 Ar/36 Ar ratios andwith low U/He, U/Ne and K/Ar ratios.

MORB
Iceland

MORB
Iceland
7.5

129Xe/130Xe

40Ar/36Ar

30000

20000

7

6.5

10000

AIR

AIR
10

11

12

13

10

Solar

11

12

13

Solar

20Ne/22Ne

20Ne/22Ne

We exclude a shallow origin for the OIB, therefore, the “primitive” noble gas signature has tobe located below the MORB source. This could be either the whole lower mantle, the D” layer orthe core itself. Wherever the OIB source is, it contains more refractory material (U, Th, K) than thedepleted mantle, source of MORB (as shown by analysis of OIB ). Therefore, the primitive signature
reservoir has to contain a higher content of primitive rare gases (3He, 22 Ne, 36 Ar, 130 Xe) than the
MORB reservoir. If this reservoir is the lower mantle, it has to be less degassed (only 10% less) thanthe degassed mantle. This is the simplest model. If the primitive reservoir is located in the core, toaccount for the high U and Th content of OIB material, subducted oceanic crust has to be involved in
OIB source. However, this material is so uranium rich that it must contain high contents of radiogenic
noble gases. A simple mixing equation between material coming from the core (core) with oceanic
crust (OC) can be written. Using X=R/Ra :XOIBsource = α × XOC + (1 − α) × XCore
Where
α=

0.4

1
Core
core
1 + [4He]
×M[4He]OC
MOC

Here XOC = 0, therefore XOIBsource = (1 − α) × XCore
Core
core
If XCore = 120 (chondrites) and XOIB = 40 (Iceland), therefore, α = 0.7. And [4He]×M=[4He]OC
MOC

20

Session 2

core
Assuming M= 0.1%, and [4He]OC = 10−5 ccSTP/g (1.5Ga old crust), we get [4He]Core =MOC
4 × 10−3 ccSTP/g.
Estimates of the helium content in the primitive mantle is between 10−4 and 10−3 ccSTP/g. This leads
to a silicate/iron partitioning between ≈ 0.1 and ≈ 1 for helium. This value is extremely high whencompared to measured values for moderate pressures.

Similar calculation can be done for a depleted lithospheric mantle (harzburgite) that could be
stored in the lower mantle and will give a similar result. A high 3 He content component is necessary,
which is not compatible with a residual mantle because helium is incompatible.
To conclude, the high U required by the U contents in OIB implies necessarily a high 3He (andother primordial noble gases) content component. The core seems to be an improbable component
because of the helium silicate/iron partionning. The model where the lower mantle is less depleted
and degassed than the upper mantle is still the most reasonable and can best explain the data. This
does not imply the lower mantle is homogeneous and less degassed at the same extent everywhere.
Some parts can be highly degassed, others, where big hotspots are coming from, may be not.

Formation and evolution of geochemical reservoirs

21

2.7 Geochemical evolution of Akagi volcano, northeast Japan: implications for interaction
between island arc magma and lower crust, and generation of isotopically variable magmas.
Katsura Kobayashi and Eizo Nakamura
Pheasant Memorial Laboratory, Institute for Study of the Earth’s Interior, Okayama University at Misasa

Major and trace element, and Sr, Nd and Pb isotopic compositions were determined for wholerock samples from the “isotopically anomalous” Akagi volcano on the volcanic front of the Northeast Japan arc. Sr and Nd isotopic compositions of phenocrysts were also analyzed together withdetermination of major and trace element compositions. Compared with the other volcanoes from thevolcanic front, the whole rock isotope compositions of Akagi show highly enriched characteristics;
87
Sr/86 Sr = 0.7060 to 0.7088, N d = -0.40 to -8.6, and 208 Pb/204 Pb = 38.4 to 38.8. The REE patterns
are characterized by HREE depletions with U-shaped patterns from middle to heavy REE, suggestingthat amphibole fractionation was induced by a reaction between clinopyroxene and H2 O-rich magmain the lower crust. The integrated isotope and trace elements systematics, and tectonic structure beneath Akagi volcano, suggest that the lower crustal assimilation by the H2 O-rich primary magma
could have been affected by the double subduction of Philippine Sea and Pacific oceanic plates. This
double subduction could have supplied larger amounts of water to the magma source region in thewedge mantle than in the case of a single subduction zone.
Significant differences in isotopic compositions are observed among phenocrysts and the coexisting melts. Such isotopic disequilibrium may have been resulted from magma mixing between an
isotopically-depleted aphyric and an enriched porphyritic magma in a shallow magma chamber. The
geochemical characteristics of these end-member magmas were retained in the lower crust, despite
differing extents of lower crustal assimilation by the H2O-rich magmas.

22

Session 2

2.8

Analytical procedure for the measurement of lithium and boron isotope composition by
High Resolution Secondary Ion Mass Spectrometer and its Application
Katsura Kobayashi1 , Takuya Moriguti1, Ryoji Tanaka1 , Kenji Shimizu2, Eizo Nakamura1
1 Pheasant Memorial Labratory, ISEI, Okayama Univ. at Misasa
2 Tokyo Institute of Technology

Lithium (Li) and boron (B) isotopes have been to be powerful tracer to understand complicate
geological processes, such as mantle-crust recycling (e.g., Ishikawa & Nakamura, 1994; Moriguti &Nakamura; 1998). We have developed analytical methods for lithium and boron isotopic composition using a High Resolution Secondary Ion Mass Spectrometer (HR-SIMS), Cameca ims-1270. Our
HR-SIMS are equipped 5 electron multipliers as multi-collection system, which reduces uncertainties
for isotope analyses caused by the instrumental fractuation and time-drift of secondary ion intensities. This apparatus, therefore, is suitable for precise determination of isotopic compositions, such aslithium, boron, lead and oxygen, with high spatial resolution (∼ 10µm) from geological materials.
In order to determine the lithium and boron isotope composition, which is comparable by conventional TIMS technique, the reliable standard to calibrate mass discrimination effect in the apparatus
is much important. Therefore, we prepared five synthetic glass standards under high-pressure condition, which have different isotopic compositions of lithium and boron with same basaltic matrix. Theirlithium and boron isotopic compositions of these glass standards were determined by the conventional
TIMS technique. The concentrations of lithium and boron were 7 and 32 µg/g, respectively.
On the HR-SIMS analysis, 10 ∼ 15 nA (∼10 to 20 µm spot size) of O- critical illuminated (Spot)primary ion beam was adopted for both lithium and boron analyses. The acceleration voltage of the
primary and secondary beam was -13 and 10 kV, respectively. In order to remove the mass interference, such as lithium hydrate and boron hydrate, ∼2000 of mass resolution (M/∆M) has been
applied. Typical sensitivities lithium (7Li+ ) and boron (11 B+ ) on this condition were ∼1000 and ∼50
cps/ppm/nA, respectively. Counting time for each cycle was 10 sec and 55 cycles were accumulated
for typical analyses. Consequently, the lithium and boron isotopic compositions of natural and synthetic basaltic glass, can be analyzed from ∼10µm spots with ∼ 1 and ∼2 ‰ (2σ) of reproducibility,
respectively. Applying this technique to the synthetic glass standards, we obtained liner calibration
carves to correct the instrumental fractionation. We are planning to investigate melt inclusions from
OIB and arc volcanics to understand the material recycling process in the earth.

Formation and evolution of geochemical reservoirs
2.9

23

A simple method for precise isotope analysis of ng-level Pb for natural rock samples
Takeshi Kuritani and Eizo Nakamura
Institute for Study of the Earth’s Interior, Okayama University

Lead isotope compositions of rock samples, as well as their strontium and neodymium data, provide important constraints on dynamic processes of the Earth’s evolution. However, analytical precision and accuracy of lead isotope determination with thermal ionization mass spectrometer (TIMS)
has commonly been worse compared to those of strontium and neodymium. The main reason is thatthe correction for the mass fractionation during measurements is not available for lead isotope analysis
because only 204 Pb is non-radiogenic in the naturally occurring isotopes. To overcome this problem
and to improve analytical precision and accuracy, “double (or triple) spike technique” has been introduced in some laboratories (Hamelin et al., 1985; Woodhead et al., 1995; Todt et al., 1996; Thirlwall,2000). However, it is still difficult for many analysts to handle spike solutions because of their highcost and relatively complicated procedures in the preparation of lead samples and estimation of error
propagation through the fractionation correction.
In this presentation we show a simple technique for precise isotope analysis of lead with TIMS,
including improved procedures of column chromatography, sample loading, and mass spectrometry.
The mass discrimination was corrected by the method of “zero-time correction” (Tuttas and Habfast,
1982; Koide and Nakamura, 1990) followed by the conventional normalization using a pure Pb reference material (NBS 981), instead of the “double spike technique”. Two-stage column chromatographyusing 100 and 10 ml columns enabled to minimize the amounts of impurities in separated lead samples, which dramatically improved reproducibility of mass fractionation during mass spectrometry
using an ion emitter made of a mixture of silicic acid and phosphoric acid. Using the present method,
analytical reproducibility of 0.02 % and 0.05 % was attained for 100 ng and 1 ng of NBS981, andalso 0.02 % and 0.06 % for 100 ng and 1 ng of Pb separated from natural rock samples. Furthermore,
we obtained the reproducibility of 0.1 % for GSJ JP-1 (peridotite), in which the Pb concentration is
as low as 0.09 ppm. The measured isotope compositions of USGS standard rocks AGV-1 and BCR-1,
along with their reproducibility, were comparable to the published values using double spike technique (Woodhead and Hergt, 2000). These observations suggest that our simple technique is highly
reliable in terms of accuracy, as well as precision, in the determination of Pb isotopic compositions of
natural rock samples irrespective of rock chemistry and sample sizes from 1 to 100 ng of Pb.

Session 2

24
2.10

Chemical compositions of a binary dilute solid solution during growth
Naoko Matsumoto∗ and Masao Kitamura∗

Institute for Study of the Earth’s Interior, Okayama University

The partitioning of elements between a solid solution and its coexisting phase is one of the mostimportant phenomena in the chemical evolution of the earth and planets. Such partitioning occurs under both equilibrium and disequilibrium conditions. In general, the chemical composition of a growing
solid solution deviates from the equilibrium. Chemical heterogeneities in a solid solution formed during the growth have been widely used to derive information about crystallization history. However,no theory, other than the Rayleigh fractionation model where the growth kinetics is ignored, has beenpresented so far to predict the chemical composition of a growing solid solution.
The composition of a crystal growing from a mother phase is mainly influenced both by masstransfer in the mother phase and by the kinetics operating at the crystal-mother phase interface. Previous models have stressed the effects of mass transfer accompanied by the assumption of equilibrium
at the interface, mainly because a comprehensive understanding of the effect of interface kinetics islacking.
The incorporation of growth units at kink sites is essential for the interface kinetics [1-3]. Recently
we propose a kink kinetic model [4] that can give both the bulk composition and the incorporating
flux per kink site for a binary ideal solid solution. In the present study, we apply the kinetic model totypical growth surfaces taking into account the effect of mass transfer.
To obtain an analytical solution, we treat the growth of a dilute (A, B) solid solution, whereconcentrations of B are sufficiently lower than those of A. The mass transfer is treated by a volume
diffusion model [5], where growth units are transferred directly to a parallel sequence of growth steps
with high kink densities.
The bulk chemical compositions and the normal growth rates are expressed as functions of thesupersaturation at the edge of the diffusion boundary layer, the frequencies of kink kinetic process, thediffusion coefficients, the thickness of the diffusion boundary layer, and the step-spacing. The relativerole of kink kinetics and volume diffusion, the orientation dependence of chemical compositions andgrowth rates are discussed in terms of these parameters. For example, the phenomenological relations
between chemical compositions and the normal growth rates are clarified: the growth sectors formedby surfaces with higher normal growth rate have chemical compositions more close to its equilibrium
value for a given supersaturation.
The formulation of chemical compositions and the normal growth rates gives a new basis forunderstanding the natural crystallization quantitatively.
References:[1] W. van Erk, J. Crystal Growth 57 (1982) 71.
[2] A.A. Chernov, Sov. Physics Uspekhi 13 (1970) 101.
[3] G.H. Gilmer and K.A. Jackson, in: E. Kaldis and H. J. Scheels (Eds.), 1976 Crystal Growth andMaterials, North-Holland, Amsterdam, 1977, p. 80.
[4] N. Matsumoto, M. Kitamura, J. Crystal Growth 222 (2001) 667.
[5] A.A. Chernov, Soviet Physics Uspekhi 4 (1961) 116.

Formation and evolution of geochemical reservoirs

25

2.11 Precise Li isotope analysis by thermal ionization mass spectrometry (TIMS) using
lithium phosphate as an ion source material and highly-yield lithium separation from natural
samples
Takuya Moriguti and Eizo Nakamura
Pheasant Memorial Laboratory, Institute for Study of the Earth’s Interior, Okayama University at Misasa,
Tottori-ken, 682-0193, Japan

A high-precision Li isotope analysis was developed for determining the Li+ ion emitted fromlithium phosphate as an ion-source material by a Re double-filament ionization method and a spot
loading technique in the thermal ionization mass spectrometry. In this method, Li isotopic fractionation is distinctly less sensitive to the filament temperature than those in the previous methods, andstable and high ion beam intensity of more than 8×10−11 A for 7 Li is obtained (Fig. 1). Furthermore,
the sample preparation is simple and the low ionization temperature (850 ◦ C) required by this method
ensures the analysis of the isotopic composition of small amounts of Li with less influence of Li contamination to the sample, compared with the previous methods. These advantages in determining the7
Li/6 Li ratio result in the in-run precision and the analytical reproducibility of ± 0.05-0.08 ‰ (2σmean) and 0.50 ‰ (RSD), respectively for 100ng of Li, and ± 0.04-0.09 ‰ (2σ mean) and 0.29 ‰(RSD), respectively for 500ng of Li
A Highly-yield lithium separation technique for rock and aqueous samples also has been established. Four separate stages of ion-exchange chromatography were carried out using organic ionexchange resin. An ethanol-HCl solution was used for complete separation of Li from Na at the thirdcolumn stage. Total reagent volume for the entire chemical process was reduced to 42 ml and 33.3 ml
for rock samples and seawater, respectively. The recovery yield and total procedural blank are 99.2
- 99.3 % and 11 pg, respectively. In this method, Rb, Sr, Sm, Nd, La and Ce can be collected after
Li elution in the first column chromatography, then separated by the following specific proceduresfor these elements. Therefore, this method makes possible multi-isotope analysis for Li-poor and restricted small amounts of samples such as meteorites and mantle materials, extending to Li isotope
geochemistry and cosmochemistry.

Session 2

26
2.12

Li-B-Pb isotope systematic study of Izu arc volcanic rocks: implications forcrust-mantle recycling at subduction zone
Takuya Moriguti and Eizo Nakamura

Pheasant Memorial Laboratory, Institute for Study of the Earth’s Interior, Okayama University , Misasa,
Tottori-ken, 682-0193, Japan

Li isotope systematics were applied to island arc volcanics from the Izu arc, Japan to investigate
geochemical processes in subduction zones. Li isotope ratios (δ 7Li) and Li/Y of the arc lavas show
clear across-arc variations, decreasing with increasing depth to the Wadati-Benioff zone (Fig 1a andb). This suggests that the amount of subduction component as a fluid added to the source region decreases with depth. δ 7Li-Y/Li systematics of the arc lavas clearly indicate a simple mixing between
two distinctive chemically homogenous endmembers, a subduction component and the mantle wedge.
Furthermore, Li-B-Pb isotope systematics allow clear discrimination between the relative contribution of altered oceanic crust (AOC), oceanic sediment and mantle wedge to arc lavas, and suggeststhat AOC is the dominant subduction component, whereas the contribution of oceanic sediment isextremely small (< 5 %) (Fig 2a and b). The contrasting physicochemical properties for Li and B
in mineral structures imply that Li may be less likely to migrate from the slab into the overlying
mantle wedge than B. Thus the Li isotopic composition in the Earth’s surface material evolved undernear-surface condition could be more efficiently introduced into the deep mantle through subduction zones than the B isotopic signature, making Li isotopes a powerful geochemical tracer for better
understanding of material recycling between crust and mantle.

Formation and evolution of geochemical reservoirs
2.13

27

Li isotope variation in quaternary basaltic rocks in northeastern Japan arc
Takuya Moriguti and Eizo Nakamura

Pheasant Memorial Laboratory, Institute for Study of the Earth’s Interior, Okayama University , Misasa,
Tottori-ken, 682-0193, Japan

Li isotope systematics show great potential as a geochemical tracer for understanding crust-mantle
recycling because of the following unique physicochemical characteristics of Li: (i) Li commonly
exists as a trace element in rocks and minerals. It is a moderately incompatible element in magmatic
processes such as partial melting and fractional crystallization, with incompatibility similar to that ofyttrium. (ii) Li is highly mobile in fluid-related processes. (iii) Li possesses two stable isotopes, 6Li
and 7 Li, with an extremely large relative mass difference. Large isotopic fractionation of Li develops
during low-temperature reactions associated with water at the Earth’s surface, resulting in a variation
in 7Li/6 Li ratio in natural samples > 30 ‰. In contrast, Li isotopic fractionation is extremely small
when temperatures exceed 350 ◦C, and thus, Li isotope fractionation is expected to be insignificant inmagmatic processes.

In the present study, Li isotope systematics were applied to northeastern Japan arc basalts in orderto understand material recycling between crust and mantle at subduction zones, together with Pb
isotope systematics. Pb/Ce and B/Nb ratios of arc basalts decrease gradually with increasing depth
to the Wadati-Benioff zone (WBZ). However, Li/Y ratios do not show gradual decrease (Fig.1). Li/Y
ratios decreases from 134 to 145km depths to the Wadati-Benioff zone but a positive anomaly in thisratio was found at around 180km depth to the WBZ. This across-arc profile of Li/Y ratio may reflect
phase changes in the subducting slab, and the breakdown of lawsonite in the slab at ∼180km depth
could be one explanation. In the lawsonite breakdown reaction (lawsonite + diopside + garnet1 =garnet2 + coesite + H2O) [1], Y is strongly partitioned into garnet2 and Li is strongly partitioned into
H2O. As a result of this reaction, therefore, the Li/Y ratio of the slab derived fluid could be higher
at around 180 km depth, causing the positive anomaly in the Li/Y ratio of arc basalts. Sr, Nd, and
Pb isotope ratios show clear across arc variation in the northeastern Japan arc lavas [2]. However, theLi isotope ratios do not show obvious across-arc variations (Fig.2). This is because the Li isotopic
compositions of mixing two endmembers for arc magmas, mantle wedge and slab-derived fluids,

28

Session 2

are extremely similar (Fig.3). The estimated Li isotopic compositions of the mantle wedge and thefluid estimated using Li and Pb isotope systematics are 2 and 4 ‰, respectively. The Li and Pb isotope
systematics indicate that the proportions of the altered oceanic crust and oceanic sediment components
in the slab-derived fluid are approximately 80 and 20 %, respectively. This ratio is consistent with thevalues previously obtained based on Sr, Nd and Pb isotope systematics [2].
[1] Schmidt, M.W. & Poli, S. Earth Planet. Sci. Lett. 163, 361-379 (1998)[2] Shibata T. & Nakamura E. J. Geophys. Res. 102, 8051-8064 (1997)Formation and evolution of geochemical reservoirs
2.14

29

Chemical and isotopic evolution of the Allende meteorite
Chie Sakaguchi and Eizo Nakamura
Institute for Study of the Earth’s Interior, Okayama University

Individual chondrules and Ca-, Al- rich inclusions (CAIs) separated from Allende meteorite havebeen analyzed for the following; (1) trace element compositions of the bulk chondrules and CAIs,(2) the constituent minerals, and (3) Rb-Sr and Sm-Nd isotope compositions. Each chondrule was cut
into two pieces; one piece was used for microprobe analyses by SEM-EDX and SIMS, and the otherpiece was used for bulk analyses of trace element compositions and for isotope analyses by ICP-MSand TIMS, respectively. The samples for bulk analyses were subjected to acid-leaching with diluted
HCl solutions, yielding leachates and residues that were analyzed separately.
The Rb-Sr isotope systematics of chondrules show an apparent isochron age of 4.26 ± 0.05 Ga
with an initial 87 Sr/86 Sr = 0.69944 ± 44, and the age obtained from 4 CAIs is 4.15 ± 0.15 Ga with an
initial 87 Sr/86Sr = 0.698914 ± 71. However, the Sm-Nd isotope compositions of the chondrules and
CAIs plot near CHUR on the isotope evolution line of 4.56 Ga. The trace element analyses indicatethat mobile elements such as the alkaline elements and Pb are enriched in the leachates relative toimmobile elements such as REE, which are only slightly dissolved into the leachates with LREE
enrichment. This LREE enrichment in the leachate can be explained by the preferential dissolutionof mesostasis and plagioclase during acid leaching. However, it is difficult to obtain the positive Rb
and Pb spikes, relative to the adjacent elements in the acid leachate trace element pattern, by the
involvement of mesostasis and plagioclase.
It is known that these components, including the matrix in Allende meteorite experienced varying
degrees of Fe-alkali-halogen metasomatic alteration resulting in the replacement of primary phasessuch as mesostasis, plagioclase, melilite and fassaite by Na-, K-, Fe- and Cl-rich phases. It is expectedthat Rb behaved similarly to these mobile elements during this alteration. Based on the Rb-Sr isotopic
analyses, the timing of the Rb disturbance in the chondrules and in the CAIs is considered to havebeen the same. On the other hand, it is known that the Rb-Sr isotope compositions of the Allende
matrix plot off of the isochron lines of the chondrules and CAIs, to the older-age side. Therefore, theRb disturbance would have occurred at significantly later stage after accretion, with Rb introduced
into the chondrules and CAIs from the matrix.
These observations suggest that the Rb-Sr isotope systematics and the concentrations of somemobile elements have been strongly affected by the metasomatic alteration of the Allende meteorite.
This alteration apparently occurred at approximately 4Ga, coincidentally corresponding to the age ofthe peak of cratering on the moon.

Session 2

30
2.15

Petrology, geochemistry and garnet and pyroxene xenocrysts of Belingwe volcanics,
Zimbabwe - indisputable evidence for crustal contamination
K. Shimizu1∗ , E. Nakamura1, and S. Maruyama2

1 The Pheasant Memorial Laboratory for Geochemistry and Cosmochemistry, Institute for Study of the Earth’s

Interior, Okayama University at Misasa, Tottori-ken 682-0122;

2 Department of Earth and Planetary Sciences, Tokyo Institute of Technology, Ookayama 2-12-1, Meguro-ku∗

Tokyo 152-8551, JAPAN<shimmy@geo.titech.ac.jp>

Belingwe greenstone belt, Zimbabwe is a large igneous province dominated by komatiitic andbasaltic rocks. 2.7 Ga Komatiite from Relaince Formation in Belingwe greenstone belt is unusuallyfresh and well studied by many geologists, petrogists and geochemists. However, its origin and tectonic setting remain controversial. Kusky and Kidd (1992) have suggested by structural researcheson the high-strain zones at the base of the volcanic sequence and discovery of ductile ultramafic mylonite from the shear zones that the thick volcanic sequence (Ngezi Group) are remnant of an Archean
oceanic plateau. Trace element ratios of melt inclusions in olivine from Belingwe komatiite also support that the komatiite have extruded as oceanic plateau (McDonough and Ireland, 1993). On the other
hand, Bickle et al. (1994) have argued that the volcanic sequence of the Ngezi Group stratigraphicallyresembles Phanerozoic flood basalt provinces extruded on continental crust. In addition, Pb and Nd
isotopic correlation in Belingwe komatiites and komatiitic basalts indicates that some samples arecontaminated up to 1 % of the basement of 3.5 Ga felsic gneiss (Chauvel et al., 1993). Petrology andgeochemistry of Belingwe komatiites were reported by many previous workers, but those of the voluminous basaltic flows in the Zeederbergs Formation which are overlying thin komatiitic layers are
not well accomplished. Studies of basaltic rocks would give important constraints on tectonic setting
and petrogenesis of Archean greenstone belts. We have also focused on komatiitic basalt and basalt
in Belingwe greenstone belt.
The detail microscopic observations have been done and least altered rock samples were selected.
Major and trace element compositions and isotopic ratios of Nd were determined for whole rocks,
residues (rock powders which are leached with 6M HCl at 80 degree Celsius for 12 hours) andclinopyroxene relicts of least altered komatiite, komatiitic basalt and basalt, and Pb isotopic ratios
were determined for whole rocks.
2.7 Ga Belingwe volcanic rocks were divided into four types (komatiite, komatiitic basalt, depleted basalt and enriched basalt) by the occurrence and texture of the rocks and concentration of
trace element of whole rocks. komatiites are slightly depleted ([La/Sm]N: 0.6 to 0.7) which are identical with previously published data, while that of komatiitic basalts are enriched ([La/Sm]N: 1.5).
Basalts in Zeerderbergs Fm. have both depleted ([La/Sm]N: 0.8) and enriched ([La/Sm]N: 1.2 to 1.6)
REE patterns. Trace element compositions of clinopyroxenes in komatiite, ([La/Sm]N: ∼0.1) komatiitic basalt ([La/Sm]N: ∼0.4), and enriched ([La/Sm]N: ∼0.3) and depleted ([La/Sm]N: ∼0.1) basalt
are equilibruim with those of whole rocks.
Pb-Pb ages of komatiitic basalt and enriched basalt are 2649 ± 89 Ma and 2654 ± 24 Ma, respectively. The obtained ages are identical to that of Chauvel et al., 1993 (2692 ±9 Ma) within errors, but
we suggest the age for Ngezi Group is 2650 Ma, because volcanics of two different types indicate
almost the same age with relatively small errors. Apparent mu values calculated with 2650 Ma are
7.9 to 8.2 for komatiite, 8.4 for komatiitic basalt, 7.7 to 8.2 for depleted basalt and 9.0 for enrichedbasalt.
Sm - Nd isochron for whole rocks of all rock types gives an age of 2887 ± 340 Ma. Whereas,
Sm - Nd isochron for clinopyroxene and residue of all rock types gives an age of 2673 ± 310Ma,
which is quite close to the Pb - Pb age. However, the age error of Sm - Nd isochron is larger thanthat of Pb - Pb isochron. This may be caused by variety of initial 143 Nd/144 Nd. The epsilon Nd at theFormation and evolution of geochemical reservoirs

31time of 2650 Ma have considerable variations for komatiites (-0.7 to 7.3) and depleted basalts (-1.0
to 12.8) and many of them are positive (depleted), whereas, the epsilon Nd values are negative forbasatic komatiites (-1.3 to -2.7) and enriched basalts (-3.3 to -2.3).
There are correlations among rock types and isotopic ratios of Nd and Pb. Samples with low(La/Sm)N values have high epsilon Nd and low mu values (i.e. komatiite and depleted basalt), whereas,
samples with high (La/Sm)N values have low epsilon Nd and high mu values (i.e. komatiitic basalt
and enriched basalt). These correlations can be generated by ∼10 % contamination of the older continental crust.
Indisputable evidence for rift-related basin of Belingwe greenstone belt were also provided in thisstudy. <100 micron garnet xenocrysts (Grs20Pyp35Alm45) were found in komatiite. Major element
content and REE pattern [(Yb/Sm)N ∼ 4] of the garnet xenocrysts suggest that they are garnet granulite origin. ∼300 micron pyroxene xenocrsts (cpx: Wo45En51Fs4, opx: Wo1En90Fs9) were foundin komatiitic basalt. Highly depleted in LREE [(La/Sm)N of cpx: 0.03, opx: 0.007] and major element characters of cpx [Na2O = 1.3 wt% and Al(6)/Al(4) = 1.8] extremely resemble to those fromsubcontinental mantle (tectosphere).
The remarkable variations of epsilon Nd for komatiites and depleted basalts can not be explainedby contamination of older crustal rock. A possibility for these remarkable variations is that the source
mantle plumes of komatiite and depleted basalt are already the heterogeneous at ca. 2.7 Ga. If thesource mantle was homogeneous, there is another possibility. As the volcanic unit of Ngezi Group
have extruded through thick tectosphere which is highly depleted, komatiite and depleted basalt may
have assimilated with the depleted peridotite to some extent, and considerable variations of isotopic
compositions in them were produced.
Acknowledgements
We are indebted to K. Kobayashi, C. Sakaguchi, T. Kuritani, R. Takana and all of the other PML
members for their analytical support and useful discussions. We would also like to thank K. Hirose
and T. Komiya for constructive discussions.
References
Bickle, M.J., Nisbet, E.G., & Martin, A., 1994, Archean greenstone belts are not oceanic crust.: Journal of Geology, v. 102, p. 121-138.
Chauvel, C., Dupre, B., & Arndt, N.T., 1993, Pb and Nd isotopic correlation in Belingwe komatiites
and basalts, in Bickle, M.J., & Nisbet, E.G., eds., Rock units and stratigraphy of the Belingwe
Greenstone Belt: the complexity of the tectonic setting, Geol. Soc. Zimbabwe Spec. Pub. 2, Rotterdam, Balkema, p. 167-174.
Kusky, T.M., & Kidd, W.S.F., 1992, Remnants of an Archean oceanic plateau, Belingwe greenstonebelt, Zimbabwe: Geology, v. 20, p. 43-46.
McDonough, W.F., & Ireland, T.R., 1993, Intraplate origin of komatiites inferred from trace elements
in glass inclusions: Nature, v. 365, p. 432-434.

Session 2

32
2.16

Formation and suppression of AlF3 during HF digestion of rock samples in Teflon bomb
for precise trace element analyses by ICP-MS and ID-TIMS
Hiroyuki Takei, Tetsuya Yokoyama, Akio Makishima and Eizo Nakamura

Pheasant Memorial Laboratory (PML) for Geochemistry and Cosmochemistry Institute for Study of the
Earth’s Interior, Okayama University at Misasa

The abundances of trace elements such as Rb, Sr, Y, Cs, Ba, REE, Pb, Th, and U in silicate rocksamples provide us important information in regarding with crust and mantle evolution. Silicate rocksamples are commonly decomposed using HF in the determination of trace elements by ICP-MS and
TIMS. When samples contain acid resistant minerals such as zircon, the decomposition should be
carried out under high temperature (> 200◦C) condition in a Teflon bomb. However, we found theformation of aluminum fluoride, AlF3, during HF digestion of felsic rock samples, when we used the
Teflon bomb at high temperature and pressure. The AlF3 incorporates trace elements, and can not
be decomposed by conventional methods such as evaporation with HClO4 . The production of thisAlF3 results in lower yields and poor accuracy in analyses of these trace elements by ICP-MS. The
formation of AlF3 is controlled by the chemical composition of the rock samples, in particular, AlF3
dose not form during decomposition of mafic samples with relatively high (Mg+Ca)/Al ratios. We
have developed a new method to suppress the AlF3 formation, in which excess Mg is added to thesample prior to acid digestion in the bomb. This new method makes it possible to accurately determine
the trace element compositions of higher-Al rock samples with lower concentrations of Mg and Ca(e.g. rhyolite and granite). In the trace element analyses by TIMS with isotope dilution techniques(ID-TIMS), AlF3 formation hinders the achievement of isotope equilibrium, resulting in erroneous
results. However, the Mg-addition method removed this problem by suppression of AlF3 formation.

Formation and evolution of geochemical reservoirs
2.17

33

Origin of the ore-forming fluids in the Hishikari gold deposits.
Hiroyuki Takei and Eizo Nakamura

Pheasant Memorial Laboratory (PML) for Geochemistry and Cosmochemistry Institute for Study of the
Earth’s Interior, Okayama University at Misasa.

The Hishikari gold deposits, which are a typical low-sulfidation adularia - silica epithermal vein
type deposit, are known for their extraordinarily high-grade ore and high production of gold. In thisstudy, the origin of the ore-forming fluids responsible for the Hishikari gold deposits was investigatedusing the trace element concentrations and Sr, Nd and Pb isotopic compositions of the ore samples.
The ore deposits occur in sedimentary rocks of the Shimanto Supergroup and/or in the unconformity
overlying Quaternary volcanic rocks of Hishikari lower andesites. We collected samples from thevicinity of the Hishikari ore deposits and from the Daisen No. 5 vein in the Honko deposit, which
occurs in both the Shimanto Supergroup and the Hishikari lower andesites, and the Yusen No. 4 veinin the Yamada deposit, which occurs in the Hishikari lower andesites.
The Sr and Pb isotopic compositions of the volcanic rocks and sedimentary rocks of the Shimanto
Supergroup are distinct, and all of the vein samples plot in a range intermediate to compositionsof those volcanic and sedimentary rocks (Fig. 1 and 2). The Sr and Pb isotopic compositions ofboth sets of vein samples are considered the result of two-component mixing of the Hishikari lower
andesites and Shimanto Supergroup end members, although Yusen vein (Yamada deposit) occurs
in the Hishikari lower andesites. Therefore, the ore-forming fluids of the Hishikari gold deposits
are considered to have been hydrothermal solutions which interacted with both the Hishikari lower
andesites and the sedimentary rocks of the Shimanto Supergroup independent of the location of theore deposits.

34

Session 2

Formation and evolution of geochemical reservoirs
2.18

35

Evaluation of coprecipitation of incompatible trace elements with fluoride during
silicate rock dissolution by acid digestion
Tetsuya Yokoyama, Akio Makishima, Eizo Nakamura
PML, Institute for Study of the Earth’s Interior, Okayama Univ. at Misasa

Determination of abundances of incompatible trace elements such as Rb, Sr, REE, Pb, Th and
U in mafic or ultramafic rock samples (e.g. MORB, peridotites) and isotope ratios of Sr, Nd andPb gives us direct information about chemical evolution of the mantle. Since these samples usuallycontain significantly lower amounts of incompatible trace elements compared to felsic rock types,
decomposition of larger sample sizes (usually >100mg) is required for their precise determination.
Numerous rock dissolution procedures by acid digestion have been developed and have been applied
to mafic and ultramafic samples in our laboratory. However, insoluble fluorides are usually producedwith sample sizes exceeding 100 mg, making it difficult to obtain homogeneous sample solutions for
ICP-MS analysis and column chemistry in hydrochloric or nitric acid media.
Classical studies by Langmyhr and Kringstad (1966) and Croudace (1980) revealed that major
elements (Na, Mg, Ca, Fe and Al) were incorporated into insoluble fluoride compounds such as ralstonite (Na0.88Mg0.88 Al1.12 (F,OH)6·H2O) during sample decomposition leading to erroneous results
during quantitative analysis. Boer et al. (1993) analyzed Sm and Nd in rock solutions which containsuch insoluble fluorides and found that the abundances were 20∼50% lower and Sm/Nd ratios weresystematically higher than accepted reference values. They concluded that such a difference reflected
precipitation of the insoluble fluoride compounds remaining in the solution and taking up differentially Sm and Nd into their crystal structure. Such selective losses should occur to all trace elements
because their affinity to fluoride compounds must be controlled by their individual character suchas ionic radius and valency. Thus, analysis of trace elements in the presence of insoluble fluorides
is likely to yield erroneous results and misleading conclusions especially in regard to discussion of
Rb-Sr, La-Ce, Sm-Nd, U-Th-Pb systematics, or REE patterns.
In this study, we have estimated quantitatively the coprecipitation of trace elements into such fluorides, and suggest an effective method of digestion that can suppress completely fluoride precipitation.
Conventional acid digestion (Methods A and B; Fig. 1) of three samples of mafic and ultramafic silicate rocks resulted in the precipitation of sticky material and very poor yields of certain trace elementin the resultant solution (Fig. 2-a and b). XRD analysis indicated that the precipitates were composedof fluorides such as CaAlF5 , CaMg2 Al2 F12 , Na0.88 Mg0.88 Al1.12(F,OH)6·H2 O and MgF2 , the formationof which depended on the major element composition of the rock sample. Coprecipitation of trace elements appeared to be strongly controlled by both ionic radius and valency of the elements as wellas the species of the host fluoride precipitate, resulting in selective losses of the elements into these
fluorides.
On the other hand, almost 100% of the trace elements were recovered (Fig. 2-c) using larger
amounts of HClO4 than is conventionally used and evaporating the sample to dryness in a step-wisefashion (Method C; Fig. 1). Using this method, white precipitates were formed as oxides of high fieldstrength elements after decomposition of the sample. Coprecipitation of trace elements of interest in
this study with the oxides is negligible except for Th for which 0.5∼3.2% by weight was coprecipitated probably as the insoluble oxide. As our method also results in negligible blank values, it can beused for both the accurate determination of trace element using ICP-MS as well as isotope analysis
using TIMS.
References
Boer, R.H., Beukes, G.J., Meyer, F.M., Smith, C.B., 1993. Fluoride precipitates in silicate wetchemistry: implications on REE fractionation. Chem. Geol. 104, 93-98.
Croudace, I.W., 1980. A possible error source in silicate wet-chemistry caused by insoluble fluorides.

36

Session 2

Chem. Geol. 31, 153-155.
Langmyhr, F.J., Kringstad, K., 1966. An investigation of the composition of the precipitates formed by
the decomposition of silicate rocks in 38-40% hydrofluoric acid. Anal. Chim. Acta. 35, 131-135.

Formation and evolution of geochemical reservoirs
2.19

37

Separation of thorium and uranium from silicate rock samples using two commercial
extraction chromatographic resins
Tetsuya Yokoyama, Akio Makishima, Eizo Nakamura
PML, Institute for Study of the Earth’s Interior, Okayama Univ. at Misasa

U-Th disequilibrium is widely recognized as a powerful tool for the understanding of numerous
geological phenomena in igneous petrology and chronology, marine geochemistry, hydrology of theearth’s surface, archaeology and palaeoclimatology. Recent developments in analytical techniques
have resulted in the use of TIMS as a substitute for the alpha counting method of Th and U isotope
measurements. The advantages of TIMS over alpha counting includes its higher precision, smaller
amount of sample, and shorter analytical procedure. In order to achieve accurate analyses, good isolation of Th and U is indispensable because ionization efficiencies of Th and U in TIMS are so lowthat interference by coexisting elements is a serious problem. For this purpose, ion exchange column
method is generally adopted by many geochemists, and most recent separation techniques, which utilize the anion exchange resin AG1X8, are basically derived from Chen et al. (1986). However, Adriaens et al. (1992) reported that not all Fe and Pb can be separated from U using this method. Since
U forms stable chloride complexes which have high distribution coefficients (Kd ∼ 103 at 9M HClcondition), purification of U is usually done in HCl media. However, Fe3+ , Co2+ , Cu2+ and Zn2+ also
form chloride complexes similar to those of U and are adsorbed on the resin simultaneously, making
it difficult to efficiently separate U in silicate samples.
Recently, two extraction chromatographic resins called U/TEVA·spec and TEVA·spec have beendeveloped and made commercially available by Eichrom Industries Inc. Compared to AG1X8, theseresins have higher selectivity towards retaining tetravalent actinides and U at lower concentrationsof HNO3 or HCl. In particular, use of the U/TEVA resin avoids the problem of interference fromcoexisting anions because this resin is of chelate exchange type. However, only a few geologicalapplications have been undertaken using these resins.
In this study, we present a new technique for the efficient separation of Th and U using a combination of U/TEVA and TEVA resins. Our method overcomes all the problems associated with preexisting separation routines using AG1X8 or U/TEVA resins. In the first column procedure with
U/TEVA·spec resin (Fig. 1), almost all elements except Th and U were eluted by 4M HNO3 . Th wasthen separated using 5M HCl, and U was finally isolated by successive addition of 0.1M HNO3 . Significant amount of Zr still remained in the Th fraction which was then further purified in the second
column stage using TEVA·spec resin. In the second procedure (Fig. 2), Zr was eluted first using 2M
HNO3, then Th was collected by 0.1M HNO3. Both the Th and U fractions obtained by these procedures were sufficiently pure for thermal ionization mass spectrometric (TIMS) analysis. Recovery
yields of Th and U exceeded 90% and total blanks were <19 pg for Th and <10 pg for U.
Our method has advantages over previous methods in terms of matrix effects, tailing problems anddegree of isolation. Since Th and U are effectively separated without suffering any matrix interference
from coexisting cations and anions, this technique can be used not only for the analysis of igneousrock samples but also for the analysis of soils, marine sediments, carbonates, phosphates and sea,ground and surface water.
References
Adriaens, A.G., Fassett, J.D., Kelly, W.R., Simons, D.S. and Adams, F.C., 1992. Determination of
uranium and thorium concentrations in soils: comparison of isotope dilution - secondary ion mass
spectrometry and isotope dilution - thermal ionization mass spectrometry. Anal. Chem., 64: 29452950.
Chen, J.H., Edwards, R.L. and Wasserburg, G.J., 1986. 238 U, 234 U and 232 Th in seawater. Earth Planet.
Sci. Lett., 80: 241-251.

38

Session 2

Formation and evolution of geochemical reservoirs
2.20

39

Precise analysis of 234 U/238 U ratio using UO+
2 ion with TIMS
Tetsuya Yokoyama, Akio Makishima, Eizo Nakamura

PML, Institute for Study of the Earth’s Interior, Okayama Univ. at Misasa

Uranium and thorium series disequilibrium is a powerful geochemical tracer especially for determining various geological processes that occur on the scales ranging from ∼106 years to a few
days, by measuring the activity of short lived radioactive nuclides (e.g. 234 U, 230 Th, 226 Ra, 228 Th, and
234Th). The 234 U-238U dating method has been applied to samples formed in marine environments,
such as corals, ferromanganese nodules, and marine sediments because sea water is in 234 U-238 U disequilibrium (234 U/238U activity ratio = 1.144 ± 0.002, Chen et al., 1986). If the system remains closed
after the achievement of 234 U-238U disequilibrium, 234 U-238U age can be determined by the following
equation:


234



U
−1=
238 U
present



234





U
− 1 exp(−λ234 · t)
238 Uinitial(Eq. 1)

where λ234 is the decay constant of 234 U, t is the time from the closure of the system to present,and (234 U/238 U) represents an activity ratio. It is clear from this equation that precise determination of234
U/238 U ratio in the sample is required in order to carry out precise dating. By the classical methodof using α counting, typical analytical uncertainties in the 234 U/238U ratio are 2∼3%, resulting in a
few % uncertainty in age dating thereby limiting the application for processes ranging from 0.1 ka to
100 ka.
In this study, we have developed a new, highly precise method for U isotopic analysis using athermal ionization mass spectrometry (TIMS) by measuring UO+
2 ions emitted from a very small
amount (100 ng ∼ 10 ng) of U. A new ionization activator, a silicic acid - diluted phosphoric acidsolution, was found to be effective not only for producing a strong and stable UO+
2 ion beam, butalso for keeping the mass fractionation during measurement small and constant. U isotope analysis
was carried out by a “dynamic multi collection” method using three Faraday cup collectors and a secondary electron multiplier in ion counting mode with RPQ (RPQ-SEM). For precise isotope analysis,baselines of the Faraday cups and a conversion factor between Faraday cups and the RPQ-SEM wereprecisely determined simultaneously with U isotope collection. Analytical precision and reproducibility of 234 U/238 U ratio ranged from 0.10 to 0.38% (2σmean) and 0.13 to 0.31% (2σ), respectively, for
100 ng to 10 ng of U extracted from a natural rhyolitic obsidian (JR-2; Fig. 1). Accuracy of our analysis was confirmed by the repeated analyses of the international U standard, NBL-145 (consistent with
NBL-112A, which is formerly NIST SRM960), and we obtained no difference in the 234 U/238U ratio
between previous results and this study (Table 1). The precision and reproducibility were improved
compared with previous studies, and thus our method is especially useful for samples with very low
abundance of U.
References
Chen, J.H., Edwards, R.L., Wasserburg, G.J., 1986. 238 U, 234 U and 232 Th in seawater. Earth Planet.
Sci. Lett. 80, 241-251.
Cheng, H., Edwards, R.L., Hoff, J., Gallup, C.D., Richards, D.A., Asmerom, Y., 2000. The half-livesof uranium-234 and thorium-230. Chem. Geol. 169, 17-33.
Luo, X., Rehkamper, M., Lee, D.-C., Halliday, A.N., 1997. High precision 230 Th/232 Th and 234 U/238 U
measurement using energy-filtered ICP magnetic sector multiple collector mass spectrometry. Int.
J. Mass Spectrom. Ion Processes 171, 105-117.
Stirling, C.H., Esat, T.M., McCulloch, M.T., Lambeck, K., 1995. High-precision U-series dating ofcorals from Western Australia and implications for the timing and duration of the Last Interglacial.
Earth Planet. Sci. Lett. 135, 115-130.

40

Session 2

Formation and evolution of geochemical reservoirs
2.21

41

Precise determination of ferrous iron in silicate rocks
Tetsuya Yokoyama, Eizo Nakamura

PML, Institute for Study of the Earth’s Interior, Okayama Univ. at Misasa

Determination of ferrous iron (Fe2+ ) and Fe2+ /Fe3+ ratio in silicate rocks is important in geological, geochemical and petrological studies for evaluation of weathering, determination of oxygen fugacity in magma, and understanding chemical equilibrium between minerals and silicate melt. Recent
analytical progress using XRF methods enables highly precise determination of total iron in silicate
rock samples with analytical error less than 0.5 % (1σ) (relative standard deviation; e.g., Takei andNakamura, in prep.). Determination of Fe2+ concentrations with the same precision as that for XRF
analyses of total iron would make it possible to more precisely address the above petrological issues.

Although high precision analyses for Fe2+ is relatively easily obtained, the accuracy of these analyses is more questionable. Fe2+ can easily be oxidized during analytical procedures, resulting in erroneously low values. In the classical method of Pratt (Pratt, 1894), which is commonly used in manylaboratories, the sample is decomposed by mixtures of HF and H2SO4 in a covered platinum cruciblethat excludes atmospheric oxygen. This relatively difficult technique often yields low concentrations
owing to the oxidation of Fe2+ . One of the most effective ways to achieve accurate determinationsof Fe2+ is Wilson’s method (Wilson, 1955) or Peters’ method (Peters, 1968), in which excess V5+ isadded to the sample before decomposition. This method is based on the following reversible reaction:
Fe2+ + V5+ = Fe3+ + V4+(1)

This reaction proceeds to the right side in the condition of strong acidity, and the amount of the V4+
reduced from V5+ is then the same as the amount of Fe2+ in the sample. Because V4+ is extremely42

Session 2

resistant to oxidation and does not back-react to V5+ during the analytical procedure, Fe2+ concentration in the sample can be determined by measuring the unreacted V5+ by adding excess Fe2+ after
sample decomposition and then titrating the unreacted Fe2+ with Cr6+ .
By reevaluating the Wilson’s method, we have developed a highly precise method for the determination of ferrous iron (Fe2+) in silicate rocks. With our method, which involves conditioning thesample solution with 5M H2 SO4 in a relatively small beaker (7 mL), the oxidation of Fe2+ or V4+
that leads to erroneous results can be completely avoided, even in 100-hour sample decompositions
at 100oC. We have measured the concentration of FeO in fifteen standard silicate rock powders provided by the Geological Survey of Japan (GSJ; Table 1). Analytical reproducibility was better than
0.5 % (1σ) for all but those samples that had small amounts of Fe2+ (<1.5 wt. % of FeO). Fourteenof these samples gave FeO contents significantly higher than the GSJ reference values. This likely
indicates that the GSJ reference values, obtained by compiling previously published data, contain alarge number of poor-quality data obtained by methods with lower recovery of Fe2+ caused by oxidation and/or insufficient sample decomposition during analyses. To achieve accurate determinations of
Fe2+ in our method, several factors besides the oxidation must be considered, including: 1) long-term
variations in the concentration of Fe2+ solution must be corrected; 2) excess use of the indicator must
be avoided; and 3) the formation of inert FeF+ complex must be avoided during titration when using
boric acid as a masking agent.
References
Peters, A. (1968) Ein neues Verfahren zur Bestimmung von Eisen(II) oxid in Mineralen und Gesteinen.
Neues Jahrb. Mineral., Monatsh. 3/4, 119-125.
Pratt J. P. (1894) On the determination of ferrous iron in silicates. Am. J. Sci. 48, 149-151.
Terashima S., Taniguchi M., Mikoshiba M., and Imai N. (1998) Preparation of two new GSJ geochemical reference materials: JB-1b and coal fly ash JCFA-1. Geostand. Newslett. 22, 113-117.
Wilson A. D. (1955) Determination of ferrous iron in rocks and minerals. Bull. Geol. Surv. G. B. 9,
56-58.

Formation and evolution of geochemical reservoirs
2.22

43

Mantle metasomatism and rapid melt ascent beneath Miyakejima volcano: evidence
from 238 U-230Th disequilibrium
Tetsuya Yokoyama, Katsura Kobayashi, Takeshi Kuritani, Eizo Nakamura
PML, Institute for Study of the Earth’s Interior, Okayama Univ. at Misasa

238

U-230Th disequilibrium of island arc volcanic rocks is a useful geochemical tracer to understand the time scales of fluid transfer, generation of melt, and its ascent in a subduction zone. Here
we present 238 U-230 Th systematics in historic and pre-historic lavas of Miyakejima volcano, which
belongs to Izu-Mariana arc located 200 km south of Tokyo. We also present major and trace element
compositions as well as their Sr and Nd isotopes to estimate the time scale of magmatic processes
beneath the island arc.
Volcanic rocks of Miyakejima consist of tholeiitic basalts and basaltic andesites (SiO2 = 48.3∼62.1%) which have erupted from earlier than 7000 B.P. to 1983 A.D. Based on major and trace element
compositions in this study and tephrochronological work of Tsukui and Suzuki (1998), the volcanic
history of Miyakejima can be divided into four stages (Stage 1: older than 7000 B.P.; Stage 2: 7000
B.P.∼2500 B.P.; Stage 3: 2500 B.P.∼ 1154 A.D.; Stage 4: 1469 A.D.∼1983 A.D.). Strong enrichmentof LILE compared to HFSE, and depleted isotope ratios of Sr and Nd are observed in all stages, implying metasomatism of MORB-type mantle by fluid related processes with little contribution froma sedimentary component as observed in other Izu-arc volcanoes (Ishikawa and Nakamura, 1994;
Moriguti and Nakamura, 1998).

A (230Th/232 Th) vs. (238 U/232Th) diagram (Fig. 1) shows 238 U-230 Th disequilibrium in eruptives
from Miyakejima with large 238 U excess ((238U/230 Th) = 1.2∼1.4), which is also obtained by theaddition of fluid enriched in U relative to Th. In this diagram, samples can be distinguished into
several different trends that correspond to each stage. Judging from the fact that all samples havealmost constant Sr and Nd isotope ratios (87 Sr/86 Sr = 0.70343∼49 : 143 Nd/144 Nd = 0.51307∼11),
they should also have had constant initial Th isotope ratio at the source region (mantle wedge) before
the addition of fluid. Although the trends of Stage 1 and 2 samples can be regarded as whole-rockisochrones that have identical initial Th isotope ratio ((230Th/232 Th)i = 1.295), it is difficult to depict asingle isochron for Stage 3 and 4 samples. Furthermore, most of the Stage 3 and 4 samples have longer

44

Session 2

time of melt transfer (7 ∼ 27 ky) than Stage 2 samples (5 ∼ 11 ky), whereas Stage 2 samples erupted
earlier than Stage 3 and 4 samples#1 . The melt transfer time of Stage 3 and 4 samples increases as
their Th contents increase, which means an increase in the degree of magma differentiation (Fig. 2).
From these observations, the trend of Stages 3 and 4 in Fig. 1 is interpreted as a mixing trend between
old, andesitic magma with higher Th and elevated (230 Th/232 Th) ratio that was differentiated from
Stage 1 magma, and newly injected basaltic magma with lower Th and lower (230 Th/232 Th) ratio.
This means that actual melt transfer time of Stages 3 and 4 is much shorter than apparent from thecalculation above.
In other arcs, contradictions exist in the melt transfer times determined by 238 U-230Th systematics (90 ky for Lesser Antilles, Turner et al. (1996); 30∼50 ky for Tonga-Kermadec, Turner et al.
(1997)) and 230 Th-226 Ra systematics (<8 ky for Lesser Antilles, Chabaux et al. (1999); ∼1 ky for
Tonga-Kermadec, Turner et al. (2000)). However, the minimum melt transfer time determined fromour 238 U-230Th systematics (5 ky) would support rapid melt transfer within several ky. In the subduction zone, therefore, channel flow in the mantle wedge is preferable for rapid melt transfer rather
than conventional permeable flow. The time constraints of melt transfer determined are important forunderstanding physical and chemical processes in the generation of melt beneath island arcs.
#1: Here, the ”melt transfer time” is the time since fluid release from the subducting slab until eruption,
including residence time in the magma chamber. It is calculated by subtracting the time after eruption from
230 Th-ingrowth time since (230 Th/232 Th) = 1.295 to the present observed value.

References
Chabaux F., Hemond C., Allegre C., 1999. 238 U-230 Th-226 Ra disequilibria in the Lesser Antilles arc:
implications for mantle metasomatism. Chem. Geol. 153, 171-185.
Ishikawa, T., Nakamura E., 1994. Origin of the slab component in arc lavas from across-arc variationof B and Pb isotopes. Nature 370, 205-208.
Moriguti, T., Nakamura E., 1998. Across-arc variation of Li isotopes in lavas and implications forcrust/mantle recycling at subduction zones. Earth and Planet. Sci. Lett. 163, 167-174.
Tsukui, M., Suzuki, Y., 1998. Eruptive history of Miyakejima volcano during the last 7000 years.
Kazan 43, 149-166.
Turner, S., Hawkesworth, C., Calsteren, P.v., Heath, E., Macdonald, R., Black, S., 1996. U-series
isotopes and destructive plate margin magma genesis in the Lesser Antilles. Earth and Planet.
Sci. Lett. 142, 191-207.
Turner, S., Hawkesworth, C., Rodgers, N., Bartlett, J., Worthington, T., Hergt, J., Pearce, J., Smith, I.,Formation and evolution of geochemical reservoirs

45

1997. 238 U-230 Th disequilibria, magma petrogenesis, and flux rates beneath the depleted TongaKermadec island arc. Geochim. Cosmochim. Acta 61, 4855-4884.
Turner, S., Bourdon B., Hawkesworth, C., Evans P., 2000. 226 Ra-230 Th evidence for multiple dehydration events, rapid melt ascent and the time scales of differentiation beneath the Tonga-Kermadec
island arc. Earth and Planet. Sci. Lett. 179, 581-593.

Session 2

46
2.23

Petrology and trace element geochemistry of eclogite xenoliths from the Colorado
PlateauTomohiro Usui1 , Eizo Nakamura1, Katsura Kobayashi1 , Shigenori Maruyama2 and Herwart
Helmstaedt3
1. PML, Institute for Study of the Earth’s Interior, Okayama University at Misasa, Tottori 682-0193, Japan
2. Department of Earth and Planetary Sciences, Tokyo Institute of Technology, 2-12-1 Ookayama, Meguro,
Tokyo 152-8551, Japan
3. Department of Geological Sciences and Geological Engineering, Queen’s University, Kingston, Ontario
K7L7 3N6, Canada

It is very important to characterize the trace element behavior in subducted oceanic crust for better
understanding of mantle/crust recycling. Eclogite xenoliths from the kimberlite diatreme breccias on
the Colorado Plateau have been thought to be fragments of subducted oceanic crust, because of theirbasaltic bulk compositions and mineralogical textures similar to ophiolitic eclogites from blueschist
terranes (Helmstaedt and Doig, 1975). The aims of this study are to verify the hypothesis that theeclogite xenoliths are the consequence of subduction zone metamorphism, and to determine the traceelement behavior during subduction of oceanic crust using the eclogite samples.
Eclogite xenoliths examined in this study mainly consist of almandine-rich garnet, Na-clinopyroxene,
lawsonite, phengite, zoisite, rutile, pyrite and zircon. Additionally, coesite, which was identified byusing a laser Raman microscopy, occurs as minute (<20µm) inclusion in garnet. The maximum temperature deduced by the method of Ellis and Green (1979) for neighboring rims of garnet and clinopyroxene ranges from 550◦C to 650◦C at 2.5GPa. Garnet crystals are extremely zoned in Mn content
from core (∼ 1.5wt%) to rim (∼ 0.3wttexture preserving growth during prograde metamorphism. The
major element compositions of clinopyroxene and lawsonite occurring as inclusions in garnet indicatethat the inclusions were crystallized at lower-grade metamorphic conditions than the same minerals
in the matrix.
Trace element compositions of constituent minerals in the eclogite xenoliths were analyzed by an
ion microprobe, Cameca ims-5f. Combining the ion microprobe data with the mineralogical observations, we evaluated the trace element behavior in the eclogite xenolith, which represents subductedoceanic crust. Based on the mass balance calculations, it is suggested that most of HREE and Zr
have been preserved in garnet and zircon, both of which are stable in the eclogite facies. Moreover,approximately 97% of the LREE and 89% of the Sr in the bulk rock are retained by lawsonite, which
decomposes during progressive metamorphism in the lawsonite eclogite facies. From these observations, it is inferred that: (i) migration of the HREE and Zr to the mantle wedge is significantly small;and (ii) migration of the LREE to the mantle wedge is significantly smaller than that of Sr, because
LREE contents in lawsonite increase with nearly constant Sr content during progressive metamorphism.
In addition, we preliminary determined ages of the zircons in garnet of the lawsonite eclogite
by a newly installed ion microprobe, Cameca ims-1270, resulting in approximately 53.4± 6.3 Ma.
Throughout the course of these observations, it is finally concluded that the lawsonite eclogite xenolith
is a fragment of oceanic crust subducted into more than 90km depth along a cold geotherm and witha subduction rate >4cm/y, underneath the North American continent.
ReferenceHelmstaedt H. H. and Doig, R., 1975. Eclogite nodules from kimberlite pipes of the Colorado Plateau- samples of Franciscan-type oceanic lithosphere. Phys. Chem. Earth 9, 95-111
Ellis, D. J. and Green, D. H., 1979. An experimental study of the effect of Ca upon garnet-clinopyroxene
Fe-Mg exchange equilibria. Contributions to mineralogy and Petrology 71, 13-22

Formation and evolution of geochemical reservoirs
2.24

47

Preparation of reliable zircon standard for U-Pb dating method by HR-SIMS
Tomohiro Usui, Katsura Kobayashi and Eizo Nakamura

Pheasant Memorial Laboratory, Institute for Study of the Earth’s Interior, Okayama Univ.

The most outstanding advance in U-Pb dating of zircon has been achieved by the developmentof secondary ion mass spectrometry (SIMS). Sensitive High Resolution Ion Micro Probe (SHRIMP)
can analyze Pb and U isotopic compositions in a small area (< 40µm), and determines precise age
to reveal the multiple growth history in a single grain (e.g. Compston et al., 1984). Recent years,
a High Resolution Secondary Ion Mass Spectrometry (HR-SIMS: CAMECA ims-1270) with multicollection system in our laboratory achieved higher analytical precision and reproducibility for Pb
isotopic ratios as well as higher spatial resolutions than SHRIMP (Kobayashi et al., 2001). However,
one of the major problems in the SIMS analysis exists in the preparation of homogeneous zircon
standard with its precise data in terms of U and Pb concentrations and Pb isotope ratios. Such data
can not be obtained only by SIMS, and thus conventional Isotope Dilution method using Thermal
Ionization Mass Spectrometry (ID-TIMS) is required.
More accurate determination of U/Pb ratios in zircon by SIMS demands several zircon standards
with different U/Pb ratios to correct the instrumental mass fractionation inevitably arising in SIMS
analysis for U-Pb zircon dating. Additionally, the zircon standard preferably has U/Pb ratios similar toan unknown sample. Since zircon is usually zoned, a reliable zircon standard requires not only precise
analysis of U/Pb ratios by ID-TIMS method but also the precise extracting technique that can pickout small and homogeneous zircon fragments from a single zircon crystal. However, conventional
extracting methods such as mechanical abrading (Krogh, 1982) and simple zircon crushing do nothave any abilities for such a precise extraction of zircon fragments, because their methods are intended
for the production of zircon standard for SIMS. In contrast with a remarkable progress of SIMS
techniques that has been applied to many geological events (e.g. Whitehouse et al., 1999), there has
been no study focusing on making reliable zircon standard that is essential part in precision andaccuracy of SIMS analysis.
In order to surmount these shortcomings, we used automatic precision dicing saw in our laboratory(DAD522, Disco ES Ltd.) can separate a crystal into ∼50µm patches under the optical microscope,
which can avoid any inclusions and crack in the crystal. This method enables simultaneous extractionof two fragments with clear spatial relationships from a single zircon crystal for SIMS and TIMS
works, respectively. Recent TIMS technique (Kuritani and Nakamura, 2001) that enabled precise
analysis of Pb isotopes with 1∼10ng of Pb is used to make our zircon standards small enough toachieve homogeneity. By combining these techniques, we prepared zircon standards for U-Pb datingby HR-SIMS, and examined the reliability of our method than previous studies.
In this study, we used two zircon standards that have different U/Pb atomic ratios in order tocorrect the instrumental mass fractionation during SIMS analysis. The first one was 91500 fragment,a zircon standard donated by Harvard Mineralogical Museum in Cambridge (USA), which is thought
to be crystallized at 1065 Ma (Wiedenbeck et al., 1995). Another zircon was from Sri Lanka thatis estimated to be crystallized at ∼560 Ma. Each zircon was cut into two pieces (∼700µm cubic)
using the automatic precision dicing saw. One part from each zircon was mounted with petropoxy154(Palouse Petro Products), and was polished for HR-SIMS standard. This was characterized regarding
their homogeneity by a optical microscope, SEM-EDX (Hitachi S-3100H, Horiba EMAX 7000) and
HR-SIMS. These analyses suggest that our zircon standards are homogeneous with respect to theirmajor element compositions as well as U/Pb atomic ratios within analytical errors. These observations
suffice their possibility for the use of the standards of precise U-Pb dating with HR-SIMS.
For ID-TIMS work, another part was decomposed in a Teflon bomb following the technique described by Krogh (1973) after the addition of the 205 Pb and 233 U spikes. The zircon solution subsequently underwent anion exchange column chemistry with a little modification of Manhes et al. (1984)

48

Session 2

for separation of Pb and U, resulting in >95% of Pb and U recoveries without any material destabilizing their ionization in the mass spectrometry. Pb and U concentration and Pb isotopic compositions
were then determined by ID-TIMS following the techniques described by Kuritani and Nakamura(2001) and Yokoyama et al. (2001). Analytical precision and reproducibility of 207 Pb/206Pb ratio were
0.002% (2σm ) and ∼0.03% (2σ), respectively, for 10ng of Pb extracted from natural zircon. Analytical reproducibility of Pb and U concentrations ranged from 0.1% to 0.5% (2σm) for 10ng of Pb and100ng of U. The total procedural blank of Pb varied from 60pg to 250pg (average: ∼160pg), whereasthat of U was ∼2.0pg. Taking into account the effect of observed blank and the precisions of Pb and U
measurements, our analytical uncertainties are 0.41% for 207 Pb/235U and 0.36% for 206 Pb/238U (1σ),
which are comparable to those in the other leading laboratories. These errors are approximately tentimes lower than those in the HR-SIMS analysis.
U-Pb analysis of the 91500 yielded the radiogenic atomic ratios of 0.07491±3 for 207 Pb/206 Pb,
1.8314±76 for 207 Pb/235U and 0.17731±64 for 206 Pb/238 U (1σ), and the 207 Pb/206 Pb apparent age of
1066±2Ma (1σ). Our result is consistent with that described in Wiedenbeck et al. (1995). On the other
hand, U-Pb analysis of the zircon from Sri Lanka yielded the radiogenic atomic ratios of 0.05887±2
for 207 Pb/206 Pb, 0.7220±30 for 207 Pb/235U and 0.08892±32 for 206 Pb/238 U (1σ), and the 207 Pb/206Pb
apparent age of 561±2Ma (1σ), resulting in 2% discordance from the calculated concordia age of
561Ma.
As described above, we have precisely determined Pb and U contents, and Pb isotopes of two
zircon standards by ID-TIMS with analytical errors comparable to previous studies. Moreover, we
carefully extracted the pristine zircon fragments for ID-TIMS and SIMS work in a single zircon crystal, and separately checked the homogeneity of the standards by SEM-EDX, HR-SIMS and optical
microscope. Therefore, it is concluded that our zircon standards are much more reliable than those
prepared in previous studies.
References
Compston, W., Williams, I. S. and Meyer, C., 1984. Journal of Geophysical Research 89, B525-534
Kobayashi, K., Usui T., Nakamura, E., 2001. International Symposium ”Transport of Material in the
Dynamic Earth”, Kurayoshi, October 2-5, 2001
Krogh, T. E., 1973. Geochimica et Cosmochimica Acta 37, 485-494
Krogh, T. E., 1982. Geochimica et Cosmochimica Acta 46, 637-649
Kuritani, T. and Nakamura, E., 2001. Submitted to Chemical Geology (revised)
Manhes, G., Allegre, C. J. and Provost, A., 1984. Geochimica et Cosmochimica Acta 48, 2247-2264
Wiedenbeck, M., Alle, P., Corfu, F., Griffin, W. L., Meier, M., Oberli, F., von Quadt, A., Roddick, J.
C. and Spiegel, W., 1995. Geostandards Newsletter, 19, 1-23
Yokoyama, T., Makishima, A. and Nakamura, E., 2001. Chemical Geology (in press)
Whitehouse, M. J., Kamber, B. S. and Moorbath, S., 1999. Chemical Geology, 160, 201-224

Formation and evolution of geochemical reservoirs
2.25

49

Trace element migration associated with evolution of metamorphic fluid duringsubduction zone-related metamorphism
Eizo Nakamura and Toshio Nakano

The Pheasant Memorial Laboratory for Geochemistry & Cosmochemistry (PML), Institute for Study of the
Earth’s Interior, Okayama University at Misasa, Tottori 682-0193, Japan

Fluids probably carry components from the slab (oceanic crust and sediment) to the mantle wedge
in deep-seated portions of subduction zones and induce partial melting of the mantle wedge to form
island arcs. Based on multi-isotope systematics including boron and lithium, island arc volcanics
are believed to originate through addition of oceanic slab components to island arc basalt sources
by aqueous fluids released from the subducting slab (e.g., Shibata and Nakamura, 1997; Moriguti
and Nakamura, 1998). Subduction zone metamorphic processes may also affect the trace element
compositions and isotope compositions of rocks subducted to greater depths in the mantle beyond arcs
(e.g., Hofmann & White, 1982). It is, therefore, important to characterize the trace element behavior in
the metamorphic fluid liberated by the devolatilization of the slab. There are, however, few constraints
on the behavior of trace elements in fluids during high-P/T metamorphism, making it difficult toquantitatively model mantle/crust recycling through subduction zones.
Boron isotope and trace element compositions of metabasites from the Sambagawa Metamorphic
Belt, representing the chlorite, garnet and biotite zones, were investigated as a function of distancefrom a boundary between metasedimentary and metabasite rocks. The B isotopic compositions ofmetabasites near the boundary are similar to those estimated for the metamorphic fluids derived fromthe sedimentary layers, and the δ 11 B values increase with decreasing B concentration and increasing
distance from the boundary. Most trace element abundances in the metabasites also decrease withincreasing distance from the boundary, correlate with changing REE patterns in the metabasites inrelation to metamorphic grades. Estimates of REE mobility in fluid by subtraction of REE in thelower grade rocks from those in the higher grade rocks indicate that REE in the metamorphic fluids
are controlled by the dissociation of specific mineral phases in the surrounding metasediments.
These observations suggest that most of the trace elements are transferred by infiltrating aqueous
fluid during metamorphism as released by devolatilization, and that the trace element compositionsof the metamorphic fluids are dependent upon the stability of important carrier minerals (e.g., garnet,
chlorite, mica and rutile) during metamorphism.

Session 250

2.26 Geochemical evolution of the Horoman peridotite complex revised by incorporation oflead isotope systematics: Implications for melt extraction, metasomatism, and compositional
layering in the mantle
Eizo Nakamura, Masaaki Tanimoto, Takeshi Kuritani and Masako Yoshikawa
The Pheasant Memorial Laboratory for Geochemistry and Cosmochemistry (PML), Institute for Study of the
Earth’s Interior, Okayama University at Misasa, Tottori 682-0193, JapanThe Horoman peridotite complex is one of the freshest orogenic peridotite complexes. On the basisof detailed petrologic and geochemical studies, it has been inferred that this complex suffered similar chemical fractionation processes, such as melt extraction, metasomatism, and subsolidus mineral
breakdown, as did other peridotite complexes [e.g., Niida, 1974, 1984; Obata and Nagahara, 1987;
Frey et al., 1991; Takahashi, 1991a, b, 1992; Takazawa et al., 1992, 1995, 1996; Yoshikawa, 1993;
Nakamura and Yoshikawa, 1995; Ozawa and Takahashi, 1995; Takazawa, 1996]. Thus this complex
should be suitable for resolving the origin and timing of chemical and isotopic variations in the peridotites, because their compositions have not been significantly disturbed by secondary alteration near
the surface.
Major and trace element compositions and isotopic ratios of Sr and Nd were determined for bulk
rocks and their constituent clinopyroxenes from the Horoman peridotite complex, Japan. Al2O3 , CaO,
and heavy rare earth elements (HREE) contents of peridotites generally decrease from plagioclaselherzolite through spinel lherzolite to spinel harzburgite, indicating simple melt extraction from a single source. However, the extremely large variations in isotopic (87 Sr/86 Sr = 0.7019 to 0.7066, N d =+ 110 to -10, 206 Pb/204 Pb=17.56 to 18.25, 207 Pb/204 Pb=15.34 to 15.52, 208 Pb/204 Pb=36.82 to 38.22)
and trace element compositions ([Ce/Yb]N = 0.006 to 4.0) cannot be explained by a simple melt extraction mechanism. The samples can be divided into two groups: one suite has depleted isotopic andlight REE (LREE) characteristics (DP), while the other suite shows enriched isotopic and LREE signatures (EP). Sm-Nd isotope systematics of whole-rock DP samples yield an isochron age of 833 ±78 Ma with an initial 143 Nd/144 Nd ratio of 0.5119 ± 2, which is identical to the isotopic compositionof mid-ocean ridge basalt (MORB) source mantle at that time. The relationship between MgO andYb abundances of whole rocks shows that melt extraction was initiated at pressures near the garnetand spinel lherzolite transition. Peridotites that formed at different depths presently occur in close
proximity to each other, sometimes within tens of meters. The chemical and isotopic signatures of theEP samples can be explained by mixing between mantle residue and an isotopically and more incompatible element enriched fluid derived from a subducted slab. This is strongly supported by Pb isotope
systematics indicating that such a metasomatic fluid have been mostly derived from subducted sediments presumably corresponding to Hidaka accretionary complex, which consists of a large amountof terriginous sediments derived from a intracontinent collision zone. These observations suggestthat the small-scale compositional layering observed in the complex may have formed in a wedge
mantle by water-enhanced thinning and folding of metasomatized peridotites which had previously
developed large-scale simple stratification as a result of melt extraction beneath a mid-ocean ridge.
Final uplift from the upper mantle occurred as a result of the uplift of the Hidaka metamorphic beltarising from the collision between the Eurasian and North American plates at approximately 20 Ma
as suggested by Yoshikawa et al. (1993). In conclusion, these observations imply that the Horomanperidotite complex is not a piece of normal convecting mantle, but is a fragment of wedge mantle
extensively deformed and metasomatized under water-bearing conditions associated with subducted
sediments.

Formation and evolution of geochemical reservoirs
2.27

51

Record in metamorphic tourmalines of subduction-zone devolatilization and boron
cycling
Gray E. Bebout1 , and Eizo Nakamura

The Pheasant Memorial Laboratory for Geochemistry and Cosmochemistry (PML), Institute for Study of the
Earth’s Interior, Okayama University at Misasa, Tottori-ken 682-0193 Japan; (1 permanent address:
Department of Earth and Environmental Sciences, Lehigh University, USA; e-mail: geb0@lehigh.edu)

Boron partitioning among fluids, tourmaline, and micas (and other B-rich phases) dictates the extent and location, and isotopic composition, of B loss in subducting slabs (and any additions to arc
sources or the deeper mantle), but the scarcity of information regarding mineral-fluid fractionation
limits efforts to quantitatively model B isotopic evolution. To clarify tourmaline’s role in B redistribution during devolatilization of subducting sediments, we have analyzed tourmaline δ 11B in several
HP and UHP suites, employing high-precision (2σmean = 0.5-0.8 ‰) ion microprobe methods.
Catalina Schist and Sambagawa Belt metasediments (sub-blueschist to amphibolite facies, <1.2
GPa) show contrasting systematics in whole-rock B content and tourmaline abundance. The Sambagawa shows uniform whole-rock B across grade, but an increasing proportion of B in tourmaline
and a decreasing proportion in micas at higher grades (Nakano and Nakamura, 2001). Catalina Schist
rocks contain less tourmaline, and at higher grades show up to 75 percent decrease in whole-rock B(Bebout et al., 1999) with correlated decreased mica B. In both suites, some higher-grade (∼epidoteamphibolite) rocks contain abundant, zoned, dravitic tourmaline. In these rocks, tourmaline cores
have lower Mg/(Mg+Fe+2 ) and Ca/(Ca+Na) and higher δ 11 B SRM951 (up to +1.6 ‰, mostly -6 to-2 ‰), and tourmalines show increased Mg/(Mg+Fe+2 ) (up to 0.825) and Ca/(Ca+Na) and decreasedδ 11 B (approaching -15 ‰) toward rims. These variations are consistent with significant prograde tourmaline growth. Some higher-grade tourmalines have thin outermost zones compositionally similar tocores and attributed to retrogradation. The δ 11B of less-abundant tourmaline in lower-grade rocks (-7
to +4.5 ‰) overlaps with, or is higher than, that of tourmaline cores in higher-grade rocks, and somecores in low- and high-grade rocks are regarded as detrital. Lago di Cignana UHP metasediments(peak 2.7-2.9 GPa; Reinecke, 1998, and van der Klauw et al., 1997, who reported coesite inclusions
in tourmaline) contain dravitic tourmaline zoned in Mg/(Mg+Fe+2), Ca/(Ca+Na), and δ 11 B. We tentatively associate cores with garnet and rutile inclusions, Mg/(Mg+Fe+2 ) up to 0.95, and δ 11 B as low
as -16 ‰ with prograde, high-P/T metamorphism, and rims with clinozoisite and quartz inclusions,lower Mg/(Mg+Fe+2 ), and higher δ 11 B up to +4.3 ‰ with overprinting during exhumation.
The HP suites illustrate contrasting B evolution reflecting the significance of sedimentary processes, bulk-rock/fluid chemistry, or kinetic factors (e.g., need for sedimentary or lower-grade metamorphic “seed crystals”) in controlling the presence and further growth of tourmaline at highergrades. In the absence of tourmaline, devolatilization can reduce whole-rock B (Catalina), whereas
tourmaline growth may serve to retain whole-rock B lost from micas and minimize change in wholerock δ 11 B (Sambagawa). At Lago di Cignana, tourmaline (with δ 11B = -16 to -9 ‰) was apparently
stable to depths of ∼90 km, and correlated major element and B isotope zoning (>16 ‰) demonstrates tourmaline’s ability to preserve complex metamorphic history spanning extremely large P-T
ranges.
Based on our results, and taking into account other recent studies of B isotope fractionation, theB lost during devolatilization from subducting sedimentary rocks (and mica-bearing altered oceanic
crust) is thought to be isotopically heavier than the residual B in the micas. This heavier B could
be sequestered by growing tourmaline (particularly in sedimentary lithologies), in which it could be
subducted to and released at even greater depths, or could be released into and mobilized in metamorphic fluids. These metamorphic fluids (containing B released from micas in subducting sedimentand altered oceanic crust) would appear to be capable of imparting the observed, relatively high-δ 11B
signatures observed in most arc lavas and producing the across-arc trends of decreasing δ 11 B observed

52
in several arcs.

Session 2

Formation and evolution of geochemical reservoirs
2.28

53

Boron isotope geochemistry of metasedimentary rocks and tourmalines in a
subduction-zone metamorphic suite
Toshio Nakano and Eizo Nakamura

The Pheasant Memorial Laboratory for Geochemistry & Cosmochemistry (PML), Institute for Study of the
Earth’s Interior, Okayama University at Misasa, Tottori 682-0193, Japan

Boron is considered to be an excellent tracer of the recycling of subducted oceanic sediment
and/or altered oceanic crust at convergent margins (Morris et al., 1990; Palmer, 1991; Ishikawa andNakamura, 1994). This is because of its extremely high mobility during both magmatic and fluidrelated processes (e.g., Seyfried et al., 1984; Ryan and Langmuir, 1993) and its extreme enrichment in
the subducted oceanic sediments and altered oceanic crust compared with unaltered mid-ocean ridge
basalts (MORB) and upper mantle materials (Spivack and Edmond, 1987; Ishikawa and Nakamura,
1992; 1993). However, B isotopic fractionation behavior during progressive metamorphism is notwell understood making it still difficult to interpret B isotope systematics of island arc and oceanicisland volcanics in relation to mantle/crust recycling through subduction zones.
In order to understand the behavior of boron (B) and its isotope fractionation during subduction
zone metamorphism, B contents and isotopic compositions together with major element compositions
were determined for metasedimentary rocks and tourmalines from the Sambagawa Metamorphic Belt,
central Shikoku, Japan. No systematic changes in whole-rock B content and isotope composition of
the metasediments were observed among the different metamorphic grades, indicating the lack of abulk fluid-rock B-isotope fractionation as a result of devolatilization.
Both modal abundance and grain size of tourmaline increase with increasing metamorphic grade.
In contrast, B contents in muscovite and chlorite decrease with increasing metamorphic grade. Theseobservations combined with mass balance calculations of B suggest the formation of tourmaline during progressive metamorphism from metamorphic fluids containing B mainly derived from muscoviteand subordinately from chlorite without allowing significant net removal of B from the metasedimentary rocks. Tourmalines in the higher-grade metasedimentary rocks have zonal structure of B isotope
and major element composition with decreasing δ 11 B and increasing Mg/(Mg+Fe) from the inner rim(core) to the outer rim. The change of Mg/(Mg+Fe) in the tourmalines with increasing grade is paralleled by similar variation in chlorite. These observations suggest that the growing tourmalines record
the progressive evolution of the B isotopic composition of the metamorphic fluid, in the outermost
rims preserving the isotope signature of peak metamorphic P-T-fluid conditions.
Based on the above observations, the δ 11B of the tourmaline is thought to have been nearly identical to that of the metamorphic fluid resulting in the “apparent” B isotopic fractionation factor between
metamorphic fluid and whole rock (α = (11 B/10 B)f luid /(11 B/10 B) wholerock ) which decreases from 1.007± 0.003 to 1.001 ± 0.003 from chlorite to biotite zone metamorphism. Such results together with theformation of tourmaline from (and sequestering of) B in metamorphic fluids may lead to less Bisotopic fractionation as a result of subduction-zone devolatilization than noted in suites containing
less tourmaline. This, therefore, makes it possible to transport B isotopic signatures, which ultimatelyreflect Earth’s surficial materials to the deep mantle, perhaps resulting in mantle B isotope anomalies
near convergent margins.

Session 2

54
2.29

Simultaneous determination of Ru, Pd, Re, Os, Ir and Pt contents and 187 Os/188 Os ratios
in geological samples using a new anion exchange chromatographic procedure
Akio Makishima, Mayumi Nakanishi and Eizo Nakamura

The Pheasant Memorial Laboratory for Geochemistry and Cosmochemistry (PML), Institute for Study of the
Earth’s Interior, Okayama University at Misasa, Misasa, Tottori-ken 682-0193, Japan

We have developed a new group separation method of Ru, Pd, Os, Ir and Pt (PGEs) and Re in
silicate samples employing less hazardous reagents and procedures. A silicate sample and isotope
spikes were mixed, and digested with HF-HBr in a Teflon bomb at 518 K after the in situ generated
Br2 treatment. Then the sample was dried and redissolved with 0.1 M HBr. After separation of fluoride residues by centrifugation, Re and PGEs were adsorbed on a 0.1 ml TEVA anion exchange resin,which has distribution coefficients (Kd s) of 2200, 16000, 1600, 5500, 4000 and 17000 for Ru, Pd,
Re, Os, Ir and Pt, respectively [Makishima, Nakanishi and Nakamura, Anal. Chem., in press]. Re and
PGEs were recovered into Re and Os fractions, by combination of HCl, HBr and HI with total recovery yields of >90 % for Ru, Pd, Os and Ir, and >70 % for Re for ∼1g silicate sample. The Re fraction
was further purified by 0.1 ml cation exchange resin, AG 50W-X8, and Re and PGE concentrations
were determined by isotope dilution (ID) ICP-MS with detection limit of <4 pg g−1 except Os. The
Os fraction was purified by the micro-distillation technique, and its 187 Os/188 Os ratio was measured
with negative thermal ionization mass spectrometry (N-TIMS). Total blanks were 2∼3 pg for Ir andOs, 8 pg for Re, and 12∼14 pg for Pd and Pt. This technique, therefore, enables us to determine
simultaneously sub ng g−1 levels of Ru, Pd, Re, Os, Ir and Pt and the Os isotope ratio in silicate
samples. Furthermore, common distillable acids were used, Sr, Ce, Nd and Pb isotope ratios shouldbe easily determined from the same sample. In addition, S can be also measured simultaneously bysimply adding S-spike before sample digestion.

Formation and evolution of geochemical reservoirs
2.30

55

Major and rare earth elements geochemistry of the Icelandic Plume: an island-wide
approach for source determination
Estelle F. Rose, A. Makishima, E. Nakamura

Institute for Study of the Earth’s Interior, Okayama University at Misasa, 682-0193 Tottori-ken, Japan

The principal distinguishing feature of Iceland among other plume-fed ocean island is its location on the Mid-Atlantic Ridge (MAR). The geochemical variations in basalts from Iceland and the
Reykjanes Ridge have for a long time been attributed to mixing between two endmembers (1) the
depleted mid-ocean ridge basalt (MORB) source and (2) the enriched Iceland mantle plume (Shilling,
1973; Hart et al., 1973; Sun et al., 1975). In contrast, the occurrence of some Iceland basalts withrelatively depleted incompatible trace element ratios and low Pb and Sr isotope ratios compared tonormal MORB (NMORB) have been interpreted as an indicator that the Iceland mantle plume transports both geochemically enriched and depleted material from the lower mantle into the upper mantle
and that the depleted MORB source does not contribute significantly to the generation of the Icelandbasalts (Thirlwall, 1995; Hards et al., 1995; Kerr et al., 1995; Fitton et al., 1997; Kempton et al., 1998;
Nowell et al., 1998). More recently it has been proposed that no depleted Iceland plume component
existed but rather an enriched low 206 Pb/204 Pb EM1-like component which could correspond to entrained subcontinental lithospheric material embedded in the North Atlantic depleted MORB source(Hanan et al., 2000). On the other hand, another recent study showed that no MORB component wasnecessary to explain the range of chemical and isotopic compositions of Icelandic lavas. It could all
be explained by melting different sections of an old oceanic crust (Chauvel and Hémond, 2000).
In this study we analyzed major and rare earth elements. By looking at samples from various locations, we propose to correct for the common mantle process that affect those lavas and decipher someconstraints on the magma sources. We present here preliminary results of this island-wide approach,since contrary to previous studies, we do not pre-select the rocks neither according to their type or
location, nor to their “primitive” chemical characteristics.
The samples from east (Snaefellsnes Volcanic Zone), south-east (Reykjanes Rift Zone) and south
Iceland (South Iceland Volcanic Zone) as well as from center-north Iceland (North Iceland Rift Zone,
were collected by T. Koyaguchi and A. Makishima in 1993 and 1994).
6 samples are alkali-basalts, 3 from the Snaefellsnes Peninsula SN4, SN5, SN6, and 3 from
Heimaey Island HE1, HE2, HE3. 3 others are basaltic andesites from Hekla HL1, HL2, HL3. All
the other samples are tholeiitic basalts.
In this sample set, no correlation was found between index of differentiation (SiO2 and Fe/(Fe+Mg))
and many incompatible trace element ratios were found. Therefore it is considered that certain traceelement ratios reflect information other than crustal processes such as crystal fractionation and assimilation. For example, we find that Icelandic lavas are characterized by large excess Sr, Ba and Rb
relative to trace elements with similar chemical behavior, which is in good agreement with previous
study (Chauvel and Hemond, 2000). However we notice an additional feature which is a Cs positive
anomaly for 8 samples that seems to correlate with an U positive anomaly (SN6, SN8, SN9, HE1,
TV1, TV2, MY2, GR1). While Rb/Cs might be expected to vary in fluid-derived deposits, we candetermine two groups of sample with constant Rb/Cs with an average of 89 ± 7 and 3 ± 0.3, respectively. The first value is consistent with mantle derived volcanic values (McDonough et al., 1992)
and is similar to what has been previously described for other Iceland rocks (Chauvel and Hémond,
2000). The second, much lower value, has never been described in Icelandic lavas and can not be interpreted as a mixing with any other reservoir because Rb/Cs is lower than any ratio known. Therefore
it could represent (1) different melting conditions during the formation of the lava or (2) an original
heterogeneity of the magma source which needs to be identified. Since this is a new observation for
Icelandic lavas we will focus on characterizing this endmember in more details.
The partition coefficients of Rb and Cs between cpx and melt are too close (Dcpx(Rb) = 0.0010 -56

Session 2

0.0041 and Dcpx (Cs) = 0.00035 - 0.0063; Shimizu, 1974) to account for the significant fractionationof Rb/Cs by melting. As for the partition coefficient for these elements between plagioclase and melt
Dplagio (Rb) = 0.13 - 0.3 and Dplagio (Cs) = 0.13 - 0.14; Villemant et al., 1988. If a source containing
plagioclase and with an initial Rb/Cs=89 undergoes melting, the lowest possible Rb/Cs ratio obtainable is 39. This is much higher than the observed Rb/Cs ≈ 3. Overall, the ratio Drock (Cs)/Drock (Rb)
has to be more than 30 to fractionate Rb/Cs significantly, and such partition coefficient are not reported in mantle silicate phases. Different melting conditions that possibly could fractionate Rb/Cs
are therefore ruled out.
Such low Rb/Cs ratio have been reported in a few xenoliths from Australia (Wilkinson and Taylor, 1980) but further constraints of these sources can be obtained by principlal component analysis(PCA).
This is an ongoing study and our preliminary results should be strengthened by Nd, Sr, Pb isotopic
analysis.
References
Chauvel C. and Hemond C. (2000) Geochem. Geophys. Geosyst.
Fitton J. G., Saunders A. D., Norry M. J., Hardarson B. S. and Taylor R. N. (1997) Earth and Planet.
Sci. Lett. 153, 197-208.
Hanan B. B., Blichert-Toft J., Kingsley R. and Schilling J.-G. (2000) Geochem. Geophys. Geosyst.
Hards V. L., Kempton P. D. and Thompson R. N. (1995) Journal of Geological society of London 152,
1003-1009.
Hart S. R. and Schilling J.-G. (1973) Nature Phys. Sci. 246, 104-107.
Kempton P. D., Fitton J. G., Saunders A. D. and Nowell G. M. (1999) LIP Reader April.
Kempton P. D., Nowell G. M., Fitton J. G., Saunders A. D. and Taylor R. N. (1998) Mineral. Mag.
62A, 759-760.
Kerr A. C., Saunders A. D., Tarney J. and Berry N. H. (1995) Geology 23, 843-846.
Makishima A. (1995). Icelandic Volcanic Rocks No.1, ISEI.
Rudnick R. L., McDonough W. F., McCulloch M. T. and Taylor S. R. (1986) Geochim. et Cosmochim.
Acta 50, 1099-1115.
Schilling J.-G. (1973) Nature 242, 565-571.
Shimizu N. (1974) Geochim. et Cosmochim. Acta 38, 1789-1798.
Sun S. S., Tatsumoto M. and Schilling J.-G. (1975) Science 190, 143-147.
Thirwall M. F. (1995) J. Geol. Soc. London 152, 991-996.
Villemant B. (1988) Contrib. to Mineral. and Petrol. 98, 169-183.
Wilkinson J. F. G. and Taylor S. R. (1980) Contrib. to Mineral. and Petrol. 75, 225-233.

Formation and evolution of geochemical reservoirs
2.31

57

Magma formation at Saba Island, Lesser Antilles: a numerical model for mantle
wedge-slab mixing according to osmium data
Sophie Alves1 , Pierre Schiano2, Bernard Bourdon3, Claude J. Allègre3
1 ISEI, 2 University Blaise Pascal, Clermont-Ferrand, France, 3 IPG-Paris, France

We have studied the osmium isotopic characteristics of samples from Saba Island, Northern Lesser
Antilles. Os concentrations are low (0.7-2.1 ppt) and isotopic ratios are very high (187 Os/188 Os = 0.198- 1.443). When the isotope ratios are plotted versus the inverse of the concentration, the data forms a
positive trend, which can be interpreted as a mixing relationship. Trace element ratios (such as Ba/Th)
and Sr, Nd and Pb isotope ratios suggest that Saba lavas are not affected by continental or sedimentmaterial neither from the down-going slab nor from the crustal basement. We then concluded thathigh Os isotope ratios in lavas may be imparted by the subducting oceanic crust contribution in the
mantle source. In order to evaluate such a hypothesis, we conducted calculations to model mantle andoceanic osmium contributions to the composition of the Saba lavas.
For the model, we considered a binary mixing process between mantle melt and oceanic crust
derived fluid, both contributing Os to the lava. Parameters such as Os concentration and isotope ratio in
the subducting crust and Os partition coefficient during mantle melting and oceanic crust dehydration
are not well known and the initial range of possible values for these parameters were chosen verylarge (e.g., for partition coefficients during melting and dehydration, respectively 0-5000 and 0-1000).
The calculation shows that it is possible to fit the data with such a calculated mixture and that we canestimate statistically likely values for Os partition coefficients: DOs (mantle melting ≈ 1000; DOs (crust
dehydration) ≈ 50. These high partition coefficients suggest that Os is more efficiently mobilized in
the crustal fluid than in the mantle melt. The following figure displays the best model obtained fromthe calculation and shows that 30 to 90% of the osmium occurring in the Saba lavas must come fromthe subducting oceanic crust.

October 4 (Thursday) 09:00 — 12:00

Session 3

High pressure mineral physics

Keynote: Takehiko Yagi, Institute for Solid State Physics, U. Tokyo
Chair: Eiji Ito, Institute for Study of the Earth’s Interior, Okayama U.

59

High pressure mineral physics
3.3

61

In situ determination of the olivine-modified spinel phase boundary in the system
Mg2SiO4 - Fe2 SiO4

Hitoshi Yamada1, Tomoo Katsura1, Toru Shinmei1, Atsushi Kubo1 , Osamu Nishikawa1 , Maoshuang
Song1, Eiji Ito1, Kenichi Funakoshi2
1: Institute for Study of the Earth’s Interior, Okayama University ,2: Japan Synchrotron Radiation Research InstituteThe 410-km seismic discontinuity is usually attributed to the phase transition of olivine to modified spinel in (Mg, Fe) 2SiO4 . From this reason, we have reinvestigated detailed phase relations of
this transition in detail by combination of high P-T in situ X-ray diffraction and compositional analysis. Starting materials were olivine solid solutions with compositions of (Mg x Fe 1-x) 2 SiO4 (x= 0.97, 0.95, 0.93, 0.90, 0.85, 0.80, 0.70). High P-T experiments were conducted using a 6-8 typemulti-anvil press “SPEED-1500” at SPring-8. Pressures were calculated from the unit cell volumeof MgO pressure maker and temperature primarily using EOS of MgO proposed by Jamieson et al.
(1982). We compressed the sample at ambient temperature, and then heated to 1600 or 1900K. We
kept the temperature with taking X-ray diffraction. After that, we quenched the sample. Compositionsof coexisting olivine and modified spinel in the recovered samples were determined using electron microprobe analyzer.

The olivine and modified spinel coexisting loops at temperatures of 1600K and 1900K are foundto be located at pressure range of 12.0∼13.8GPa and 13.0∼15.4GPa, respectively. The dP/dT slopeof the transition in Mg2 SiO4 , which is estimated by extrapolation of the binary loop, is found to be
steeper than that given by Katsura and Ito (1989). The phase relations obtained here suggest thatthe temperature at the 410km depth should be 1850K. However, because of uncertainty of EOS, we62

Session 3

still have 100K uncertainty of temperature estimation. The depth interval of olivine-modified spinel
coexisting regions in the mantle should be 20km if the composition of mantle olivine is (Mg0.89 Fe0.11)
2 SiO4 . The seismic observations showing less than 10km thickness of the 410km discontinuity cannot
be simply explained by applying the phase diagram of olivine-modified spinel transition obtained
here.

High pressure mineral physics
3.4

63

Garnet-ilmenite-perovskite transitions in the system Mg4 Si4O12 -Mg3Al2 Si3O12 : phaseequilibria, calorimetry and application to the mantle
Masaki Akaogi1∗ , Akira Tanaka1 , and Eiji Ito2
1 Dept. of Chem., Gakushuin Univ., Toshima, Tokyo 171-8588
2 ISEI, Okayama Univ., Misasa, Tottori 682-0193

1. Introduction
Recent high-pressure phase equilibrium experiments in simple systems and natural mantle rocks
showed that at normal mantle temperature majorite garnet transforms to perovskite in pyrolitic mantle
at the depth of the uppermost lower mantle (Irifune et al., 1996; Wood, 2000; Kubo and Akaogi,2000). However, phase transition of majorite at lower temperature regime such as in subducting slabs
is not fully clarified. We have examined phase relations at 1000 ◦ C in the system Mg4 Si4O12 (En)Mg3 Al2Si3O12 (Py) which represents a model system for majorite garnet in the mantle. Post-garnet
transitions of Mg3 Al2 Si3O12 pyrope have been studied at temperature down to 800 ◦ C. Enthalpies ofilmenite solid solutions in the En-Py system have also been measured to calculate the phase relations
among garnet, ilmenite and perovskite. These results are used to examine stability of ilmenite solid
solution in the mantle, particularly in subducting slabs.
2. Experimental Methods
High-pressure phase equilibrium experiments were carried out using multianvil apparatus. Starting materials were glasses of (1-x)MgSiO 3-xAl2O3 compositions (x=0, 0.05, 0.1, 0.15, 0.2, 0.25) andcrystalline pyrope. The experimental pressure and temperature ranges were 19-27 GPa and 800-1000◦
C, respectively, and run duration was 3-6 hrs. The recovered samples were examined by microfocus X-ray diffraction technique, and lattice parameters of ilmenite solid solutions were measured bypowder X-ray diffraction method. Calorimetric measurements were performed using a Calvet-type
microcalorimeter at 978 K. Enthalpies of (1-x)MgSiO 3-xAl2O3 ilmenite solid solutions (x=0, 0.1,
0.2, 0.25) were measured by differential drop-solution calorimetry method in lead borate solvent.
3. Results and Discussion
The results of phase equilibrium experiments on Mg3 Al2Si3 O12 have shown that pyrope dissociates into perovskite solid solution and corundum solid solution at about 26 GPa at temperature above
about 1100 ◦ C, while pyrope first transforms to ilmenite and subsequently to perovskite below about1100 ◦ C. At 1000 ◦ C , MgSiO3 ilmenite is stable at 20-24 GPa and Mg3 Al2Si3 O12 ilmenite at about
25-27GPa. These two endmembers form single phase ilmenite solid solution in the whole compositional range of the En-Py system at about 20-27 GPa at 1000 ◦ C (Fig. 1). Garnet solid solution andperovskite solid solution are stable, respectively, below and above the stability field of ilmenite. The
stablity of ilmenite solid solution at 1000 ◦ C in the entire compositional range is very different fromthat at 1600 ◦ C by Kubo and Akaogi (2000) in which ilmenite is stable only in the En-rich composition. This indicates that stability of ilmenite solid solution is very sensitive to temperature and its
field expands rapidly with decreasing temperature.
Enthalpies of ilmenite solid solutions measured by differential drop-solution calorimetry increase
with Al2O3 content in ilmenite. By combining the data of ilmenite solid solutions in this study withthose of garnet and perovskite solid solutions (Akaogi and Ito, 1999), enthalpies of garnet-ilmenite
and ilmenite-perovskite transitions in Mg3 Al2Si3 O12 are estimated to be 150 16 and 99 27 kJ/mol,
respectively. Using the above enthalpy data of Mg3 Al2Si3O12 and those of MgSiO3 (Akaogi and Ito,
1993), the phase relations in the En-Py system have been calculated. The calculated phase boundaries
are generally consistent with those by the experimental determination, and it is confirmed that theilmenite stability field expands to Py-rich composition with decreasing temperature.
Here we assume that mantle composition is pyrolite and that transformations of majorite garnet in
the mantle can be approximated by those of the En-Py system. We choose two adiabatic temperature

Session 3

64

Fig. 1. Phase relations in the En-Py system at 1000 ◦ C.

Fig.2. Zero-pressure density vs. depth.

High pressure mineral physics

65

distributions in which temperature at the 660 km depth is 1600 ◦ C in the normal mantle and 1000 ◦C
in descending slab. Based on the phase relations of the En-Py system at 1600 ◦ C (Kubo and Akaogi,2000), it is suggested that majorite transforms to perovskite at about 660-720 km depth in the normal
mantle (Fig. 2(A)). In contrast, the present results on En-Py system at 1000 ◦ C suggest that majorite
first transforms to ilmenite at about 610-680 km, and subsequently to perovskite at about 700-740 km.
As shown in Fig. 2(B), in the descending slab, a sharp density increase associated with the post-spineltransition exists at about 690 km and two broad regions of density increase, one above and the otherbelow the post-spinel transition are observed.
References
Akaogi, M. and Ito, E. (1993) Geophys. Res. Lett., 20, 1839-1842.
Akaogi, M. and Ito, E. (1999) Phys. Earth Planet. Inter., 114, 129-140.
Irifune, T., Koizumi, T. and Ando, J. (1996) Phys. Earth Planet. Inter., 96, 147-157.
Kubo, A. and Akaogi, M. (2000) Phys. Earth Planet. Inter., 121, 85-102.
Wood, B. J. (2000) Earth Planet. Sci. Lett., 174, 341-354.

Session 366
3.5

MD simulation of the density and P- and S-wave velocity jumps across the 410km and
660km seismic discontinuities
Masanori Matsui
Faculty of Science, Himeji Institute of Technology

The seismic wave discontinuities near 410 and 660 km depths in the mantle have been attributed
to the pressure-induced phase transformations from (Mg,Fe)2 SiO4 olivine to (Mg,Fe)2 SiO4 modifiedspinel, and from (Mg,Fe)2 SiO4 spinel to (Mg,Fe)SiO3 perovskite plus (Mg,Fe)O magnesiowustite,
respectively. In recent publications (Matsui, 1999; Matsui et al., 2000; Matsui, 2000), we have shownthat the molecular dynamics (MD) simulation with realistic potential models, including the breathingshell model for O ions, is very successful in reproducing accurately the observed crystal structures andelastic constants of MgO periclase (Pc), the three Mg2SiO4 polymorphs, olivine (Ol), modified-spinel(Ms) and spinel (Sp), and MgSiO3 perovskite (Pv) over wide T and P ranges, including those found in
the deep Earth. Here we apply MD simulation to predict the density and P- and S-wave velocity jumps
between Mg2 SiO4 Ol and Ms, and between Mg2SiO4 Sp and MgSiO3 Pv plus MgO Pc at the high-Tand high-P conditions corresponding to the 410 and 660 km discontinuities, respectively. We thencompare these simulated density and P- and S-wave velocity jumps with representative seismologicalmodels reported for the 410 and 660 km discontinuities, to study the mineral composition for the
mantle, and to assess the seismological models.
For both the 410km and 660 km discontinuities, the MD simulation for mineral composition
models containing about 50 to 60 vol% (Mg,Fe)2SiO4 component is found to be compatible with arecent model SF99(Shearer and Flanagan, 1999) within nearly 1 sigma limit of SF99 for all density
and P- and S-wave velocity jumps. However, for both the 410km and 660 km discontinuities, our MD
computed density jumps, either for the pyrolite or the piclogite composition, are substantially smaller
than the values by the PREM(Dziewonski and Anderson, 1981) or the ak135 model(Kennett et al.,1995).
References
Dziewonski, A.M., Anderson, D.L., 1981. Phys. Earth Planet. Inter. 25, 297-356.
Kennett, B. L. N., Engdahl, E. R., Buland, R., 1995. Geophys. J. Int. 122, 108-124.
Matsui, M., 1999. Phys. Earth Planet. Inter. 116, 9-18.
Matsui, M., 2000. Phys. Earth Planet. Inter. 121, 77-84.
Matsui, M., Parker, S.C., Leslie, M., 2000. Am. Mineral. 85, 312-316.
Shearer, P. M., Flanagan, M. P., 1999. Science 285, 1545-1548.

High pressure mineral physics
3.6

67

Mineralogy of the lower mantle by the combined method of a laser-heated diamond anvil
cell experiment and analytical transmission electron microscopy
Kiyoshi Fujino1∗, Nobuyoshi Miyajima1, Yohei Sasaki1 , Reki Odawara1 , Hisayuki Ogawa1 ,Nagayoshi Sata2, Takehiko Yagi2
1 Hokkaido University, 2 ISSP, University of Tokyo∗<fujino@ep.sci.hokudai.ac.jp>

1. Introduction
For the study on mineralogy of the Earth’s lower mantle, a laser-heated diamond anvil cell (DAC)
experiment is indispensable because it is almost the only method to produce static high pressures corresponding to the Earth’s lower mantle. However, characterization of the synthesized materials is noteasy because these materials are very small in amount and very fine in grain size. Analytical transmission electron microscopy (ATEM) is suitable for characterization of such materials. With ATEM
we can analyze both structures and chemistries of these materials in a nanometer scale. Therefore, thecombined method of a laser-heated DAC experiment and ATEM is very powerful for the mineralogyof the lower mantle. However, both techniques have several problems to be improved. We present theoutline of a newly installed YLF-laser heated DAC system and the attempts to overcome problems in
AEM analyses of ultrahigh pressure materials, along with the preliminary results on the mineralogyof the lower mantle.
2. Newly installed YLF laser-heated DAC system
The newly installed DAC system consists of double-sided laser heating by a YLF laser and spectroradiometric temperature measurement. The YLF laser is operated in TEM01(donut) mode with
output power of maximum 53 W. The laser-heated spot size ranges from 20 to 150 micrometer indiameter. Temperature profiles of the heated sample are displayed nearly in real time by fitting thethermal radiation to the Plank radiation function. More detailed data of the heating stability, the precision of measured temperatures and temperature profiles will be given in the meeting.
3. Analytical transmission electron microscopy
Although ATEM is very powerful to analyze synthetic ultrahigh pressure materials by DAC [1]and also natural ultrahigh pressure minerals [2], analyses by ATEM have problems as follows [3];1) Ultrahigh pressure materials are very weak against an electron beam and easily become amorphous.
2) In ion-thinning to prepare ATEM foils of these materials, elements are selectively removed
from the surface area of the foils.
3) A properly spread (weak) electron beam must be used to prevent the selective removal of elements from the foil during the measurement of characteristic X-ray intensities.
To overcome the problem 1), we are installing a high-sensitive TV camera system for electron
microscope by which we can observe electron images and diffraction patterns of the ultrahigh pressurematerials under the very weak electron beam.
To resolve the problem 2), we proposed to use the experimentally determined k-factors to calculatecompositions [3]. We are now examining another way to resolve this, ultramicrotomy instead of ionthinning to prepare ATEM foils. Ultramicrotomy was originally developed to prepare ultrathin TEM

68

Session 3

foils in biology. In this method samples embedded in polymeric resin are sectioned into very thin foils
with a diamond knife. So far, the DAC samples could be sectioned into 40 nm thick foils, and EDS
analyses of MgSiO3 perovskite showed an nearly ideal chemical formula.
4. Preliminary results on the mineralogy of the lower mantle
To explore the solubility relations between (Mg,Fe)SiO3- and CaSiO3- perovskites under the lower
mantle conditions, phase relations of the system CaMgSi2O6 -CaFeSi2O6 under the lower mantle conditions and chemical compositions of the product phases have been examined.
In the analyses of the recovered samples synthesized at 20-40 GPa and 1700-2300 K by a laserheated DAC using diopside-hedenbergite solid solutions as starting materials, MgSiO3 perovskite
coexists with CaSiO3 perovskite in the diopside composition. Their EDS analyses showed that the Cacontent in MgSiO3 perovskite was within 1 mol%, but the reliable composition of coexisting CaSiO3
perovskite has not yet been obtained. The obtained Ca content of CaSiO3 perovskite is far lower thanthat of the ideal chemical formula.
References[1] Miyajima, N., Fujino, K., Funamori, N., Kondo, T., and Yagi, T. (1999) Garnet-perovskite transformation under the Earth’s lower mantle conditions: An analytical transmission electron microscopy study. Phys. Earth Planet. Inter., 116, 117-131.
[2] Tomioka, N. and Fujino, K. (1997) Natural (Mg,Fe)SiO3 -ilmenite and -perovskite in the Tenham
meteorite. Science, 277, 1084-1086.
[3] Fujino, K., Miyajima, N., Yagi, T., Kondo, T., and Funamori, N. (1998) Analytical electron microscopy of the garnet-perovskite transformation in a laser-heated diamond anvil cell. Proceedings of the U.S.-Japan Seminar, AGU, Washington, 409-418.

High pressure mineral physics
3.7

69

Mineralogy of subducted basaltic crust (MORB) in the lower mantle
S. Ono1, E. Ito2
1

IFREE, JAMSTEC, 2 ISEI, Okayama Univ.

1. Introduction
Recent seismological studies [1-3] indicate the presence of seismic structures in the lower mantle.
Laboratory experiments indicate that the pressure-induced phase transitions of the olivine component of the mantle occur at about 13.5 GPa and 23.5 GPa, these phase transitions being thought tobe responsible for the seismic discontinuities at 410 km and 660 km depths. However, an adequate
explanation of the seismic observations has not yet been forthcoming. Peridotitic material converts
to an assemblage of Mg-rich and Ca-rich perovskites and magnesiowtite by a depth 750 km, andthis lithology probably persists deep in the lower mantle [4]. Therefore, we explore experimentallywhether phase transitions in subducted oceanic crust (MORB) might be responsible for the deeper
seismic observations.
Recently, we have developed a multianvil apparatus using sintered diamond anvils, which provides
stable conditions for long-duration runs to 50 GPa and 3000 K. Here we report an experimental study
on the phase relations in MORB composition over a pressure range from 25 to 37 GPa, i.e. depthsof 700-950 km in the mantle. The results provide insight into seismic observations and dynamics ofdeeply subducted slabs in the lower mantle.
2. Experimental methods
Sintered diamond anvils with truncated edge length of 2 mm were used in combination withmagnesia octahedron. Pressure calibration was carried out using several fixed points up to higher
than 33 GPa. We further confirmed pressures up to 40 GPa at room temperature, based on equationsof states of Au and MgO (used as pressure markers) by means of in situ X-ray diffraction using
a synchrotron radiation source, SPring-8. High temperature generation was performed by putting asmall cylindrical rhenium heater in the pressure medium of magnesia together with a thermal insulatorof LaCrO3 . Powdered sample of the MORB composition was put directly into the rhenium heater.
After loading to the desired pressure, the sample was heated at 1973 K for at least 3 hours. The
recovered sample was polished and examined by an electron probe-micro analyzer (EPMA). The
lattice parameters of the phases were measured by micro-focused X-ray diffractometry.
3. Results and discussion
Runs were conducted at 25-37 GPa and 1973 K in a multianvil press. The mineral assemblage
observed at 25 GPa was majoritic garnet + stishovite + Ca-perovskite. We confirm that at 30 GPa
majoritic garnet completely disappeared, and the MORB composition crystallized into an assemblageof Mg-perovskite + Ca-perovskite + stishovite + aluminous phase with the calcium ferrite type structure. At pressures between 30 and 37 GPa we encountered no further phase changes. These results
are generally consistent with those of Kesson et al. [5], who examined the high pressure and temperature mineral chemistry of the MORB by transmission electron microscopy on run products from thediamond anvil cell.
We estimated that the change in proportions of major phases in the MORB composition for depths
from 200 to 1000 km in the mantle [6]. A conspicuous change in the mineral assemblage occurs at
depths of 700-800 km, where, corresponding to the disappearance of majoritic garnet, Mg-perovskite
and the aluminous phase appear as major phases and the proportions of Ca-perovskite and stishovite
also increase rapidly. It should be noted, however, that in the region beyond 800 km, no further change

70

Session 3

in the mineralogical constitution of MORB was encountered.
The density of MORB under lower mantle conditions was calculated for various temperatureswith an appropriate equation of state using suitable thermoelastic parameters. The average mantle isdenser than MORB at the top of the lower mantle owing to the transition of spinel to a dominantly
perovskitic lithology at the 660 km discontinuity. However, the density of the MORB increases rapidly
at 700-750 km as a consequence of the change from a majoritic garnet-dominant assemblage to one
dominated by perovskite, which is then denser than surrounding average mantle.
The density profile of MORB is definitely expected to intersect the average mantle density at
depths between 1500 and 2000 km (depending on the temperature within the subducted slab). The
density crossover indicates the presence of a neutral buoyancy region, and the MORB portion of the
subducted slab might stagnate at that depth. Indeed, static neutrally buoyant oceanic crust might also
be responsible for the recently observed transition in seismological heterogeneity at a depth around1600 km. To derive more definite conclusions as to the fate of the entire subducted slab deep in thelower mantle, however, we have to carry out density calculations for the whole slab, including thelayers of depleted peridotite and sediment.
References[1] H. Kawakatu, F. Niu, Seismic evidence for a 920-km discontinuity in the mantle, Nature 371(1994) 301-305.
[2] R. D. van der Hilst, H. Kàrason, Compositional heterogeneity in the bottom 1000 kilometers ofEarth’s mantle: toward a hybrid convection model, Science 283 (1999) 1885-1888.
[3] S. Kaneshima, G. Helffrich, Dipping low-velocity layer in the mid-lower mantle: evidence forgeochemical heterogeneity, Science 283 (1999) 1888-1891.
[4] S. E. Kesson, J. D. Fitz Gerald, J. M. Shelley, Mineralogy and dynamics of a pyrolite lower mantle,
Nature 393 (1998) 252-255.
[5] S. E. Kesson, J. D. Fitz Gerald, J. M. Shelley, Mineral chemistry and density of subducted basaltic
crust at lower-mantle pressures, Nature 372 (1994) 767-769.
[6] S. Ono, E. Ito, and T. Katsura, Mineralogy of subducted basaltic crust (MORB) from 25 to 37
GPa, and chemical heterogeneity of the lower mantle, EPSL 190 (2001) 57-63.

High pressure mineral physics

71

3.8 In situ X-ray diffraction measurements for some mantle minerals to 40 GPa using a
combination of synchrotron radiation and multianvil apparatus with sintered diamond anvils
Tetsuo Irifune∗, Yuichiro Sueda, Takeshi Sanehira, Hisanobu Naka, Toru Inoue, Yuji Higo, Takayuki
Ueda, Ken-ichi Funakoshi, and Wataru Utsumi∗ Geodynamics Research Center, Ehime University, 2-5 Bunkyo-cho, Matsuyama, Ehime 790-8577 Japan

Phase transformations in some mantle minerals have been studied at pressures 20 - 40 GPa andtemperatures to 1800K using SPEED-1500 at SPring-8. High pressure generation was performed
using sintered diamond (SD) cubes of 14 mm edge length, manufactured by Sumitomo Electric Industries Co. Ltd. In situ X-ray diffraction measurements were made by the energy dispersive system equipped at the BL04B1 beamline of SPring-8. Pressure was monitored by unit cell changes of
Au using an adequate equation of state, while the temperature measurement was based on e.m.f. of
W97Re3-W75Re25 thermocouple. Twin sheet heaters of Re was used, and the sample was directly
enclosed in MgO pressure medium. We made several such in situ X-ray diffraction measurements for
MgAl2O4 and Ca0.5 Mg0.5 SiO3, which are both important constituents of the upper mantle peridotite,as the first attempt in our SD experiments at SPring-8.
MgAl2 O4 spinel was found to transform to CaFe2 O4 -type phase at pressures above 25 GPa via anassemblage of MgO periclase and Al2 O3 corundum at lower pressures. No evidence for the formationof the CaTi2 O4 phase, which was reported recently on the basis of diamond anvil cell experiments,
was obtained at pressures to 40 GPa and temperatures to 1800K in the present experiments. The
presence of -MgAl2O4 reported earlier was also not confirmed, and it is thus most likely that the
CaFe2 O4 phase is the only stable phase under these P/T conditions.
Ca0.5 Mg0.5 SiO3 diopside transformed to an assemblage of CaSiO3 and MgSiO3 perovskites atpressures to about 35 GPa and temperatures to 1500K, confirming our recent results based on both
quench and in situ X-ray diffraction measurements at lower pressures (<25 GPa). The bulk moduli(K) of both of these perovskites were estimated based on the present and some existing in situ X-ray
diffraction data, avoiding the effects of the deviatric stress produced in the sample. The preliminary
data analyses yield the K0 values of 230 GPa for CaSiO3 perovskite and 260 GPa for MgSiO3 perovskite, respectively.

Session 372
3.9

Exploration of β-Fe using sintered diamond anvils

Atsushi Kubo, Eiji Ito, Tomoo Katsura, Toru Shinmei, Hitoshi Yamada, Osamu Nishikawa,
Maoshuang Song, Kenichi Funakoshi
Institute for Study of the Earth’s Interior, Okayama University

Based on experiments using the laser-heated diamond-anvil cells (DAC), a new polymorph of Fe,
called β phase, was reported [1, 2]. As to the structure of β-Fe, Saxena et al. [3] have claimed the
double hexagonal closest packing (dhcp), while Andrault et al. [4] have persisted the orthorhombic(space group Pbcm). Both groups [3, 4], however, reported almost the same P-T stability field forβ-Fe; i.e., higher than 35 GPa and ca. 1500K. On the other hand, Yoo et al. [5] observed the dhcp
structure under the conditions of 15 − 40 GPa and 1300 − 2000 K. Shen et al. [6] reported thatthe dhcp phase was never found at high temperature but found in temperature-quenched product atpressures lower than 60 GPa. In order to clarify if the β phase is really existent, it is desirable to do
experiments using a multi anvil apparatus (MAA), which is superior to DAC in terms of larger sample
volume, smaller deviatoric stress, and less temperature gradient through the sample. Funamori et al.
[7] first studied P-T stability of Fe by in situ X-ray diffraction using MAA, and concluded that βphase was not found up to 32 GPa and 1500K. In the present study, we explored β-Fe by in situ Xray observation up to conditions higher than 40 GPa and 1900 K using MAA equipped with sintereddiamond (SD).
Cubic SD anvil assembly was squeezed by a press SPEED-1500 installed at SPring-8. The SD
cubes used in the present study were of 14 mm edge length with truncated edge length (TEL) of 1.5 or
2.0 mm. Several experimental runs were carried out. Figure 1 shows a cross section of a representativecell assembly. A fine powder mixture of Fe + MgO (1:1 in wt.) was put into a semi-sintered MgO
sleeve of 0.6 mm outer diameter (OD), 0.3 mm inner diameter (ID) and 0.5 mm length. The sleeve
was set at the center of a cylindrical Re heater of 0.7 mm OD, 0.65 mm ID and 1.85 mm length. Veryfine powder (< 1 micron) of diamond was packed into the remaining spaces inside the heater. The
Re heater was placed inside a LaCrO3 thermal insulation sleeve that was embedded in an octahedral
pressure medium made of MgO + 5 wt% Cr2 O3. Temperature was monitored using W97%/Re3%W75%/Re25% thermocouple in contact with the central surface of the Re heater. High-pressure insitu X-ray diffraction study was performed by energy-dispersive method using synchrotron radiation
X-ray. Incident white X-ray cut to 0.05 mm in horizontal and 0.1 or 0.2 mm in vertical was irradiatedto the sample from the direction parallel to the axis of the Re cylindrical heater (Fig. 1). The press
position was adjusted to make incident X-ray shoot the sample and the MgO sleeve independently.
The horizontally diffracted beam was introduced to the Ge-SSD through the collimator of 0.05 mm
width to ensure the diffraction angle 2Θ = 6.0◦ . By this method, diffraction patterns of the sample
together with weak reflections from diamond powder were collected (Fig. 2). The generated pressure
was determined by measuring the volume of MgO of the sample sleeve based on the pressure scaleof Jamieson et al. [8]. Diffraction patterns of MgO sleeve and the sample were acquired alternatively
at the same press load and temperature for typically 5 and 10 minutes, respectively. The recovered
samples were polished to make chemical analyses by electron probe micro analyzer (EPMA).
Here we describe the results of a typical run S614 (TEL 2.0 mm). The heating and cooling cycles
were performed up to 1700 and 1900 K at press loads of 662 and 803 tonf, respectively.
In the first cycle at 662 tonf, Fe initially assumed  phase at 37.7(4) GPa and 300 K. On heating,
γ phase appeared at 1400 K and 38.4(2) GPa. The peaks from γ phase intensified in accord withweakening of those from  phase to 1700 K. However, the peaks of  phase still survived at 1700K
presumably due to temperature gradient in the sample. On cooling, the diffraction peaks of γ phasebegan to weaken accompanying with the increase of  peaks at 1300 K and 36.9(1) GPa. However,the diffraction peaks of γ phase still remained even at 300K and ca. 34 GPa. In addition to the above
observation, many small peaks were noticed on increasing temperature higher than 1100K that wereHigh pressure mineral physics

73

thoroughly assigned to diffraction peaks of either FeO (wüstite) or Fe3C (cementite) [9].
On isothermal compression from 662 to 803 tonf at 300 K, γ phase was always observed with phase. The diffraction pattern of FeO slightly changed at 675 tonf and 300 K; i.e., the (111) and(220) reflections became weak and broad presumably due to splitting of these peaks while the intensity of (200) reflection unchanged. These changes indicate phase transition of FeO from B1 to therhombohedral structure [10].
The second heating cycle at 803 tonf was started from 300K and ca. 36 GPa. On heating, nochange in diffraction pattern was observed up to 1300 K. However, the growth of reflections fromγ phase was recognized at 1600K and 41.1(2) GPa, associating with the reduction of  phase, andbecame prompt with increasing temperature (Fig. 2). It should be noted that the diffraction patterns in
Fig. 2 show the larger thermal expansivity of  than that of γ. Phase transition in FeO (rhombohedral
to B1) was also observed at 1000 K and 39.1(4) GPa. The sample was quenched at 42.2(3) GPa and1900 K.
EPMA analyses of the run products revealed that Fe grains and MgO in the sleeve remained as
pure Fe and MgO, respectively. However, marginal portion of MgO grains in the sample changed
to ferropericlase (Fp) showing chemical zoning of 0 − 19 mol% FeO. The FeO component of Fp
should have derived by oxidation of the outer layer near the surface of Fe grains. Existence of Fe3C
was confirmed in the very limited region between the sample and diamond powder by analyzing the
product after aluminum coating.
The results obtained in the present study indicate that the −γ boundary is approximated by T (K)= 670 + 18P (GPa) over pressures from 14 to 42 GPa. We obtained no evidence for the presence of βphase with the dhcp or the orthorhombic Pbcm structure up to 42.2(3) GPa and ca. 1900 K. Therefore,
the present study do not support the stability of β phase under conditions higher than 35 GPa and ca.
1500 K claimed by DAC experiments [1-5].
References[1] Saxena et al., Science 260, 1312 (1993)[2] Boehler, Nature 363, 534 (1993)[3] Saxena et al., Science 269, 1703 (1995)[4] Andrault et al., Science 278, 831 (1997)[5] Yoo et al., Science 270, 1473 (1995)[6] Shen et al., GRL 25, 373 (1998)[7] Funamori et al., GRL 23, 953 (1996)[8] Jamieson et al., in High-Pressure Research in Geophysics, pp. 27-48, CAPJ, Tokyo (1982)[9] Scott et al., GRL 28, 1875 (2001)[10] Fei and Mao, Science 266, 1678 (1994)

Session 3

74

Fig. 1 A schematic drawing of the sample assembly used in the run S614.

Fig. 2 Diffraction patterns obtained at 803 tonf and 1300K (lower) and 1900K (upper). P: periclase, D:diamond, C: cementite, W: wüstite.

High pressure mineral physics
3.10

75

High-pressure phase relationships of FeS

Satoru Urakawa∗, Masayuki Hasegawa, Junji Yamakawa, Ken-ichi Funakoshi, Wataru Utsumi∗

Department of Earth Science, Okayama University

Phase relationships for FeS polymorphs have been studied at the pressure of 16-20 GPa and thetemperature of 300-1350 K by in situ X-ray observation. Experiments were carried out by an energy
dispersive X-ray diffraction method using SPEED-1500 system at BL04B1 of SPring-8. FeS transforms from monoclinic phase to NiAs-type phase via hexagonal phase, with increasing temperature
at 16-20 GPa. Contrary to Fei’s prediction [1], we found that NiAs-type phase of FeS is stable at 18
GPa. The stability fields will extend the higher pressure than 20 GPa. Our results correspond with
the NiAs-hexagonal phase boundary determined by Kusaba et al [2, 3]. Assuming the straight NiAshexagonal phase boundary, we evaluate the location of the NiAs-hexagonal-liquid triple junction at
39.5 GPa and 2300 K.
References[1] Fei, Y. et al., Science, 268, 1892, 1995.
[2] Kusaba, K. et al., Geophysical Monograph 101, 297, 1998.
[3] Kusaba, K., et al., Journal of Physics and Chemistry of Solids, 59, 945, 1998.

Session 3

76
3.11

Experimental determination of compressional wave velocity of olivine aggregate at 1
GPa up to 1273 K
Yoshitaka Aizawa

Institute for Study of the Earth’s Interior, Okayama University , aizawa@misasa.okayama-u.ac.jp

The velocity of the compressional wave (Vp ) was measured for olivine aggregate at temperatures
up to 1273 K at 1 GPa using a piston-cylinder high-pressure apparatus. The samples used in thepresent study were hot-pressed olivine aggregate with uniform chemical composition. The value of Vp
at 1 GPa and an ambient temperature is 8.42 km/s, which is identical to that in previous measurementsof olivine single-crystals. The temperature derivative of Vp (dVp/dT= -4.7 × 10−4 km/s/K) obtained
from the present experiments at 1 GPa is smaller than the previously reported value at an ambientpressure. The temperature gradient for constant velocity (dT/dZ)c allowing the effect of anelasticity
is calculated to be 4 K/km. Taking the temperature gradient in the upper mantle into account (15K/km within the lithosphere and 0.6 K/km within the asthenosphere), it is suggested that the low
velocity zone in the upper mantle could be caused by a drastic change in the temperature gradient at
the lithosphere/asthenosphere boundary.

High pressure mineral physics
3.12

77

Temperature dependence of elastic moduli of polycrystalline wadsleyite and ringwoodite
N. Mayama∗ , I. Suzuki, T. Katsura, A. Yoneda, I. Ohno∗

Earth Science, Okayama University

We synthesized and sintered polycrystalline specimens of wadsleyite (beta-phase) and ringwoodite (gamma-phase) using the San Carlos olivine with composition of ((Mg0.91Fe0.09 )2SiO4) at highP/T conditions. We measured the resonance frequencies of the sphere specimens of beta- and gammaphases for 13 and 8 isotropic oscillatory modes at ambient condition, and the 9 and 6 modes in themwere traced up to 700K. We obtained the adiabatic bulk modulus, KS =167.7(2)GPa and shear modulus, G=106.13(4)GPa for beta-phase, and KS =185.0(1)GPa, G=118.10(5)GPa for gamma-phase, respectively at 298K. The average slope of the moduli were found to be dKS /dT= -0.0176(1) GPa/K,
dG/dT= -0.0156(3) GPa/K for beta-phase and dKS /dT= -0.0188(1) GPa/K, dG/dT= -0.0155(3) GPa/K
for gamma-phase, respectively. The numerals in parentheses indicate probable errors in the last place.
The compressional (vp ) and shear (vs ) wave velocities of the high pressure specimens as well as
the olivine phase under upper mantle conditions were calculated with the aid of the modified Watchman’s equation. Comparing these results with seismic Earth models, we estimated the olivine content
around the 410km-discontinuity, to be about 60 and 45 volume % for vp and vs , respectively. With thesame way, the olivine content was also estimated from 30 to 50 volume % for shear wave at 520kmdiscontinuity. If olivine content is uniform below 410km-depth and has the volume ratio estimatedabove, the jumps of seismic velocities and acoustic impedance at 520km-discontinuity were estimated to be about 1.5% and 2.5%, though some of seismic Earth models do not indicate discontinity
at 520km-depth.

Session 3

78
3.13

X-ray diffraction study of CaO at high pressure and temperature: equation of state of
CaO
Takaya Nagai, Koichi Kittaka, Takamitsu Yamanaka

Department of Earth and Space Science, Osaka University, 1-1 Machikaneyama, Toyonaka, Osaka 560-0043,
JAPAN

Pressure-volume measurements have been performed for CaO at P=0-73 GPa and T=295-643
K using an externally heated diamond anvil cell and angular dispersive X-ray diffraction method.
Heating was accomplished with a Pt/Rh-resistance coil heater placed surrounding diamond anvils andtemperature was measured by almel-chromel thermocouple attached to the surface of a diamond anvil.
Temperature difference between sample and thermocouple was calibrated and it should be within ±3
K. More homogeneous temperature distribution in sample chamber can be expected rather than laser
heating method. Pressure was determined by ruby fluorescence method at 295 K and by the equationof state of gold (Anderson et al., 1989) at high temperatures. Silicon grease was used as pressuretransmitting media and pre-indented rhenium gasket was used. Angular dispersive X-ray diffraction
measurements were carried out using a 12 kW Molybdenum-target rotating anode X-ray generator
with a graphite monochrometer (Rigaku RU-200) and an imaging plate system (Rigaku Raxis IV++)
as a detector. The lattice parameters of CaO and gold (pressure marker) were determined by a profilefitting method, since some diffraction lines of CaO and gold are partially overlapped.
It has been known that the pressure-induced phase transition occurs from the B1 to the B2 structure
on CaO. We confirm that the transition is reversible and the transition pressure is between 61 and 63
GPa and at 295 K. We also confirm that the transition pressure is almost independent of temperatureand this can be explained that the B1-B2 transition is driven by elastic instability rather than bythermal activation as discussed in Tsuchiya and Kawamura (2001).
Room temperature equation of state parameters of the B1 phase are identified as KT 0 = 117(1)GPa with KT 0 = 4 (fixed), whereas the third-order Birch-Murnaghan equation of state yields KT 0 =125(2) GPa and KT 0 = 3.6(1). These parameters are in good agreement with those of Richet etal.(1988). High temperature data sets at 488, 584 and 643 K were analyzed by high-temperature
Birch-Murnaghan equation of state. KT 0 = 3.6 was assumed constant and zero-pressure thermal
expansion α(0, T) of Okajima (1978) was used. We obtained (∂KT 0/∂T )p = −0.0114(4) GPaK−1
and this value is comparable with the value (∂KT 0/∂T )p = −0.0147(4) GPaK−1 of the molecular
dynamics simulation (Matsui, 1998).
References
Anderson, O.L., D.G. Isaak and S. Yamamoto (1989) J. Appl. Phys., vol.65, 1534-1543.
Matsui, M. (1998) J. Chem. Phys., vol.108, 3304-3309.
Okajima, S. (1978) Master thesis of Okayama University.
Richet, P., H.K. Mao and P.M. Bell (1988) J.Geophys.Res., vol.93, no. B12, 15279-15288.
Tsuchiya, T. and K. Kawamura (2001) J. Chem. Phys., vol.114, 10086-10093.

High pressure mineral physics
3.14

79

Thermal equation of state of an omphacite

Yu Nishihara, Eiichi Takahashi, Tomohiro Iguchi, Keisuke Nakayama, Kyoko Matsukage, Miki
Shirasaka and Takumi Kikegawa
Earth and Planetary Sciences, Tokyo Institute of Technology

Omphacite is one of the most abundant minerals in subducted mid oceanic ridge basalt (MORB)
at the pressure condition of 3 - 15 GPa. In order to understand the behavior of MORB in the deep
mantle, knowledge about physical properties of omphacite would be impotent. In situ synchrotron
X-ray diffraction experiments were conducted on an omphacite, which was synthesized at 5 GPa
and 1773 K, at simultaneous high-pressures and high-temperatures by using multi-anvil apparatus
MAX-III, and pressure-volume-temperature data were collected up to 10 GPa and 1000 K. A fit to thehigh-temperature Birch Murnaghan equation of state yielded, isothermal bulk modulus KT 0 = 120(2)GPa, with assumed pressure derivative of bulk modulus K’T = 5.0, temperature derivative of bulk
modulus (dKT /dT)P = - 0.014(5) GPa K−1 , and volumic thermal expansivity with α = a + bT withvalues of a = 1.6(4) × 10−5 K−1 and b = 0.9(7) × 10−8 K−2 , when equation of state of NaCl by Brown(1999) is adopted for pressure scale. The derived KT 0 value is consistent with the linear interpolations
from those of diopside and jadeite in literatures.

Session 3

80
3.15

Static compression of Fe-Si alloy

Naohisa Hirao, Tadashi Kondo, Eiji Ohtani, and Takumi Kikegawa
Tohoku University

Seismic data indicate that the density of the Earth’s outer core is about 10% lower than the densityof iron at the core prssures and temperatures (e.g., (1, 2)). Light alloys with iron are assumed to bepresent to account for the density deficit of the core. H, C, O, S, and Si are considered as the candidatesof the light elements (e.g., (3)). Therefore, the properties of iron compounds at high pressures provide
constrains on the composition of the Earth’s inner core as well as outer core. However, there are little
experimental data at the core pressures on the properties of these iron compounds. Static compression
experiment on Fe-Si alloy was carried out up to 50 GPa at room temperature (4). We have conductedhigh pressure experiments to clarify the structure and stability of Fe-Si alloy, Fe2.3 Si, under highpressures.
High pressure experiments for Fe2.3 Si have been conducted to the pressures above 70 GPa at room
temperature by a Mao-Bell diamond anvil cell (DAC). Angle dispersive x-ray diffraction experiments
at high pressure were performed using a synchrotron radiation source and an imaging plate detector
at the BL-13A beamline of PF. The sample was loaded directly into the hole of a rhenium gasket.
No pressure medium was used in experiments to obtain the maximum sample volume. Pressure wasdetermined by the equation of state for rhenium to evaluate the pressure of the exposure area in x-ray
experiments (5).
Fe2.3 Si alloy has a metallic DO3 structure of Fe3Si showing a body-centered-cubic (bcc) phase.
No evidence for the phase transition was obtained at pressures to 74 GPa in this study. Using thesecond-order Birch-Murnaghan equation of state with a pressure derivative of bulk modulus, 4, the300 K bulk modulus of Fe2.3 Si yields the value of 190 GPa.
References(1) Birch, F., J. Geophys. Res., 69 4377, 1964.
(2) Dziewonski and Anderson, Phys. Earth Planet. Int., 25, 297, 1981.
(3) Poirier, J. P., Phys. Earth Planet. Int., 85, 319, 1994.
(4) Knittle and Williams, Geophys. Res. Lett., 22, 445, 1995.
(5) Vohra et al., Phys. Rev. B, 36, 9790, 1987.

High pressure mineral physics
3.16

81

Theoretical estimations of contribution of electrons to the thermal expansion of Au andPt metals
Taku Tsuchiya and Katsuyuki Kawamura
Dept of Earth and Planetary Sci., TITech, JAPANIntroduction
Au and Pt are often used as the pressure calibration standard in high-temperature high-pressure in-situ
experiments. However, some recent studies suggest that the pressure values measured in high-T P insitu experiments are not in agreement with those determined in quench experiments.1−3 As the originof this discrepancy, uncertainty of pressure effect on a thermocouple, deficiency of used pressure
scale itself, etc. are inferred. Especially, in the case that the equation of state of metal constructed
only using low-temperature experimental data is used as a pressure scale under high-T P conditions,
there is possibility of some underestimation of pressure value. The fact that the effect of the electron
expected to become important under high temperature is hardly taken in these equations of state
causes the underestimation.
However, this electronic effect to the thermal expansion of Au and Pt and its pressure dependence
have never been studied at all. It is quite important to clarify whether this effect is negligible or notunder the temperature pressure conditions of the Earth’s mantle. In the present study based on the solidstate physics and the electronic structure calculation method, we predict contributions of electrons to
the thermal pressures of gold and platinum and discuss the necessity for possible improvement of the
known equations of state under high-T P conditions.

Theory and calculation
According to the solid state physics,4 the contribution of electrons to the thermal pressure in metals
is represented based on the volume dependence of electronic free energy as follows,




1
T2
.
Pe (V, T ) = γe β
2
V

(2)

Here, β and γ are the electronic specific heat coefficient and the electronic Grüneisen parameter,
respectively. These are written as1(3)
β = π 2kB2 N(EF )3γe =∂ ln β.
∂ ln V(4)

And N(EF ) is density of states (DOS) at the fermi level. Consequently, Pe (V, T ) can be estimatedfrom volume V and N(EF ) which can be calculated by the electronic structure of metal. In thisstudy electronic structures of Au and Pt were calculated by the first principles method within thedensity functional theory (DFT) and the local spin density approximation (LSDA). It is known thatthe relativistic effect of valence electrons becomes important in the heavy element after the 5th period.
In this study full-relativistic calculations were performed to correctly include the effect of spin-orbit
coupling. The pressure values at 0K were determined by least-squares fitting of calculated static latticeenergies and volumes to the third-order E-V Birch-Murnaghan equation of state.

Session 3

82

Results and discussion

Calculated thermal pressures of Au and Pt are 249.1(V/V0 )−0.3T 2 Pa (Au) and (379.4−39.2 ln V )(V/V0 )6.2−0.7 ln V T
Pa (Pt), respectively. T dependencies of Pe at V/V0 = 1.0, 0.9 and 0.8 are drawn in the figure, whichindicate some interesting and important results. (1)Au has larger electronic thermal pressures thanthose of Pt under all volume ranges that V/V0 < 1. (2)The volume dependence of Pe is positive ingold but negative in platinum. This difference can be explained by the position of Fermi levels. Since
all 5d orbitals are full in gold, the Fermi level exists in the s band above the d band and the pressure
dependence of Pe is small. On the other hand, 5d orbitals are not full of platinum and the Fermi levelcrosses the d band. This results in the large pressure dependence of Pe . (3)At high-temperature corrections of Pe to the known equations of state may be important. Estimated values of Pe at 1500K and
25 GPa are 0.39 GPa and 0.59 GPa, respectively. However, these are not so large enough to explain
the deviation of pressure values of in situ observations and quench experiments.

References
1. Irifune et al. (1998)
2. Ono et al. (2000)
3. Hirose et al. (2001)
4. Ashcroft and Mermin Solid State Physics (Harcourt)

Pe / GPa

2

1

0
0

Au V/V 0 =1.0
0.9
0.8
Pt V/V 0=1.0
0.9
0.8

1000

2000

T/K

3000

High pressure mineral physics
3.17

83

Thermal diffusivity and thermal conductivity of olivine and garnet under pressures up
to 8 GPa and at temperatures up to 1000 K
Masahiro Osako1, Eiji Ito2 and Akira Yoneda2
1 National Science Museum,Tokyo <sako@kahaku.go.jp>

2 Institute for Study of the Earth’s Interior,Okayama University

To contribute toward study on the dynamic aspects of the Earth, we have measured the thermal
conductivity and the thermal diffusivity of mantle candidate minerals under high pressures and at hightemperatures.
A pulse method using one dimensional heat flow has been applied to simultaneous thermal diffusivity and thermal conductivity measurements. The sample, made in the form of three identical thindisks, is instantaneously heated by a pulse current in a thin heater inserted into one joint of the stacked
disks. Change of temperature is observed on the other joint by a thermocouple. Thermal diffusivity isdetermined from decay time of the temperature profile, and thermal conductivity is calculated fromthe rise of temperature and an power of the current.
Experiments were performed using a uniaxial split-sphere high-pressure apparatus (USSA-1000)
at ISEI. A magnesia octahedron with a edge length of 18 mm was used as a pressure medium intungsten carbide anvils with a truncation of 11 mm. Dimensions of the sample were 4 mm in diameterand 1 mm in height. A furnace composed from a pair of metallic planar heater was installed in theoctahedron for high temperatures. Fig. 1 shows the sample assembly.

Fig. 1. Sample assembly used in this work.

Measurements were made on garnet and olivine single crystals under pressures up to 8.2 GPa.
The garnet sample from Bahia, Brazil, has a composition of 73 % almandine and 25 % pyrope in molar fractions. This kind of garnet belongs to the cubic crystal system and reveals no heterogeneity inthermal conductivity, however, the sample disks were cut so that the end of their surface was perpendicular to one crystallographic axis. The olivine sample from northern Pakistan shows a idomorphic
shape and has a composition of 93% forsterite and 7% fayalite in molar fraction. Olivine reveals highanisotropy in thermal conductivity or thermal diffusivity (Fig. 2), therefore, we need to prepare setsof disk specimens for the three different crystallographic axes.

84

Session 3

Fig. 3 shows the thermal diffusivity and thermal conductivity of the olivine for b-axis under highpressures at room temperature. Error bars come from uncertainty of the thickness of specimens (0.001mm) and from limits of resolutions for time and emf on the temperature profile. Extrapolated valuesof thermal conductivity or thermal diffusivity to zero pressure agree with those obtained by othermethods. Reproducibilities of measurements are within 7 per cent for thermal diffusively and within
15 per cent for thermal conductivity. In the case of olivine for the direction of b-axis the pressure
dependences on thermal diffusively and thermal conductivity is 5 per cent increase per 1 GPa . On
the other hand the garnet is shown that the pressure derivatives of thermal diffusivity and thermal
conductivity tend to decrease as pressure increases, whereas the olivine does not reveal such tendency.
At 5 GPa the garnet has a pressure dependence of 2 per cent a 1 GPa for thermal diffusivity and 3 percent a 1 GPa for thermal conductivity.
High temperature measurements of thermal diffusivity and thermal conductivity up to 1000 K
were made at 8.2 GPa, the terminal pressure in this experiments. The thermal diffusively or the thermal
conductivity of garnet decreases moderately with temperature increase unlike normal insulators in theT-inverse law of thermal conductivity. The same effect on garnet is seen in previous measurements
under zero pressure.
From our results, we can estimate the thermal diffusivity or the thermal conductivity of major
constituents in the upper mantle. In the uppermost mantle garnet could have a thermal diffusivity of
0.9 ∼ 1.0×10−6 m2 s−1 and a thermal conductivity of 2.6 ∼ 2.8×W/mK. Those values would become
0.8 ∼ 0.9 × 10−6 m2s−1 and 2.8 ∼ 3.5 × W/mK at around 400 Km depth, the bottom of the upper
mantle.
References[1] L. D. Dzhavadov. High Temp.-High Pressures, 7: 49-54 (1975).
[2] M. Osako. Bul. Natl. Sci. Mus., Ser E, 20: 1-7 (1997).
[3] H. Fujisawa, N.Fujii, H. Mizutani, H. Kanamori and S. Akimoto. J. Geophys. Res., 73: 4727-4733(1968).
[4] T. Katsura. Geophys. J. Int., 122: 63-69 (1995).
[5] A. E. Beck, D. M. Dahbha and H. H. Schloessin. Phys. Earth Planet. Inter., 17: 35-53 (1978).
[6] G. H. Schaermeri. In: W. Schreyer (Ed.), High Pressure Researchs in Geoscience, E. Schweizerbart’sche Verlangsbuchbandlung, Stuttgart, pp. 349-373 (1982).

High pressure mineral physics

85

Fig. 2. Anisotropic thermal diffusivity of olivine at ambient pressure measured by an Angstrom method [2].

Fig. 3. Thermal diffusivity and thermal conductivity of olivine (parallel to b-axis) versus pressure comparedwith other results [3][4][5][6]. Circles and dotted lines denote thermal diffusivity, and squares, solid lines andcross denote thermal conductivity. Dotted symbols on the zero-pressure line are the thermal diffusivity in Fig.
1 and the calculated thermal conductivity.

Session 3

86
3.18

Non-hydrostatic thermodynamics of anisotropic materials: geological implications forpressure solution
Akira Yoneda

Institute for Study of the Earth’s Interior, Okayama University , Misasa, Tottori-ken 682-0193, Japan

Introduction
The occurrence of earthquakes and presence of highly deformed rock textures are unequivocal
lines of evidence that non-hydrostatic stress and strain play an important role in geological processes. Despite its significance in geoscience, the basic thermodynamic theory for systems in a nonhydrostatic state has not been fully constructed yet, because the current formulation of Gibbs freeenergy, G, under non-hydrostatic stress has been debated for many years (Macdonald 1957; Kamb
1961; Kumazawa 1963; Ito 1966; Ida 1969; Paterson 1973; Robin 1974; Bayly 1987; Dahlen 1992;
Shimizu 1992). In this paper, I present a new formulation of G in a non-hydrostatic stressed state
within the elastic region or thermodynamically reversible region. The formulation is then applied tointerpreting the pressure solution process (e.g., Evans and Kohlstedt 1995).
Formulation
The Gibbs free energy, G, under non-hydrostatic stress is formulated uniquely as

G = F − σij · ωij ωij = V KT Sijkl δkl + ∆+ij ,where F is the Helmholtz free energy, σij the stress, ωij the tensorized volume newly defined,
V the volume, KT the isothermal bulk modulus, Sijkl the elastic compliance, and ∆+ij the isochoricstrain. For the first-order approximation, the above equation can be further simplified as
dG = −KT dV,
where dV is the volume change for any given thermodynamically reversible process, including
both hydrostatic and non-hydrostatic processes. An experimental result regarding a geological processof pressure solution supports the present formulation quantitatively. Further, the present formulation
is successfully used to explain thermodynamic equilibrium in pressure solution process.
An experimental result supporting the present formulation
Elias and Hajash (1992) conducted a quantitative experiment controlling confining pressure andpore-fluid pressure on compaction of quartz sand; they determined chemical potential of quartz component in fluid from its concentration. Figure 1 compares their result with theoretical predictions. In
their work, effective stress, Pe , is defined by subtraction of pore fluid pressure, Pf , from confiningpressure, Pc , or Pe = Pc - Pf . Therefore positive Pe yields compressive state for quartz grains. During
the experiment, Pf was kept constant at 345 bar and temperature at 150◦C. Pe = 0 bar, or, Pc = 345 bar
was the reference state to measure chemical potential change. The Gibbs-Kamb theory cannot explain
experimental results at all. At lower stress, experimental result is in good agreement with prediction
by uni-axial stress, while at higher stress it approaches to that by hydrostatic pressure. It is quite probable that stress state evolves from uni-axial state to tri-axial state with proceeding compaction caused
by increasing effective stress. It is obvious that their result is consistent with the present formulation,
at least in order of magnitude, and definitely against the Gibbs-Kamb theory.

High pressure mineral physics

87

Figure 1 Molar chemical potential, or Gibbs free energy, variation induced by increasing effective stress; ∗ :experimental data; GK: prediction by Gibbs-Kamb theory, dµ∗ = Pf dV + dF ; NH: that by the present
formulation for uni-axial stress, dG = −KT dV = (1/3)V dPe ; HY: that by hydrostatic compression,dG = −KT dV = V dPe .

Session 3

88
3.19

Rates of cation diffusion in olivine and silicate perovskite under mantle conditions
Christian Holzapfel1, D.C. Rubie1 , S. Chakraborty2, F. Langenhorst1 , D.J. Frost1
1 Bayerisches Geoinstitut, Universität Bayreuth, 95440 Bayreuth

2 Institut für Geologie, Mineralogie und Geophysik, Ruhruniversität Bochum, Universitätsstr. 150, 44780

BochumDetermining the kinetics of mineral reactions and phase transformations is essential to understanddynamic and geochemical processes occurring in the Earth. One important area of study in this respectis the determination of the rate of chemical diffusion in mineral phases because slow diffusion ratesin oxides and silicates often control reaction kinetics and can result in compositional zoning duringchemical exchange. From measured chemical zonation patterns it is in principle possible to deduce
information about the time scales of processes, temperatures and pressures. In this study we investigate diffusional behavior in olivine and (Mg,Fe)SiO3 perovskite at high pressures and temperaturesusing the multianvil apparatus.
To investigate diffusion in olivine, single crystal diffusion couples were employed consisting ofpure synthetic forsterite (Mg2 SiO4) and natural San Carlos olivine ( (Mg1.88,Fe0.12 )SiO4), containedin Au and Ni capsules. The oxygen fugacity was either directly buffered at the Ni-NiO buffer (Nicapsules) or was calibrated (Au capsules) by comparison with experiments at known oxygen fugacity.
Observed diffusion profiles are 8 to 25 µm long and can be measured with high accuracy with theelectron microprobe. In a first set of experiments we determined diffusion coefficients at a constanttemperature of 1400◦C and at pressures between 6 and 12 GPa and in a second set at a constantpressure of 12 GPa and temperatures between 1350◦C and 1550◦C. An activation volume of 5.5
cm3 /mole was determined for Fe-Mg exchange, valid along the Ni-NiO buffer up to the upper pressure
limit of olivine stability, which is consistent with former studies at lower pressure [1, 2]. We determine
an activation energy of 282(47) kJ/mole at 12 GPa which is consistent with previous estimates at 1 bar.
Over the range of investigated conditions, no change of slope of either ln D vs 1/T or ln D vs P was
observed, suggesting a single diffusion mechanism in the investigated pressure-temperature regime.
For Ni diffusion the activation volume is 6(1) cm3/mole and the activation energy is 352(72) kJ/mole.
Mn diffusion coefficients are similar in magnitude to those for Fe-Mg exchange but a direct evaluationof the activation volume is difficult because of large analytical scatter due to low Mn concentrations.
These results confirm that divalent cations with similar ionic radii have similar diffusional properties.
Although (Mg,Fe)SiO3 perovskite is the dominant phase of the lower mantle and therefore probably the most abundant mineral in the Earth, there has been only one previous study of diffusion in thisphase (Si self-diffusion in MgSiO3 ) [3]. In this study we have investigated Fe-Mg interdiffusion inpolycrystalline perovskite diffusion couples at 24 GPa and 1750-2100◦C. Samples with different ironcontents (XM g = 0.95 and 0.90) have been synthesized prior to the diffusion experiments. Experiments
were performed at low fO2 in the presence of metallic Fe. Because of extreme slow diffusion, concentration profiles are typically only 100 to 600 nm long. For analysis we use analytical transmission
electron microscopy which, in addition to determining the diffusion profile, also allows a direct structural observation of the diffusion zone. In one sample an apparent influence of subgrain boundaries
on the diffusion profile was observed. Samples at very high temperatures (> 1800◦C) are intensively
twinned, but so far line measurements and elemental maps show no effect of twin lamellae crossing
the interface on diffusion profiles. At 1750◦ C and 24 GPa a diffusion coefficient of 3.66 × 10−20
m2 /sec is derived from our measurements. To explore possible effects of oxygen fugacity we are currently performing experiments at higher fO2 in Re capsules
References[1] Farber et al. (2000):JGR 105(1), 513-529

High pressure mineral physics[2] Misener (1974), In Hoffmann et al.:Geochemical Transport and Kinetics, 117-129, Carnegie Inst
[3] Yamazaki et al.(2000): Toriumi (2000): Phys Earth Planet Int 119, 299-309

89Session 3

90
3.20

Carbon self-diffusion in diamond

KT Koga, MJ Walter, E Nakamura, and K Kobayashi
ISEI, Okayama University at Misasa, Misasa, Tottori, JapanThe formation ages of diamonds have been inferred from various observations, yet they are notconclusive (Richardson et al., 1984, Rudnick et al., 1993; Shimizu and Sobolev, 1995; Pearson etal., 1998). Notable differences in estimated ages are 1 to 2 billion years old using a Sm-Nd model
isochron determined for silicate inclusions in diamonds (Richardson et al., 1984), and less than 120
million years old inferred from the trace element zoning in garnet inclusions in diamonds (Shimizu
and Sobolev, 1995). Advances in micro analytical techniques demonstrated that carbon isotopic ratio
in diamond can show heterogeneity down to the tens of microns (Harte and Otter, 1992, and Hauri etal., 1998). Experimental determination of carbon diffusivity in diamond adds a strong constraint to theduration for which carbon isotopic zoning can persist under the pressure-temperature conditions ofcratonic lithosphere. We have conducted the diffusion experiments at 10GPa and 1800◦ C. Our initial
result of a carbon self-diffusion coefficient is approximately 3×10−21 m2/s in IA diamond annealedfor 24 hours.
Experiments are designed and conducted as follows. The starting material is a small chip of natural
IA diamond which was used as an anvil for DAC apparatus. To minimize the deformation related
crystallographic imperfection, we have used undeformed, faceted girdle portion of a diamond. A thin,approximately 600◦ , isotopically enriched 13 C is vapor deposited on top of polished diamond surface.
Experimental conditions were obtained using a split sphere-type multi-anvil apparatus at ISEI.
Octahedral shaped MgO+5%Cr2O3 was used as a pressure medium, and 50µm thick Re foil was used
as a cylindrical heater. A 13 C coated diamond piece was then packed with in MgO powder and MgOspacers in the central part of the Re heater. The temperature was monitored by a W74 Re26 − W95 Re5thermocouple. After the anneal, the diamond surface was recovered by dissolving MgO in HCl. The
diamond was mounted and surface roughness was determined by a Dektak surface profilometer.
The Diffusion profiles were determined by Cameca 5f at PML, ISEI, Okayama University. We
have used the depth profiling technique on diamond, which is to raster the ion beam from a direction
parallel to the direction of diffusion, to obtain a high resolution diffusion profile. Focused oxygen
beam was rastered over 50µm square area and electronic gate was used to collect the signal from thecental, approximately 35µm square, region. The spattered secondary ions were accelerated to 4.5kV
and an electron multiplier counted the 12 C + ions for 1s and 13 C + ions for 4s. No energy filtering wasapplied, and the mass resolution was approximately 500. Molecular interference of 1H 12 C + to 13 C +was insignificant in our measurements since the range of 13 C/12 C ratio for diffusion experiments far
exceeds that of nature. We also have obtained approximately natural level abundance ratio of 13 C/12 C
in a natural, untreated diamond.
The largest uncertainty of the diffusion experiments is due to mixing of diffusant layer caused bytopography at the diamond surface. Over a distance of 100µm along the surface, a range of 40nm highfrequency topography is common. This topography prevents the determination of diffusion coefficientfrom a profile shorter than 40nm long. Thus the diffusion profiles determined from short duration,
and/or low temperature experiments does not provide significant information. Our 24h annealing at
1800◦C produced a diffusion profile over 120nm in depth while the surface topography ranged over
60nm. The minimum possible diffusivity is 7×10−23 m2/s which is calculated from a profile only due
to surface mixing. Considering small differences between depth profile range and surface topography,
the present result needs to be confirmed with longer duration experiments and with significantly longer
diffusion profile than the surface topography. Until those determination, we only infer the value as an
initial estimate.
The implication to the age of diamond from this result is limited since there is no understandingof the activation energy of diffusion from one data point. However, using the results of ab inito cal-High pressure mineral physics

91

culations for the energy of vacancy formation and migration (Swalin, 1961; Bernholc et al., 1988;
Bernholc et al., 1992), it is possible to extrapolate the initial result to the condition of subcontinental
lithosphere which is around 800-1200◦C. Using 750kJ/mol activation energy from above estimate,diffusivity is about 7 to 17 order of magnitude smaller than 3×10−21 m2 /s reported here. In one billion
years, the equilibration distance will be from 3µm to 30pm. Thus, according to the current inference,
the reported carbon isotope zoning in diamond could persist for durations on the order of the age of
the Earth.
Bernholc J., Antonelli A., Del Sole T. M., Bar-Yam Y., and Pantelides S. T. (1988) Mechanism of
self-diffusion in diamond. Phys.Rev .Lett . 61, 2689-2692.
Bernholc J., Kajihara S. A., Wang C., and Antonelli A. (1992) Theory of native defects, doping anddiffusion in diamond and silicon carbide. J .Mater .Sci.Eng. B11, 265-272.
Harte B. and Otter M. (1992) Carbon isotope measurements on diamonds. Chem.Geol. 101, 177-183.
Hauri E., Pearson D., Dulanova G., and Milledge H. (1998) Microscale variations in C and N isotopes
within mantle diamonds revealed by SIMS .7thInternationalKimberliteConference, 317-319.
Pearson D. G., Shirey S. B., Harris J. W., Carlson R. W., (1998) Sulphide inclusions in diamonds fromthe Koffiefontein kimberlite, S Africa: constraints on diamond ages and mantle Re-Os systematics. EarthPlanet.Sci.lett., 160, 311-326.
Richardson S. H., Gurney J. J., Erklank A. J., and Harris J. W. (1984) Origin of diamonds in oldenriched mantle. Nature 310, 198.
Rudnick, R. L., Elderidge, C. S., and Bulanova, G. P. (1993) Diamond growth history from in situ
measurement of Pb and S isotopic compositions of sulfide inclusions. Geology, 21, 13-16.
Shimizu N., and Sobolev N. V., (1995) Young peridotitic diamonds from the Mir kimberlite pipe.
Nature, 375, 394-397.
Swalin R. A. (1961) Theoretical calculations of the enthalpies and entropies of diffusion and vacancy
formation in semiconductors. J .Phys.Chem.Solids 18, 290-296.

Session 3

92
3.21

Stored strains and boundary characters of dynamically recrystallized quartz aggregates
Osamu Nishikawa
Institute for Study of the Earth’s Interior, Okayama University

The strain energy and grain boundary misorientation recorded in the growing and the consumed
grains during dynamic recrystallization process are fundamental information, respectively about thedriving force of the grain boundary migration and the grain boundary character. I have investigatedstrains stored in grains and grain boundary misorientations for naturally sheared and recrystallized
quartz aggregates (Hatagawa mylonite) on the basis of detailed analyses of the crystallographic orientations and the microstructures.
Analysis of crystallographic orientation distribution patterns using a SEM-EBSP system and TEM
observations revealed that {101̄1}< a >slip system was dominantly operated, and (0001)< a > and{101̄0}< a > slip systems also operated subordinately during deformation. Grain sizes and shapes inquartz aggregates exhibit bimodal distributions: coarse ribbon grains of which their long axes oriented
parallel to mylonitic lineation forming banding structures, and fine equant grains which are prevailed
in highly strained portions. Serrated grain boundaries are well developed in ribbon grains. The grains
favorably oriented for {101̄1}< a > and {101̄0}< a > slip tend to grow to large ribbon grains.
Microstoructural features of these grains are smooth boundary and low intra-grain misorientation(lower than 5◦ ), which is regarded as strain free. These obsevations suggest that grains favorablyoriented for active slip systems ({101̄0}< a > and {101̄1}< a >) rapidly grew at the expense ofdifferently oriented neighboring grains during deformation.
Strains stored in grains have been investigated in more detail by analyzing dislocation sub-structuresusing a TEM. Free dislocation densities and mean subgrain size throughout the samples are 108 cm−2
orders and 1.5µm respectively, and no significant differences are found among grains by their orientation. However, misorientations at the subgrain boundaries are significantly dependent on the crystallographic orientation of host grains. Grains favorably oriented for {101̄0}< a > and {101̄1}< a >slips have the subgrain boundaries with small misorientation by 2◦ , while larger misorientation exceeding several degree are common in other grains. The results of TEM observations indicate thatonce dislocation density exceeds a certain critical value, for example 108 cm−2 orders in the present
case, newly multiplied dislocations are absorbed into subgrain boundaries with rotation of subgrain.
Therefore, the difference of strain energy stored in grains could describe only by those of subgrainboundary energy, which depends on degree of misorientation.
Misorientations between neighboring grains have been calculated in order to investigate the grainboundary characters. For the grains favorably oriented for {101̄0}< a > and {101̄1}< a > slips, inaddition to frequent occurrence of low angle boundaries, those close to Dauphiné twin relationshipwhose twined parts are related by a rotation of 180◦ around c-axis are also found very often. It has
been believed that Mechanical Dauphiné twinning are frequently introduced in α-quartz during plastic
deformation, and cause r{1011} poles to orient parallel to the compression axis rather than z{0111}
pole. However, no preferred orientation of r and z poles was found in the samples. Therefore, it isconcluded that grain boundary characters in the present samples are completely controlled by a- andc-axes preferred orientations, and not by mechanical twining.

High pressure mineral physics
3.22

93

In-situ observation of plastic deformation at high pressure and temperature: A
preliminary study
Masami KanzakiInstitute for Study of the Earth’s Interior, Okayama Univ., Misasa, Tottori 682-0193 Japan;
mkanzaki@misasa.okayama-u.ac.jp

1. Introduction
Deformation experiment utilizing multi-anvil apparatus is difficult, because strain and stress can
not be measured directly by conventional techniques. Therefore no quantitative deformation experiment has been conducted to date. In this study, we developed in-situ sample strain measurement
technique based on X-ray radiography under high pressure and temperature conditions. In order totest this technique, we measured deformation of polycrystalline forsterite (Mg2 SiO4) under compression.
2. Experimental
Experimental setup for X-ray radiography is similar to our in-situ viscosity measurement (Kanzaki
et al, 1998). SPEED1500 press at SPring-8 was employed to generate high pressure and temperature.
18M MgO cell was used. Rod shaped sample (d=1.9mm, l=1.9mm) was obtained by sintering polycrystalline forsterite at 1 atm and 1500 ◦C. Pair of hard alumina pistons is placed at top and bottomof the sample. The sample is covered by 25 micron thick platinum foil. Since platinum absorbs more
X-ray compared to silicate, it acts as a sample shape marker for radiography. In this way, we could
monitor sample shape during the experiment (see Fig. 1). With increasing pressure, part of sample
is eclipsed by WC anvils. Neverthless, we can observe not only height but also diameter, so sample
volume can be calculated. The cell was first compressed to 3 GPa at room temperature. Then it washeated up to 1200 ◦C. After annealing the sample for 2 h, the sample is isothermally compressed to
6.5 GPa (3.5GPa/8h). During the experiment, X-ray image of the sample has been recorded in digital
video tape. We also monitored guide block displacement and anvil gap length during the experiment.

X-ray radiographs of sample before (left) and after (right) the compression. Inclined rectangular area at thecenter is forsterite sample. Arrows show compressional stress directions by hard alumina pistons. Note thatsignificant reduction of sample length by compression.

94

Session 3

3. Results and Discussion
Fig. 2 shows change in sample shape during the deformation experiment. At initial compression atroom temperature, the sample length reduced up to 20%. We noted that there is no significant change
in diameter. Therefore reduction in length is mostly due to reduction of sample porosity. During theheating, sample length reduced slightly with increase of diameter. So at this stage, deformation without volume change has been taking place. At ramping stage, we observed sample length reduction,
but it was much smaller than that of initial compression. Part of this reduction should be due to elastic
deformation. Judging from bulk modulus of forsterite, sample strain due to plastic deformation atwhole ramping stage would be about 1%.
Our preliminary study shows that by directly observing the sample shape, quantitative informationof deformation can be obtained. This is significant improvement over previous indirect studies. Previous studies relied on guide block displacement as an indicator of sample length change. However, ourdata revealed that guide block displacement is not good indicator at all, especially at heating stage.
Present result showed that plastic deformation at ramping stage is very small. This result is alreadypredicted by Durham and Rubie (1998), from their analysis of recovered MgO pressrue medium. In
order to obtain better strain data at ramping stage, high resolution CCD camera and better sample arrangement such as employed by Karato and Rubie (1997) must be employed. We observed small but
real deformation at heating stage. This is not reported before, but stress relaxation of sample during
heating by diffraction technique has been reported. Thus combinding stress and strain measurements,
stress relaxation behavior as a function of temperature could be fully explored.
References:
Durhan and Rubie (1998) in “Properties of Earth and Planetary Materials”, AGU.
Kanzaki et al. (1998) SPring-8 User Experimental Report No.2,1998A, p78.
Karato and Rubie (1997) J. Geophy. Res., 102, 20111-20122.

Height and width (diameter) changes during the compression of polycrystalline forsterite. Data at very initialstage (t≤2000s) contain large error. Note large height reduction at initial compression and following smallchange at ramping.

High pressure mineral physics
3.23

95

Conduction mechanism of fayalite and gamma-Fe2 SiO4: electrical conductivity jump byolivine-spinel transition
D. Sakamoto∗, A. Yoshiasa, O. Ohtaka, T. Yamanaka and K. Ota∗ Earth and Space Science, Graduate School of Science, Osaka University

Fayalite and gamma-Fe2SiO4 are an end member of (Mg1−X FeX )2SiO4 olivine and ringwoodite
solid-solutions, respectively. Bradley et al. (1963) mentioned that the electric conductivities of fayalite
and gamma-Fe2SiO4 were induced by electron hopping, Fe2= Fe3++e.
Xu et al. (1998) measured the electric conductivities of olivine, wadsleyite and ringwoodite underupper-mantle conditions. The 2 orders of conductivity jump was found at the olivine-wadsleyite transition, while the conductivity jump was not found at the wadsleyite-ringwoodite transition. It has notbeen solved yet why ringwoodite (or gamma-Fe2SiO4) exhibits higher conductivity than olivine (orfayalite).
In this study the difference in conduction mechanism of fayalite and gamma-Fe2SiO4 was examined by in-situ measurement of electric conductivity and Seebeck coefficient under high pressure.
Noted that an electron (or hole) mobility can be estimated by electric conductivity and Seebeck coefficient. The mobility is important factor to decide conduction mechanism, because a small poralon
has a low mobility (Emin 1975).
Electric conductivities of fayalite and gamma- Fe2 SiO4 at 6.2 GPa were fitted to the above equation and the fitting results yield the values of =1.99( 0.02) and Ea=0.396( 0.002) eV (376∼670K) for
fayalite =1.99( 0.02) and Ea=0.396( 0.002) eV (534∼1057K) for gamma- Fe2 SiO4, respectively.
A hole mobility of gamma-Fe2SiO4 at 500K, 1000K and 1500K can be estimated as
8.53×10−8 m2V−1 S−1 , 1.32×10−6 m2 V−1 S−1 and 3.29×10−6 m2V−1 S−1 , respectively. These values
are much lower than 8×10−5 m2V−1 S−1 (Emin 1975). This feature shows the hopping conduction isdominant conduction mechanism in gamma-Fe2SiO4 (spinel). It is assumed that an electric conductivity of fayalite exhibits intermediate character between hopping conduction and band conduction.
On the other hand, it is considered an electric conductivity of gamma-fayalite exhibits hopping conduction. Hence, it is concluded that conductivity jump between ringwoodite (or gamma-Fe2SiO4) andolivine (or fayalite) is caused by a difference of conduction mechanism.

Session 3

96
3.24

Electron energy-loss spectroscopy of garnet-perovskite high-pressure assemblages

Nobuyoshi Miyajima1 , Falko Langenhorst1 , Daniel Frost1, Dave Rubie1 and Takehiko Yagi2
1 Bayerisches Geoinstitut, Universitat Bayreuth, D-95440 Bayreuth, Germany
2(Nobuyoshi.Miyajima@uni-bayreuth.de)Institute for Solid State Physics, University of Tokyo, Kashiwa 277-8581, Japan

Introduction: The valences of iron and its cation sites in (Mg,Fe)SiO3-Al2O3 perovskite are
generally considered to have significant effect on the physical properties of the earth’s lower mantle(e.g., electrical conductivity and cation diffusion). Recently, the high relative Fe3+ concentration in(Mg,Fe)(Si,Al)O3 perovskite was determined in the relation to the Al content and the possibility ofa significant concentration of oxygen vacancies was demonstrated, using Mössbauer spectroscopy[McCammon, 1997] and electron energy-loss spectroscopy (EELS) [Lauterbach et al, 2000 & Gloter
et al., 2000]. The latter has much potential for determination of redox states and coordination ofchemical elements in high-pressure samples in which nanometer spatial resolution is often required
due to the small size of the coexisting phases. In this study, iron oxidation states in garnet-perovskite
transformation have been investigated using analytical transmission microscope (ATEM) equippedwith parallel electron energy-loss spectrometer (PEELS). We discuss the Fe3+ / Fe ratio and thesubstitution mechanism of trivalent species of aluminous silicate perovskite.
Experimental: A natural orthopyroxene and two natural garnets with different chemical compositions were used as starting material (Table 1). The synthesis of a (Mg,Fe)(Si,Al)O3 perovskite andmajorite assemblage from natural orthopyroxene has been performed at 24.5 GPa and 2223 K using
multi-anvil apparatus (for orthopyroxene starting material, run no. H1488). On the other hand, thehigh-pressure transformations from garnet to perovskite have been carried out at about 35 GPa (forpyrope, no. Prp01) and 65 GPa (for almandine, no. Alm01) using laser-heated diamond anvil cells.
The Fe L2,3 -edge electron energy-loss near-edge structure (ELNES) of product phases has been examined by ATEM operating at 200kV with PEELS and energy-dispersive X-ray spectrometer (EDS).
Fe L2,3 ELNES is a sensitive probe of the Fe3+ / Fe ratio which can be measured with a small
electron probe, enabling the separate acquisition of spectra of silicate perovskite and the coexisting
phases in a mixed assemblage. Quantification of the spectra followed the procedure described by van
Aken et al. (1998), using an empirically calibrated universal curve.

Results and Discussions: The chemical composition and Fe3+ / Fe ratio of all analyzed samples
are summarized in Table 1. A typical Fe L2,3 ELNES spectrum for silicate perovskite and coexisting
majorite from multi-anvil samples are illustrated in Fig. 1. Based on a simple fingerprint techniqueusing a reference Fe2+ spectrum of the starting material orthopyroxene, the Fe L2,3 edge spectrum
for majorite consists mainly of Fe2+ and perovskite indicates primarily a mixed iron valence compound of Fe2+ and Fe3+ . This result suggests that in the silicate perovskite-majorite assemblage ofHigh pressure mineral physics

97

Fig. 1. Typical Fe L2,3 -edge ELNES of silicate perovskite and majoritic garnet.

Fig. 2. Relation between Fe3+ /Fe and Al composition for silicate perovskite.

98

Session 3(Mg,Fe)SiO3-Al2O3 system, Fe3+ has strong affinities with the perovskite phase under the same pressure, temperature and oxygen fugacity. Perovskites transformed from two garnets also contains thehigh relative Fe3+ concentration, compared with the starting material garnets. In order to elucidate therelation between Fe3+ / Fe and Al composition for Al-rich perovskite, the compositions were plotted on Fig. 2, followed the Fig. 6 in Lauterbach et al. (ibid). Although the relative low Al-perovskite
coexisted with majorite is plotted in a linear trend with their data, Al-rich perovskites are plotted outof the trend. This preliminary result might suggest a significant change of the following two mecha3+
3+
nisms of incorporation of trivalent species into MgSiO3 perovskite: (1) 2Si4+
B = 2(Al , Fe )B + VO
2+
4+
3+
3+
and (2) MgA + SiB = FeA + AlB . Which mechanism of Eq. (1) or (2) becomes favorable for Al3+
and/or Fe3+ in the lower mantle under the conditions of much higher pressure and the different bulk
compositions such as silica-rich bulk composition (e.g., subducted slab) or MgO-rich composition?
We are therefore currently examining samples in Fe-free system of MgSiO3-Al2O3 for comparison.
References
Gloter, A., Guyot, F., Martinez, I. & Colliex, C. (2000): Electron energy-loss spectroscopy of silicate
perovskite-magnesiowustite high-pressure assemblage. American Mineralogist, 85, 1452-1458.
Lauterbach, S., McCammon, C.A., van Aken, P., Langenhorst, F. & Seifert, F. (2000): Mössbauer and
ELNES spectroscopy of (Mg,Fe)(Si,Al)O3 perovskite: a highly oxidized component of the lower
mantle. Contrib Mineral Petrol, 138, 17-26.
McCammon, C.A. (1997): Perovskite as a possible sink for ferric iron in the lower manetle. Nature,
387, 694 - 696.
van Aken P.A., Liebscher, B., Styrsa, V.J. (1998): Quantitative determination of iron oxidation statesin minerals using Fe L2,3 -edge electron energy-loss near-edge structure spectroscopy. Phys Chem
Minerals, 25, 323-327.

High pressure mineral physics
3.25

99

NMR evidence for a new water dissolution mechanism in depolymerized silicate melts:
results for hydrous diopside composition
X. Xue∗, T. Kawamoto, M. Kanzaki∗ Institute for Study of the Earth’s Interior, Okayama University , Misasa, Tottori 682-0193 JapanThe dissolution of water in silicate melts significantly affects their phase relations, and physical/thermodynamic properties. The mechanism of water dissolution is thus a subject of intensive interest in earth sciences and has been studied by various spectroscopic techniques. However, almost all
the studies thus far have been on relatively polymerized systems, from fully polymerized silicates (e.g.
SiO2) to disilicate compositions (e.g. Na2 Si2O5 ). These studies have shown that water is dissolved ineither molecular H2 O or SiOH at least for Al-free systems. More depolymerized silicate melts, suchas metasilicate compositions are also geologically important, because mantle-derived magmas are
known to become less silica- rich with depth of generation. Such magma could be particularly important during the early evolution of the earth’s mantle.
We have studied a hydrous diopside (CaMgSi2 O6) glass (quenched melt) using 29 Si and 1H NMR
in an attempt to understand the water dissolution mechanism in a depolymerized system. The sample
contains about 3.85 wt% H2O and has been synthesized at 1 GPa and 1500C in a piston cylinder
apparatus. Both single-pulse 29 Si and 1 H-29 Si cross-polarization (CP) MAS NMR yield a single,
asymmetric peak near -82 ppm, resembling that of anhydrous diopside glass. The 1 H-29 Si CP time
constant has been found to be 0.39 ms, similar to those of SiOH in other hydrous silicate glasses andcrystals. Thus at least some of the dissolved water are in the form of SiOH. The 1 H NMR spectra(both single pulse and spin-echo) contain two partially resolved peaks at 1.3 and 4.5 ppm. The 4.5
ppm peak is broader and asymmetric with a broad shoulder near 14 ppm. This peak and its shoulder
are similar to those observed in more polymerized hydrous silicate glasses, such as BaSi2O5 , and canbe ascribed to SiOH and/or molecular H2 O with a range of H-bonding strengths. The 1.3 ppm peak isnarrower and somewhat taller than the 4.5 ppm peak. The chemical shift of this peak indicates that itis due to OH groups that are not hydrogen-bonded. Peaks at this position has not been observed in any
Al-free hydrous silicate glasses thus far and are not likely to be due to SiOH or molecular H2 O. On theother hand, MgOH in low-pressure hydrous magnesium silicates, such as talc, and CaOH in calcium
silicate hydrates (CSH) are known to yield similar 1H NMR chemical shifts. We may thus assign thispeak to MgOH/CaOH, suggesting that a substantial part of the dissolved water are in the form of
free OH. Therefore the water dissolution mechanisms in silicate melts/glasses seem to change frommolecular H2O+SiOH in more polymerized systems to free OH groups pus SiOH/molecular H2O inmetasilicate compositions. Any model for water dissolution in depolymerized silicate melts/glasses
must take this fact into consideration.

Session 3100

3.26 Ab initio calculation of the 17 O and 1 H NMR parameters for various OH groups:
Implications to the speciation and dynamics of dissolved water in silicate glasses
X. Xue and M. KanzakiInstitute for Study of the Earth’s Interior, Okayama Univ., Misasa, Tottori 682-0193 Japan;
mkanzaki@misasa.okayama-u.ac.jp

Ab initio molecular orbital calculations have been carried out for silicate, aluminosilicate andaluminate clusters, in order to study the NMR characteristics of various types of hydroxyls (OH)that are possibly present in hydrous silicate glasses and melts. The clusters have been optimized
with the density functional theory (B3LYP/6-311+G(2df,p)), and their NMR parameters calculated
at HF/6-311+G(2df,p). Our calculations suggest that the 17 O and 2 H quadrupolar coupling constants(Cq O , Cq H ) and 1 H chemical shift (δ iH ) of SiOH, AlOH and bridging OH’s (SiOHAl, AlOHAl)
all show good correlation with the O-H and O-H... O distances. The calculated Cq O , Cq H and δ iH
values agree well with those of experimental data for OH groups with similar O-H... O distances incrystalline phases. Our calculation suggests that hydroxyls with stronger hydrogen-bonding tend toyield smaller Cq O , Cq H , and larger δ i H . SiOH and bridging OH’s of comparable hydrogen-bonding
strengths give similar 17 O and 1 H (2H) NMR parameters. AlOH have a tendency not to form strong AlO-H... O type hydrogen-bonding, and thus give relatively large Cq O , Cq H and small δ iH . On the basisof these calculation results, together with information for hydrogen-bonding strengths estimated fromexperimental vibrational spectra and 1 H NMR data, we were able to predict 17 O NMR parameters
for hydroxyls in hydrous silicate glasses. The observed 17 O NMR peaks for silica gel and hydrous
albite glass, that have been attributed to SiOH, are significantly narrower than expected from Cq O ,suggesting that at least some of the SiOH, if present, must be non-rigid. The observed broad 17 O
NMR peaks for hydrous albite and alkali silicate glasses, originally attributed to molecular H2 O,
could equally well be ascribed to rigid hydroxyls with no or weak hydrogen-bonding.

High pressure mineral physics
3.27

101

The effect of sulfur on silicate network structure in sodium silicate glasses
T. Tsujimura(1) , X. Xue(2) , M. Kanzaki(2) , A. Kitakaze(1) and Y. Kudoh(1)(1) Department of Mineralogy, Petrology and Economic Geology, Graduate School of Science, Tohoku
University, Sendai, 980-8943, Japan(2) Institute for Study of the Earth’s Interior, Okayama University, Misasa, Tottori, 682-0193, Japan

When the sulfur contents in silicate melts exceed the saturation limit, an immiscible sulfide melt
segregates from silicate melts. Chalcophile elements partition preferentially into the sulfide melt effecting major changes in the trace and minor element distributions of the silicate melt and possibly
generating metal-sulfide deposits (Naldrett et al., 1978; Ohmoto and Goldhaber, 1997). Thus, sulfur
plays a crucial role in magmatic system although its contents in silicate melts are less than 1 wt.%
(Poulson and Ohmoto, 1990; Tsujimura and Kitakaze, in prep.). Information about the effect of sulfur
on the silicate melt network is therefore necessary in order to understand their physical and chemical properties. 29 Si magic angle spinning nuclear magnetic resonance spectroscopy (MAS NMR)
work has demonstrated that the technique is useful in characterizing silicon coordination environment(Stebbins, 1988; Maekawa et al., 1991). The purpose of this study is to provide constraints on the effects of sulfur on the structure of sodium silicate glasses (quenched melts) using 29 Si and 23 Na MAS
NMR.
Starting glasses of (Na2 O(40) SiO2(60) and Na2 O(30) SiO2(70) ) compositions were prepared fromreagent grade Na2 CO3 and SiO2 at ambient pressure. Mixed starting materials were carefully melted
in Pt crucible at 1000o C for 5 h and then quenched in water. All samples were clear glasses. 0.2
wt. % Gd2 O3 was added in all glasses to reduce the relaxation time of the 29 Si nucleus. For thesulfur-containing glass syntheses, the silicate glass and native sulfur were well mixed and sealed in
an Au capsule with graphite powder. The ratio of the silicate glass to sulfur is 90:10 in weight. The
samples were melted at 1000o C and 1-2 kbar and the quenched to glasses in an Ar-mediated internallyheated pressure vessel with a rapid quenching device at Tokyo Institute of Technology (SMC-2000).
We made MAS NMR measurements with a Varian UNITY-INOVA400NB spectrometer and a 5mm
Jakobsen-type MAS probe at a Larmor resonance frequency of 79.5 and 105.9 MHz for 29 Si and 23 Na
respectively. External chemical shift standards of tetramethyl-silane (TMS) for 29 Si and 1M aqueous
NaCl for 23 Na were used.
For the Na2O(40) SiO2(60) composition, the 29 Si MAS NMR spectrum of the starting, sulfur-freeglass consists mainly of two partially resolved peaks centered at -77.2 and -86.8 ppm, that are attributable to Q2 and Q3 species, respectively. Whereas in the spectrum of the sulfur-bearing glass, theintensity of the Q2 peak decreases markedly. For the Na2 O(30)SiO2(70) composition, the 29 Si MAS
NMR spectrum of the starting, sulfur-free glass consists of a main peak centered at -89.3 ppm, attributable to Q3 species, and a shoulder on the low frequency (more negative chemical shift) side,
attributable to Q4 . Whereas for the spectrum of the sulfur-bearing glass, the relative intensity of theshoulder due to Q4 species increases compared with that of the sulfur-free glass. Changes in the 23 Na
MAS spectra upon sulfur dissolution are subtle. These results suggest that the dissolution of a small
amount of sulfur (< 1 wt.%) into silicate melts result in a marked increase in the degree of polymerization of the silicate network.
References:
Maekawa et al. (1991) J. Non-Cryst. Solids, 127, 53-64.
Naldrett et al. (1978) Can. J. Earth Sci.,15, 715-728.
Ohmoto and Goldhaber (1997) in ”Geochemistry of Hydrothermal Ore Deposits” VIII, 517-611.
Poulson and Ohmoto (1990) Chem. Geology, 85, 57-75.
Stebbins (1988) J. Non-Cryst. Solids, 106, 359-369.

Session 3

102
3.28

Water speciation in basaltic melts: constraints from water solubility measurements
Shigeru Yamashita

shigeru@misasa.okayama-u.ac.jp, Institute for Study of the Earth’s Interior, Okayama University , Misasa,
Tottori 682-0193, Japan

Water is a major volatile component dissolved in magmas at depth, and its presence can promote a significant change in the physicochemical properties of silicate melts. Many spectroscopic
studies have shown that water is dissolved into quenched silicate melts (glasses) as both OH groups
and molecular H2O, through a coupled reaction: (1) H2 O (vapor) = molecular H2 O (melt) and (2)molecular H2 O (melt) + O (melt) = 2OH (melt) (e.g., Stolper, 1982). However, recent advance in theunderstanding of relaxation kinetics of hydrous silicate melts has revealed that the relaxation rates are
so fast that the water speciation measured by spectroscopy in quenched glasses represents, to somedegree, rapid re-equilibration upon quenching (e.g., Dingwell & Webb, 1990).
Here I present results from an indirect approach to investigate the water speciation in basaltic
melts along the solubility surface for total water (both OH groups and molecular H2 O as H2 O) in themelts at magmatic temperature. Basaltic melt (Izu-Oshima volcano, ∼53 wt% SiO2) was equilibrated
with pure water vapor at 1473 K and at pressure between 10 and 80 MPa in an internally heatedpressure vessel, and the total water content in the quenched glass was determined by IR spectroscopy.
This dataset can be modeled adequately on the basis of ideal mixing of quasi-chemical oxygen species
in the system OH (melt) - molecular H2 O (melt) - O (melt). In combination with available volumetricdata of water (Ochs & Lange, 1999), I used an iterative approach to fit the K(2) value (equilibrium
constant of water speciation reaction (2) defined by XOH 2 / (Xmol.H 2O XO )) to the solubility data. The
iteration converged upon ln K(2) = 0.32±0.22, and with this optimum value, the model reproduces
the solubility data over the pressure range studied. The ln K(2) value obtained in basaltic melts is∼1.7 magnitude greater than that obtained by the same procedure in rhyolitic melts (∼77 wt% SiO2)
at 1473K and at pressure between 22 and 100 MPa (ln K(2) = -1.38±0.18, Yamashita, 1999). Thisimplies that a greater fraction of dissolved water has been converted to OH groups in more basic
silicate melt.
References
Dingwell , D. B. & Webb, S. L. (1990). Relaxation in silicate melts. Eur. J. Mineral. 2, 427-449
Ochs, F. A. & Lange, R. A. (1999). Scienece 283, 1314-1317
Stolper, E. M. (1982). Contrib. Mineral. Petrol. 81, 1-17
Yamashita, S. (1999). J. Petrol. 40, 1497-1507.

High pressure mineral physics
3.29

103

Pressure dependence of the emf of thermocouples: Comparison of W-Re and Pt-Rh
thermocouples with P-measurements by in situ X-ray
Tomohiro Iguchi∗, Eiichi Takahashi, Yu Nishihara
Department of Earth and Planetary Sciences, Tokyo Institute of Technology

One of the serious problems in high-temperature-high-pressure studies is the pressure dependenceof the emf of thermocouples. When two different kinds of thermocouples exposed to identical temperatures and elevated pressures, there is a discrepancy in temperature reading. We measured thetemperature discrepancy and the pressure using the in situ X-ray observations. Experiments were
performed using a DIA-type, cubic-anvil apparatus. Experimental results were obtained in the pressure and temperature ranges of P=1.8-10.6GPa and T=26-1000C. W5%Re/W26%Re thermocouples
recorded 12C higher temperature than Pt/Pt13% Rh thermocouples at 10.6GPa and 1000C.

Session 3

104
3.30

Infrared Observation on Low Temperature Phase Transformations of Ice
M. Song1, K. Aoki2, H. Yamawaki2, H. Fujihisa2, and M. Sakashita2

1 Core Research for Evolutional Science and technology Japan Science and technology Corporation,
Kawaguchi, Saitama 332-0012, Japan
2 National Institute of Materials and Chemical Research 1-1 Higashi, Tsukuba, Ibaraki 305-8565, JapanIce exhibits a very rich phase diagram with 12 known crystalline phases and many metastable
and amorphous phases. The high pressure behavior of ice is of fundamental importance to condensed
matter physic and planetary sciences. Many intrguing phenomina of ice have been revealed in recent
years, such as pressure-tuned amorphization, hydrogen-bond symmetrization, Fermi resonance, softmode behavior. In the low temperature and low-pressure region, the physical behavior of ice becomes
extremely unusual. When pressed at 77 K, ice Ih transformed into high-density amorphous (HDA) ice
at 1.0 GPa [1], and then HDA ice will further transform into a ice VII-like phase at about 4 GPa at
80K [2], which is confirmed by a neutron diffraction measurement to be metastable phase quenchable
to ambient pressure below 100K [3]. This particular phase-metastable ice VII exists at temperature
below 180 K in the thermaldynamic stable field of ice VIII [2, 3]. When compressed at below 150 K,
ices II, VI also transit into metastable ice VII [4, 5]. The orientation disorder in metastable ice VII
is speculated to be static [3], completely different from that dynamical one in ice VII. When heatedat ambient pressure, metastable ice VII will transit into LDA ice and then ice Ih [3]. Metastable ice
VII transited into ice VIII upon heating above 2 GPa [2, 3], but Raman measurements of metastable
ice VII up to 30 GPa at 77 K showed no phase transformation [2], although theoretical calculation
predicted that under isothermal compression metastable ice VII should finally transform into ice VIII
[6]. The phase stability of metastable ice VII is still not completely determined. Here we employ
low-temperature infrared absorption technique to study the stability of this phase. A detailed study ofmetastable ice VII may sheed light on our understanding of the structures of ice VII and VIII, andprovides new refinements on the potentials of ice.
Infrared spectroscopic absorption observation of ice was performed in a P-T range of 40-300 K
and 0.2-50 GPa. The pressure of sample was around 3-10 GPa at ambient temperature after sampling.
We cooled down sample to 230-240 K in a cryostat with KBr windows, and then unloaded pressure
to 0.1-1.5 GPa to get ice VI, V, II, Ih. The IR spectra of ice Ih, II, V, VI, VII, VIII were carefully observed and cpmpared with previous studies. Both ice II, VI were successfully cooled down below 150
K. When compressed at temperature below 120 K, ices II and VI transformed into metastable ice VII
probably due to mechanical instability as ice Ih-HAD transformation [7]. The metastable ice VII show
an IR absorption profile almost same as that of ice VII at ambient temperature. Under isobaric heating,metastable ice VII transits in to ice VIII and shows a general trend to decrease the phase transition
temperature by about -2 K/GPa from about 170K at 6GPa to 120 K at 24 GPa. This confirms that theintensification of dipole-dipole interaction in ice contribute to the reorientation of water molecules inmetastable ice VII to form an orientationally-ordered antiferroelectric configuration of ice VIII. Under isothermal compression below 100 K, metastable ice VII shows a temperature-insensitive phasetransition into ice VIII at 30 GPa. The mechanisms of ice II, VII-metastable VII-VIII remains to befurther studied by theoretical calculation and experiment.
References[1] O. Mishima, L. D. Calvert & E. Whalley, ’Melting ice’ I at 77 K and 10 kbar: a new method ofmaking amorphous solids. Nature 310, 393-395 (1984).
[2] R. J. Hemley, L. C. Chen & H. K. Mao, New transformations between crystalline and amorphous
ice. Nature 338, 638-640 (1989).
[3] S.Klotz, J.M.Besson, G. Hamel, R.J. Nelmes, J.S. Loveday & W.G. Marshall, Metastable Ice VII
at Low Temperature and Ambient Pressure, NATURE 398, 681 (1999).

High pressure mineral physics

105[4] J. L. Finney, C. Lobban, S. Klotz & J. M. Besson, In Collected Abstracts of the XVII Congress
and General Assembly (ed. Griffin, J. F.) C-534 (IUCr, Seatle, 1996).
[5] Finney, J. L., Lobban, C., Klotz, S., Besson, J. M. & Hamel, G. in ISIS 96-Rutherford Appleton
Laboratory Report RAL 96-050 (ed. Taylor, A. D.) Vol. A17 (Chilton, UK, 1996).
[6] J. S. Tse, and M.L. Klein, Phys. Rev. Lett. 58, 1672-1675 (1987).
[7] J.S. Tse and D.D. Klug, C. A. Tulk, I. Swainson, E. C. Svensson, C.-K. Loong, V. Shpakov, V. R.
Belosludov, R. V. Belosludov & Y. Kawazoe, The mechanisms for pressure-induced amorphization of ice Ih. Nature 400, 647 (1999).

October 4 (Thursday) 13:30 – 16:30

Session 4

Dehydration of subducting slabs and water storage inthe mantle

Keynote: Max Schmidt, CNRS, Clermont-Ferrand/ETH Zürich
Chair: Eiji Ohtani, Tohoku University

107Dehydration of subducting slabs and water storage in the mantle
4.1

109

Dehydration vs. melting vs. dissolution: Where and how to generate a mobile component
during subduction
Max W. Schmidt
Institute for Mineralogy and Petrology, ETH, Sonneggstr.10, Zuerich 8092 Switzerland;max.schmidt@erdw.ethz.ch

Prograde metamorphism of subducting lithosphere causes a series of mineralogical reactions
which commonly result in a mobile phase. The mobile phase is either a low-density fluid or a highdensity, possibly supercritical fluid or a silicate or carbonatite melt. This contribution reviews thegeneration (conditions and amounts) of the various mobile phases and also examines the restite(s) tobe subducted to large depths.
Somewhat simplifying, 5 different regimes can be recognized to depth-equivalents of 10 GPa:(i) Low to medium P, low T (< 650 C, < 2.5 GPa) where hydrous phases are abundant and dehydration reactions are often perpendicular to typical subduction P-T paths, thus, dehydration
rates are high. Any subducted lithosphere will go through this first stage.
(ii) Low to medium P, high T where amphibole and biotite are the principal hydrous phases to melt
at fluid absent conditions. Adakitic magmas (slab melts, Na>K) are likely to origin at suchconditions (1-2.5 GPa, 800-1000 C).
(iii) Medium to high P, low T (500-800 C, 2.5-10 GPa) where hydrous phases are already largely
reduced in volume and dehydration reactions are often subparallel to possible P-T paths. In thisrange fluids become increasingly rich in dissolved matter.
(iv) Medium P, high T (> 700-800 C, 2.5-5 GPa) where phengite is the principal hydrous phase to
melt and melts have K>Na.
(v) High P, high T (>800 C, >5 GPa) where the solvus between fluid and melt is closed and a supercritical fluid might dissolve hydrous phases. Above 10 GPa (8 GPa in metasediments) phengite
decomposes to the anhydrous phase K-hollandite, however at such pressures potassium could
already be extracted by melting or by supercritical fluids.
Most dehydration and decarbonatization reactions are continuous and, in a given bulk composition, result in a combination of continuous and stepwise devolatilization. In the real world, the following factors contribute to the continuous nature of the devolatilization signal: highly variable bulk
compositions in the sedimentary and mafic layers as well as possibly different degrees of depletionin the peridotitic layer(s); inhomogeneous distribution of carbonates vs. hydrates resulting in a inhomogeneous X(CO2) in the fluid phase, high temperature gradients within the subducting lithosphere,
and finally, possible kinetic hindering of reactions and thus widening of reaction zones.
Contrasting the above petrological effects, a mechanism of fluid/melt ascent through channeledflow or hydrofractures could focus a broadly distributed fluid. These latter mechanisms are largely in
discussion.
Although the exact pressures and temperatures for the closure of the solvus between fluid and melt
are poorly defined, there is little doubt that at high pressures the solvus closes and the mobile phasebecomes supercritical. Especially potassium may entirely dissolve in small quantities of fluid and it isdoubtful that much of the potassium could reach large depths. The transport properties of supercriticalfluids for key trace elements is expected to differ significantly from low-pressure, low-density fluids
but little experimental data are available.

110

Session 4

Carbonates present within the subducting oceanic crust influence dehydration reactions but contrary to what is intuitively believed, often enlarge the stability field of hydrous phases. This is especially true in mafic and pelitic compositions where melting reactions are generally shifted to higher
temperatures in the presence of carbonates. Carbonates may remain stable in a siliceous melt andwith increasing temperature form a carbonatite melt coexisting with a silicate melt. At low temperatures fluids from carbonate-saturated lithologies are generally CO2 -poor (especially with magnesite
or dolomite X(CO2 ) is mostly < 15%) and the only efficient decarbonatization processes are either
flushing with aqueous fluids from below or melting at relatively high temperatures.
The layered structure of an oceanic lithosphere originating from fast spreading ridges (i.e. theentire circumpacific and thus most actual subduction zones) and temperature gradients within the
subducting lithosphere define the relation between amount and depth range of dehydration in themafic and ultramafic layer and possible melting of the sedimentary and mafic layers. In particular, thelayered structure causes fluid:rock ratios in the sedimentary layer (as fluids from the underlying mafic
and peridotitic layers must rise through the sediments) of >1:1 (integrated from 10 to 70 km depth)
before melting temperatures are reached. Thus, some of the trace elements commonly considered to
be only efficiently mobilized by melts (e.g. Be, Th) can also be quite effectively mobilized by fluids.
The layered structure is likely to cause significant amounts of fluid to be present in the sedimentarylayer to ca. 70 km depth. To this pressure, fluid-present sediment melting is likely to dominate overfluid-absent sediment melting, however at larger depths, dehydration below the sediments is generally
insufficient to saturate significant amounts of sedimentary melt.

Dehydration of subducting slabs and water storage in the mantle
4.2

111

Circulation and Redisribution of Water in the Earth’s Interior
Eiji Ohtani

Institute of Mineralogy, Petrology, and Economic Geology, Tohoku University, Aza-Aoba, Aramaki, Aoba-ku,
Sendai 980-8578 Japan; ohtani@mail.cc.tohoku.ac.jp

Water stored in descending slab is transported into the depths greater than 200 km when the slab
temperatures are lower than c.a. 600◦ C at around 6 GPa (e.g., (1)). In cold slabs with local water
enrichment greater than 1 wt.% of the bulk rock, hydrous wadsleyite and hydrous ringwoodite are
the main water reservoirs and excess water is stored in superhydrous phase B Mg10 Si3O14 (OH)4in the transition zone. Phase Egg, AlSiO3OH is also a candidate for water reservoir in the sediment component in slabs in the transition zone. Our high pressure experiments revieled that a decomposition reaction of superhydrous phase B into phase G(=D) and periclase; Mg10 Si3 O14 (OH)4 =(5/3)Mg1.2Si1.8 H2.6 O6+8MgO at 30 GPa and 1000∼1100◦C, and a decomposition of phase Egg intodelta-AlOOH and stishovite, AlSiO3OH = AlOOH+SiO2 at 23 GPa and about 1000◦C. delta-AlOOH
is a new high pressure polymorph of AlOOH with a structure similar to delta-FeOOH (2). No water
release occurs by these reactions. Hydrous phases, phase G and delta-AlOOH, are not stable under thenormal geotherm conditions, and are stable only in the cold slabs descending into the lower mantle.
These phases can play a crucial role for water transport into the lower mantle by slab subduction.
We can expect two dehydration zones in the descending slabs; a shallow dehydration zone in the
200km depth, and a deep dehydration zone in the depths around 900∼1200 km, where dehydrationis originated from decomposition of superhydrous phase B, phase G, and delta-AlOOH. There is alayered structure in the water storage potential along the normal mantle geotherm; i.e., the upper andlower mantles have a relatively low water storage potential (e.g., (3)), whereas wadsleyite and ringwoodite in the transition zone have a larger potential. Thus, the fluid formed in the deep dehydration
zone will move upwards and will be stored in the transition zone.
We have clarified the boundary the reaction 3Fe+H2O–>FeO+2FeH using the diamond anvil cell.
FeO and FeH is stable at temperatures above 1000◦ C at 35 GPa, and the phase boundary of the reaction
has a positive slope. The reaction indicates that iron hydrite FeH is important during the accretional
stage of the Earth. We observed that Fe and water is stable at relatively lower temperatures and higher
pressures, whereas FeO and FeH are stable at higher temperarures. This observation indicates thatiron and water reacted to form FeO and FeH in the hot accreational stage of the early Earth, and FeH
can be transported into the core by the reaction (e.g., (4,5) ). On the other hand, during cooling stageof the core, dissolved FeH and FeO in the core can react to form Fe and water. H2 O may be formedby this process at the core mantle boundary.
Thus two processes, i.e., water transport from ocean to the mantle by slab subduction and supplyof water from the core to the mantle provide re-distribution of water from the ocean and core to themantle transition zone during secular cooling of the Earth.
References(1) T. Kawamoto et al. (1996). Earth Planet. Sci. Lett., 142, 587-592.
(2) A. Suzuki et al. (2000). Phys. Chem. Minerals, 27, 689-693.
(3) Bolfan-Casanova et al. (2000). Earth Planet. Sci. Lett., 182, 209-221.
(4) Fukai, Y. (1992). in High-Pressure Research: Application to Earth and Planetary Sciences, edited
by Y. Syono and M.H. Manghnani, pp. 373-385, TERRABUB/AGU, Tokyo/Washington D. C..
(5) Okuchi, T. (1997). Science, 278, 1781-1784.

Session 4

112
4.3

An experimental study on the Mg-Al granulites and production of granitic melts in
K2 O-MgO-Al 2O3-SiO2-H2O system
Ganesha Atni Venkatramiah, J.A.K. Tareen
Department of Geology, University of Mysore, MYSORE-570 006, INDIA<ganmysore@hotmail.com>

Ultrahigh Temperature (UHT) crustal metamorphism has been considered as that type of metamorphism which occurs at P-T conditions of more than 7 kb and 900◦C respectively. The key minerals
represent this UHT metamorphism includes sapphirine, spinel, cordierite, orthopyroxene, sillimanite,garnet, osumilite and the worldwide occurrence of these has been discussed in detail [1]. Experimental work on dehydration melting of biotite in natural system shows that large amount of granitic melt
produces depending on the starting bulk composition [2-5]. The stability of biotite in KFMAS(H)system has been studied experimentally [6-9] and hitherto there are no reports on the dehydration
melting of phlogopite in Fe-free KMASH system above 950◦C/6 kb. We report here the compositionof melts along with restites obtained by vapour-absent melting of phlogopite in the P-T range of 6-9
kb and 1000-1250◦C respectively.
All experiments were carried out in KMASH system using bulk composition 2Phl-6Sil-9Qtz*,
corresponding to Mg-Al granulite. The starting mineral phases used were synthetic ideal phlogopite,sillimanite and natural quartz in the ratio of 2:6:9 respectively. Experiments were conducted using
Internally Heated Pressure Vessel under vapour-absent condition and the water available in the system
was solely from the structural water of phlogopite. Table-1 lists the bulk composition, experimental
details, mineral phases encountered and composition of melts.
The major mineral phases obtained in the given P-T range are Spr-Crd(H)-Al.En-Os. Sapphirine
is of 221 type and enstatite is highly aluminous (up to 14 wt% of Al2 O3). Interestingly, an hexagonalcordierite instead of orthorhombic cordierite was observed as new phase. The melts obtained is highly
silicic (up to 75 wt% of SiO2) and overall chemical analysis are comparable and consistent with thereported melts [2,4]. The melts of the present study show positive trend in MgO, Al2 O3 and negative
trend in SiO2 with respect to increasing temperature. Further, there is an inverse correlatin between
MgO and Al2 O3 vs SiO2. The clustering and overlapping of melts obtained and that of the reported
melts [2,4] on (Al2O3+K2 O)-MgO-SiO 2 triangular diagram (Fig-1) indicates that ’granitic melts’would be produced from pelitic rocks irrespective of XMg in biotite. But, the quantity of melt andrestites at given P-T condition definetely varies with the bulk composition and water in the system.
*Abbreviations are after Kretz (1983), Os-Osumilite
Table-1. Experimental details and composition of melts and bulk.
Expt. Temp (◦ C)
Duration Product
MgO
No.
Pressure (kb) (Hours)
K2 O
BULK COMPOSITION: 2Phl-6Sil-Qtz (+H2 O-1.27)
4.01
48
49
26B
40
35
43
37A

1150/7.40
1050/8.00
1190/9.00
1200/6.00
1050/6.00
1200/7.00
1250/6.00

24
46
24
34
24
64
24

Spr-Crd(H)-En-Os-L
Spr-Crd(H)-Sil-En-L-Os
Spr-Crd(H)-En-Os-L
Spr-Crd(H)-En-Os-L
Spr-Crd(H)-Sil-En-Qtz-L
Spr-Crd(H)-En-Os-L
Spr-Crd(H)-L

7.34
5.62
7.04
6.81
7.24
8.34
5.61

Oxide%Al2 O3 SiO2 Total
9.81 30.29 53.97
Melt Composition
1.75 13.77 72.00
2.55 14.91 73.46
2.06 13.88 75.13
2.82 14.61 73.31
2.59 14.94 70.42
1.86 14.26 73.48
4.85 15.93 69.95

99.35
94.86
96.54
98.11
97.55
95.19
97.94
96.34

Dehydration of subducting slabs and water storage in the mantle

113

Fig-1: (Al2 O3 -K2 O)-MgO-SiO2 diagram showing starting bulk composition and experimental melts. B1:
Starting bulk of the present study. B2, B3 and B4 are bulk used for experimentation by previous workers[2,4,7]. L3 denotes melts of present study. M1 and M2 are melts of previous workers in the natural systems[2,4]. Note the overlalpping of melts [2,3] on the present study in the KMASH system.

References[1] Harley, S.L, Geol.Soc.London Spec.Publ., 138, 81-107 (1998)[2] Vielzeuf, D. and Holloway,J.R., Contrib.Mineral.Petrol., 98, 257-276 (1988)[3] Patino Douce,A.E. et al., Earth Plan.Sci.Letter, 97, 290-315 (1990)[4] Patino Douce,A.E. and Johnston,A.D., Contrib.Mineral. Petrol., 107, 202-218 (1991)[5] Rutter,J.M. and Wyllie,P., Nature, 331, 159-160 (1988)[6] Carrington,D.P. and Harley, S.L., Contrib.Mineral.Petrol., 120, 270-291 (1995a)[7] Carrington,D.P. and Harley, S.L., J.Met.Geol., 13, 613-625 (1995b)[8] Audibert,N. et al., J.Met.Geol., 13, 331-344 (1995)[9] Motoyoshi et al., Eur.J.Mineral., 5, 439-440 (1993)[10] Kretz, R., Am., Mineral., 68, 277-279 (1983)

114

Session 4

4.4

Earthquakes at depth, their enabling mineral reactions, and transport of H2O into thedeep mantle
Harry W. Green, II

Institute of Geophysics and Planetary Physics and Department of Earth Sciences, University of California,
Riverside, CA 92521 USA<harry.green@ucr.edu>

Over the last decade, extraordinary progress has been made in visualizing the interior of the earth
using seismic wave speeds and tomographic inversion techniques. However, we still know relatively
little about the mechanism(s) by which earthquakes can be triggered and propagate at depth. Only
two mechanisms have been experimentally verified as potential earthquake trigger mechanisms at
high pressure: dehydration embrittlement and transformation-induced faulting. During this period,
many aspects of the latter have been worked out in the laboratory, hence reasonably stringent tests on
the viability of this mechanism can be made. However, the former mechanism has been little studied
since its discovery in the 1960s; discussions of its potential importance as an earthquake mechanism
have been qualitative and uncertain. Both mechanisms require enabling mineral reactions under stressthat produce local pockets of low-viscosity ”fluid” that lead to mechanical instability. Dehydration
embrittlement, potentially capable of explaining all earthquakes at depth, requires hydrous phasesthat dehydrate under conditions appropriate to explain the continuous distribution of earthquakes withdepth in subduction zones. Only recently have candidate high-pressure hydrous phases been identified
and their stability fields mapped out but still only in simple systems and whether hydrous phases exist
to sufficient depth in subducting lithosphere is unknown. Similarly, transformation-induced failure canonly function under conditions of metastable transformation of olivine to wadsleyite and ringwoodite
in the mantle transition zone, yet evidence for the presence of metastable olivine in the cores ofsubduction zones has been equivocal. Therefore, it is not clear whether or not the necessary enablingmineralogy for either mechanism is present under all conditions where it might operate.
Several subduction zones have clearly defined double seismic zones at intermediate depths, withthe lower zone located as much as 40 km below the top of the slab. Hydration of the oceanic lithosphere at ridges is unlikely to extend below 10-12 km and only a very small percentage of xenoliths
from oceanic islands contain even small crystals of hydrous phases. Therefore, if the lower seismiczone is a result of dehydration, there must be a mechanism operating at or near oceanic trenches that
introduces water to sufficient depth. Although there is no direct evidence of such a mechanism, it has
been proposed that fault zones of great normal faults outboard of trenches might provide a pathway
for introduction of water and consequent growth of hydrous phases. Moreover, the correlation is now
quite striking between the stability field of antigorite, the hydrous phase most likely to be present at
depth in the oceanic lithosphere, and models of sub-duction zones. In particular, the convergence of
double seismic zones at 200-250 km is readily fit by thermal models to the breakdown of antigorite toolivine + enstatite + H2O. This evidence, although indirect, is a powerful suggestion that these earthquakes may be triggered by this reaction. If antigorite does exist at least discontinuously distributedin the subducting lithosphere and if the cold core of the subduction zone remains at a temperatureof less than ∼600 ◦ C at 200 km, then the part of the slab colder than that will pass into the stability
field of Phase A, the first of a succes-sion of increasingly dense hydrous phases. However, the earthquake generating power of these phases is questionable because all of the reaction boundaries likely
to be crossed by a particle path in a subduction zone result in increasingly hydrous phases, hence
fluid could be released only transiently during the reaction, an unlikely process. The disappearanceof double seismic zones at ∼250 km and the very low frequency of all earthquakes at depths of 300400 km may be a consequence of this fact. Thus, the combination of the phase equilibria of hydrous
peridotite and the presence of earthquakes deep in slabs at intermediate depths suggest that H2O ispresent in many subducting slabs and that however much H2O makes it to 200 km is likely to beDehydration of subducting slabs and water storage in the mantle

115

carried undisturbed into the deep mantle.
New evidence that may indicate extensive preservation of metastable olivine in the transitionzone of the Fiji region (Chen and Brudzinski, Science, 2001) reinforces the transformation-induced
faulting hypo-thesis for triggering of earthquakes below 400 km. However, in hydrous slabs thermal
consequences of transformation-induced faulting, perhaps enhanced by shear heating, could raise thetemperature in the growing fault/slip zone and activate brucite dehydration at depths greater than 500
km. Thus, earthquakes below 500 km could be composite, potentially explaining the very large very
deep earthquakes about which there has been controversy over the last several years. This two-stepprocess for these earthquakes would also be consistent with the observations that they occur in warmerslabs and are frequently double events, with a smaller event preceding the larger event. In this model,
even though the Phase D + brucite dehydration boundary extends to pressures appropriate for the
lower mantle, earthquakes would stop at the base of the transition zone because the transformationinduced trigger would cease to operate at that boundary.
The model of composite earthquakes is testable in the laboratory whereas the inferred presence ofhydrous phases in shallow subducting lithosphere should be testable by seismological methods.

Session 4

116
4.5

Slab dehydration and earthquakes: Seismological evidence
Dapeng Zhao

Geodynamics Research Center, Ehime University, Matsuyama 790, Japan<zhao@sci.ehime-u.ac.jp>

A growing body of evidence suggests that fluids (water, magma, and gases) are intimately linked
to a variety of earthquake faulting processes. These include the nucleation, propagation, arrest andrecurrence of earthquake ruptures, fault creep or slow earthquakes, and the long-term structural andcompositional evolution of fault zones. In this report I present seismological evidence for the influences of fluids and magma on the generation of large earthquakes in the crust and the subducting
oceanic slabs under the Japan Islands. The relationship between seismic tomography and large crustal
earthquakes (M 5.7-8.0) in Japan during a period of 116 years from 1885 to 2000 is investigated and it
is found that most of the large crustal earthquakes occurred in or around zones of low seismic velocity.
The low-velocity zones may represent weak sections of the seismogenic crust. The crustal weakening
is closely related to the subduction process in this region. Along the volcanic front and in back-arc
areas, the crustal weakening may be caused by active volcanoes and arc magma resulting from theconvective circulation process in the mantle wedge and dehydration reactions in the subducting slab.
In the forearc region of southwest Japan, fluids are detected in the 1995 Kobe earthquake source zone,
which may have contributed to the rupture nucleation. The fluids may originate from the dehydrationof the subducting Philippine Sea slab. The recent 2001 Geiyo earthquake (M 6.8) occurred at 50 km
depth within the subducting Philippine Sea slab, and it may be also related to the slab dehydration
process. A detailed 3-D velocity structure is determined for the northeast Japan forearc region using
data from 598 earthquakes that occurred under the Pacific Ocean with hypocenters well located withsP depth phases. The results show that strong lateral heterogeneities exist along the slab boundary,which may represent asperities and results of slab dehydration and may affect the degree and extent of
the interplate seismic coupling. These results indicate that large earthquakes do not strike anywhere,
but only anomalous areas which may be detected with geophysical methods. The generation of a large
earthquake is not a pure mechanical process, but is closely related to physical and chemical propertiesof materials in the crust and upper mantle, such as magma, fluids, etc.
References
Zhao, D., H. Kanamori, H. Negishi, D. Wiens, 1996. Tomography of the source area of the 1995 Kobe
earthquake: Evidence for fluids at the hypocenter? Science 274, 1891-1894.
Zhao, D., H. Negishi, 1998. The 1995 Kobe earthquake: Seismic image of the source zone and its
implications for the rupture nucleation. J. Geophys. Res. 103, 9967-9986.
Zhao, D. T. Mizuno, 1999. Crack density and saturation rate in the 1995 Kobe earthquake region.
Geophys. Res. Lett. 26, 3213-3216.
Zhao, D., F. Ochi, A. Hasegawa, A. Yamamoto, 2000. Evidence for the location and cause of large
crustal earthquakes in Japan. J. Geophys. Res. 105, 13579-13594.

Dehydration of subducting slabs and water storage in the mantle
4.6

117

Transportation of H2O and melting in subduction zones
Hikaru Iwamori

Department of Earth and Planetary Sciences, University of Tokyo

Abstract
Recent knowledge concerning the phase relationships of hydrous peridotitic and basaltic systems
allows us to model fluid generation and migration in subduction zones. Here I present (1) numericalmodels for the transportation of H2O and melting in subduction zones, in which generation and migration of aqueous fluid, its interaction with the convecting solid, and melting are considered, based on
the phase relationships, (2) predictions of the corresponding seismic structures based on the calculated
distribution of the fluids (aqueous fluids and melts), (3) comparison between the model predictions
and 3-D seismic tomographic images beneath the Japanese islands, and (4) analyses on distributionof the volcanoes and the volcanic chains.
The model calculation suggests that in northeast Japan, nearly all the H2 O expelled from the subducted Pacific plate is hosted by serpentine and chlorite just above the plate, and is brought down to∼150 km. Breakdown of serpentine and chlorite at these depths results in the formation of a fluid column through which H2 O is transported upwards, and results in the initiation of melting in the mantle
wedge beneath the backarc (Iwamori, 1998). The seismic tomographic studies suggest the existence ofsuch a melting region beneath the backarc (Iwamori and Zhao, 2000). In central Japan, the subducted
Philippine Sea plate overlaps the subducted Pacific plate. This geometry causes slow thermal recoveryof the subducted Pacific plate, resulting in dehydration reactions at levels (200-300 km) deeper thanin northeast Japan, and deflection of the volcanic chain towards the backarc side (Iwamori, 2000). The
model also suggests that, beneath central Japan, a part of H2 O subducted is carried to depths greaterthan 300 km by phase A. In contrast, in southwest Japan, where a relatively hot part of the Philippine
sea plate (Shikoku basin) subducts, the dehydration reactions are predicted to occur at relatively shallow levels (<100 km depth). The seismic tomographic image supports well the predicted distributionof the fluids beneath the volcanic front to the forearc region (Zhao et al., 2000). These comparisons
between the model predictions and the observations suggest that the thermal structure (determined byage and subduction velocity) of the subducting plate strongly controls the distribution of the aqueous
fluids and melts in subduction zones through the position of the dehydration reactions.

References: Iwamori, EPSL 160, 65-80, 1998; Iwamori, EPSL 181, 41-46, 2000; Iwamori, EPSL 181,131-144, 2000; Iwamori and Zhao, GRL 27, 425-428, 2000; Zhao et al., GRL 27, 2057-2060, 2000.

118

Session 4

4.7

Apatite concentration in the Finero phlogopite-peridotite, Italian Western Alps, as thelate product of subduction-related mantle metasomatism
Tomoaki Morishita∗ , Shoji Arai, Akihiro Tamura∗ Earth Sciences, Kanazawa University

Introduction
The Finero phlogopite-peridotite is though to be an example of highly metasomatized peridotite
by slab-derived melts (or fluids) (Zanetti et al., 1999; Prouteau et al., 2001). The Finero phlogopiteperidotite is, therefore, very important to understand evolution of mantle wedge due to interaction
with slab-derived melts or fluids. We found apatite concentrations in the Finero phlogopite-peridotite.
Here we present occurrence of apatite and other petrographical characteristics, and discuss the originof the apatite concentrations in the context of mantle metasomatism on a mantle wedge.
Petrography
The phlogopite-peridotite consists of harzburgite and dunite with small amounts of chromitite andclinopyroxenenite. Apatite and carbonate are randomly distributed in the pholopite-peridotite (Zanetti
et al., 1999). Apatite concentration was found in thin (< 1 cm in thickness) layers being relatively
white in color. The host peridotite has a granoblastic texture consisting mainly of olivine, amphibole,orthopyroxene and spinel with minor amount of clinopyroxene, phlogopite and sulfide minerals. Veryweek foliation is recognized in the host peridotite by elongation of amphibole. The thin layers areslightly oblique to the foliation of the host.
The thin layers consist of fine-grained minerals (a few to 100 µm in size), olivine, orthopyroxene,
amphibole, spinel, phlogopite, apatite, carbonate and sulfide minerals. Each apatite grain is 10-100µm in size and concentrates as veins parallel to the thin layers. Apatite concentration exceeds morethan 10 modal % in locally. Carbonate minerals are commonly found in the apatite-rich vein. The
thin layers are rich in fine-grained orthopyroxene (< 50 µm). Distinctive low-Al, Ca orthopyroxene
locally occurs as a thin film, < 10 µm in width, to be network at the grain boundary of fine-grained
olivine grains (50 µm ). Orthopyroxenes characterized by low Al contents were also reported from
apatite-bearing domains in the Finero (Zanetti et al., 1999). Cr-spinels in the thin layers are chemicallyheterogeneous within each grain (Cr2O3 & Fe2 O3-poor part and Cr2O3 & Fe2O3 -rich part).
Discussion
The presence of low-Al orthopyroxenes was reported in peridotite xenoliths, which are thought tobe subarc upper mantle metasomatised by slab-derived fluids (Blatter & Carmichael, 1998; Smith etal., 1999, Arai & Kida, 2000; McInnes et al., 2001). The thin layers with apatite concentration wasformed through a very local reaction between peridotite with (CO2 + H2 O)-bearing melts (or fluids),
probably derived from a subducted slab. Composition of localized melts (or fluids) may evolve in local
situation where amphibole or phlogopite are precipitating with loss of K2O, Al2O3 and increasing of
CO2 /H2 O ratio from the primary melts (or fluids), leading to be unmixing of SiO2-rich and CO2rich melts (or fluids) (Zanetti et al., 1999). Apatite and carbonate were precipitated from CO2 -richvarieties. Crystallization of orthopyroxene, representing addition of silica to mantle peridotite is amechanism for metasomatic alteration of mantle wedge peridotite.
Reference
Arai, S. & Kida, M. (2000). The Island Arc, 9, 458-471.
Blatter, D. L., & Carmichael, I. S. E. (1998). Geology, 26, 1035-1038.

Dehydration of subducting slabs and water storage in the mantle

119

McInnes, B. I. A., Gregoire, M., Binns, R. A., Herzig, P. M. & Hannington, M. D. (2001). Earth
Planet. Sci. Lett., 188, 169-183.
Prouteau, G., Scaillet, B., Pichavant, M., Maury, R. (2001). Nature, 410, 197-200.
Smith, D., Riter, J. C. A. &Mertzman, S. A. (1999). Earth Planet. Sci. Lett., 165, 45-54.
Zanetti, A., Mazzucchelli, M., Rivalenti, G. & Vannucci, R. (1999). Contrib. Mineral. Petrol., 134,
107-122.

Session 4

120
4.8

High internal pressure retained in fluid inclusions of mantle-derived materials
Hiroyuki Kagi
Laboratory for Earthquake Chemistry, Graduate School of Science, University of Tokyo

Mantle-derived geological samples give us direct information on the physicochemical state of
fluid of mantle origin. In this talk, our recent studies on the excess internal pressure of fluid inclusions
in natural diamond and in mantle xenolith will be introduced.
Natural minerals of high-pressure origin have the potential to retain high internal pressures on materials trapped within these crystals. Inclusions in natural diamonds, as demonstrated from the blue
shift in vibrational spectra, retain several gigapascals of high pressure (Kagi et al. (2000), Min. Mag.,64, 1089-1097, and references therein.). The internal pressure retained in inclusions gives us a clue
to decipher the depth where the inclusions were trapped. Some kinds of natural diamonds contain“mantle-derived” fluid inclusions. It is essential problem to discuss the depth (pressure) where thefluids were trapped for understanding chemistry of mantle-fluids. We investigated NIR-IR spectroscopic properties of water contained as fluid inclusions of cuboid diamonds. The results indicatedthat a cuboid diamond contained two kinds of water; ice VI retaining approximately 1.9 GPa and liquid water at ambient pressure. Recently, we have measured mapping analysis of infrared absorption
spectra on the thinly sliced cuboid diamonds. The preliminary results demonstrated that the highpressure bearing fluid distribute heterogeneously in the crystal suggesting that the diamond trapped
mantle-fluid at different growth stage of diamond. Natural diamonds trap fluids at mantle depth (pressure) where diamond is thermodyamically stable, but some part of fluid might be trapped at shallower
depth during the growth stage of the diamond.
Internal pressure of fluid inclusions can serve as a sensitive strength-marker for mantle minerals.
Recently, we investigated fluid inclusions in mantle xenolith from Far Eastern Siberia (Yamamoto etal., EPSL submitted) by micro-Raman spectroscopy and microthermometry. The samples contained
pure CO2 inclusions and melt inclusions. We measured Raman spectra for each CO2 inclusion after
separating to mineral species. Two lines are observed approximately at 1286 cm−1 and 1388 cm−1 onRaman spectra of CO2. Both the intensity ratio (R) and energy difference (∆) between these two peaks
are known to show pressure dependence. The changes in two parameter with pressure are very subtle for natural samples, however, Raman barometry for CO2 inclusions have not been established forgeological purpose. We determined R and D precisely by applying the numerical resolution enhancement technique, then Raman barometer was calibrated with conventional homogenized temperatureusing cryogenic microthermometry. The results indicated that retained pressure in the CO2 inclusions
showed surprisingly systematic gradation dependent to mineral species; sp>opx∼cpx>>ol. The difference in the pressure reflects the deformations of minerals at mantle condition after trapping fluidinclusions. This is a direct observation of strength of mantle minerals to plastic deformation which
was projected onto the retained pressure of fluid inclusions.

Dehydration of subducting slabs and water storage in the mantle
4.9

121

In situ observation of supercritical behaviors between H2O fluids and subduction zone
magmas.
Tatsuhiko Kawamoto and Kyoko Matsukage
Institute for Geothermal Sciences, Kyoto University, Beppu, 874-0903, Japan,
kawamoto@bep.vgs.kyoto-u.ac.jp, matsukage@bep.vgs.kyoto-u.ac.jp

Visual observation of melting relationships between H2 O fluids and natural andesitic and rhyolitic
melts were conducted in a range of pressure from 0.4 to 1.8 GPa and temperature up to 1050 degree
C. The experiments were conducted with a Bassett type externally heated diamond anvil cell (Bassett
et al., 1993, Rev. Sci. Inst.). Starting materials were the Fuji 1707 andesite and rhyolite glasses thatare geochemically characterized by calc-alkaline rock series and medium K andesites, and represent
typical felsic magmas erupted in subduction zones. Supercritical behavior was observed in a range ofchemical composition between andesite (62 wt. % SiO2) and rhyolite (70 wt. % SiO2). The available
experimental data suggest a higher supercritical temperature at equivalent pressure between andesite
and H2 O fluids than that between rhyolite and H2 O fluids. This is in contrast to the suggestion by
Bureau and Keppler (1999, Earth Planet. Sci. Lett.) based on their data in the system of H2 O fluid andnepheline (39.4 wt. % SiO2), jadeite (58.8 wt. % SiO2), and haplogranitic melts (73.6 wt. % H2 O)
with two exceptions of albite (65.3 wt. %) and Ca bearing haplogranitic melts (71.6 wt. % SiO2 ).
Chemical fractionation between H2O-rich fluids and silicate-rich melts with respect to silicatecomponents were determined by electron microprobe analysis of silicate glasses quenched fromsilicate-rich melts. Na2 O and K2O are likely to enter the H2O-rich fluid and CaO, MgO and FeO
are preferentially partitioned into silicate-rich melt. The same preferential distribution of Na2 O wasalso observed between albite melt and H2O fluids and in this case SiO2 is preferentially partitioned
in the H2 O-fluid. Chemical transport of Na2O and K2 O by H2 O-rich fluid in H2 O-saturated magma
chamber was suggested based on dynamic degassing experiments in andesite system (Sakuyama andKushiro, 1979, Contrib. Mineral. Petrol.). The knowledge of partitioning coefficients of trace elements
between H2 O fluid and silicate melts can be investigated with the present method.

Session 4

122
4.10

Geochemical Consequences of Deep Subduction of Oceanic Crust: Insights from
Melting Experiments on Hydrous Metabasalt at 3-16 GPa
Robert P. Rapp1 and N. Shimizu

1 Department of Geosciences and Mineral Physics Institute, State University of New York at Stony Brook, ESS

Building, Stony Brook, New York 11794-2100 USA

Subduction transports hydrated, geochemically-altered basaltic oceanic crust through the islandarc and back-arc regimes, into and sometimes through the mantle transition zone at 660 km depth, tobe incorporated into the deep mantle source region for so-called ’HIMU’ ocean island basalts, suggesting that crustal recycling occurs on a whole-mantle scale. Along the way, breakdown of hydrous
minerals in the slab release fluids which contribute to island arc volcanism and metasomatism of the
sub-arc and back-arc mantle, while transforming the slab to a large-ion lithophile (LILE) elementdepleted, ultrarefractory crystalline residue. Experimental studies have shown that minerals such as
phengite and lawsonite can be stable in basaltic compositions up to pressures of 10-11 GPa (Schmidt,
1996); given the broad diversity in the dynamics of subduction, a broad range in slab geotherms ought
to be expected, so that fluid release associated with the breakdown of hydrous minerals can take placeat depths up to 300 km (e.g., “hot slabs” will undergo dehydration at lower pressure than “cold slabs”.

Figure 1. Primitive-mantle normalized trace element abundance patterns for “pristine” and“mantle-hybridized” slab melts, produced experimentally, and selected adakite arc lavas.

It can be readily demonstrated that dehydration reactions that involve low-degree partial melting of the slab are far more efficient at mobilizing and fractionating trace elements than dehydration
reactions taking place under sub-solidus conditions (e.g., Kogiso et al., 1997 and this study). We
have conducted a series of dehydration melting experiments on a natural, altered MORB composition,attempting to track along the solidus for dehydration melting at 3-11 GPa, using the ion microprobe to characterize the geochemical characteristics of low-degree slab melts and their crystalline
residues as a function of pressure. Dehydration melting at pressures less than 4-5 GPa are controlledDehydration of subducting slabs and water storage in the mantle

123

by the breakdown of amphibole and zoisite, and produce Na-rich granitoid liquids in equilibrium witheclogite residues containing rutile; low-degree partial melts closely resemble “adakite” arc lavas in“hot” subduction zones (Fig. 1).
The characteristic depletions in such elements as Ti, Nb and Ta in the experimental “slab melts”
at < 4 GPa are attributable to the the low solubility of rutile at temperatures less than 1100◦C. Hybridization of “pristine” slab melts by reaction with mantle peridotite in the overlying wedge produces
Mg-rich adakites with trace element abundance patterns parallel to those of the (low-Mg#) adakites.
These lavas form a compositional continuum with high-Mg andesites (HMA), which we interpret either as superheated hybridized adakites, or the product of partial melting of adakite-metasomatized
sub-arc mantle.
At pressures between 5 and 11 GPa, dehydration melting reactions are controlled by the breakdown of lawsonite and phengite (Schmidt, 1996; Ono, 1998); low-degree melting of the slab inthis pressure range produces K-rich granitoids, due to the preferential partitioning of Na in residual clinopyroxene with increasing pressure. Rutile is not part of the residual assemblage, and as aresult Nb and Ta abundances in low-degree slab melts increase sharply above 5 GPa. Liquids at 5-11
GPa are strongly enriched in incompatible trace elements such as Rb, Ba, U, Th, Sr, and the LREEs,
and strongly depleted in HREEs and Y; as a consequence the eclogitic residues of melting are nearlycompletely stripped of LILEs (including U and Th) by the time they reach depths of 300 km or
so. It is these chemically-depleted, ultrarefractory residues that would be incorporated into the OIB
source region. At pressures of 14-16 GPa, tenuous supercritical fluids coexist with majorite garnetand clinopyroxene assemblages (i.e., not garnetite); with trace element partitioning behavior betweenhydrous fluid and these crystalline residues remain wholly unconstrained.
References
Defant, M.J., Richerson, P.M., DeBoer, J.Z., Stewart, R.H., Maury, R.C., Bellon, H., Drummond,
M.S., Feigenson, M.D. and T.E. Jackson (1991) Dacite genesis via both slab melting and differentiation: petrogenesis of the La Yeguada volcanic complex, Panama. J. Petrology 32, 1101-1142.
Kay, S.M., Ramos, V., and M. Marquez (1993) Evidence in Cerro Pampa volcanic rocks for slab
melting prior to ridge-trench collision in southern South America. Jour. Geol. 101, 703-714.
Kogiso, T., Tatsumi, Y., and S. Nakano (1997) Trace element transport during dehydration processes
in the subducted oceanic crust: 1. Experiments and implications for the origin of ocean island
basalts. Earth Planet. Sci. Lett. 148, 193-205.
Morris, P.A. (1995) Slab melting as an explanation of Quaternary volcanism and aseismicity in southwest Japan.Geology 23, 395-398.
Ono, S. (1998) Stability limits of hydrous minerals in sediment and mid-ocean ridge basalt compositions: implications for water transport in subduction zones. J. Geophys. Res. 103, 18253-18267.
Schmidt, M.W. (1996) Experimental constraints on recycling of potassium from subducted oceanic
crust. Science 272, 1927-1930.

Session 4

124
4.11

The fate of subducted carbonate within oceanic crusts
Miki Shirasaka and Eiichi Takahashi
Tokyo Institute of Technology

The Carbonate-bearing eclogite and marble in the UHP metamorphic belt suggest that the sedimentary carbonate goes down to at least 100 km (c.f. Schertl and Okay, 1994) and it probably reaches
to about 200 km depth (Shirasaka et al., 2001). If the subducted carbonate melts, the CO2 -rich melt
occurs and is probably affect the magmatism and metasomatism in arc system. On the other hand,the solid carbonate is a probable carrier of CO2 into the deep mantle. In order to investigate the fateof carbonate within slabs, we carried out a series of experiments in MORB-CaCO3 system at 2.7-8.0
GPa.
The phase relation of MORB-CaCO3 system is shown in Figure 1. The solidus lies between
1200◦C and 1300◦C at 2.7-3.0 GPa, falls to 900-1000◦C at 4-5 GPa, and increases in temperature
with increasing pressure at 5-8 GPa. As the reaction Clinopyroxene + Aragonite = Garnet + Dolomite
shifts to the right with increasing pressure, the solidus markedly decreases at 3.5-4.0 GPa.
The carbonatitic melt coexists with clinopyroxene + garnet + coesite + rutile at >3.5 GPa. The
melt changes its composition with increasing temperature (increase in Ca, Si contents and decreases in
Mg content). In suprasolidus above 4.0 GPa, the melt composition is Ca-enriched dolomite, Ca/(Ca+Mg+Fe)= 0.6-0.7, and less SiO2 (< 1 wt.%), Al2O3 (< 1 wt.%) and Na2 O (< 0.1 wt.%) contents. Although
the melt composition at far solidus is undefined, it seems to have higher SiO2, K2 O and Na2 O contents
than those at lower temperature.
Carbonate phase in subsolidus changes with increasing pressure. At 2.7-3.5 GPa and <1100◦C,
aragonite or calcite is stable. Above ∼3.5 GPa, dolomite is stable at higher temperatures, and it decomposes to aragonite + magnesite with increasing pressure and/or with decreasing temperature.
Moreover a portion of aragonite decreases with increasing pressure. As the exchange reaction

Mg3 Al2Si3 O12 + 3 CaCO3 = Ca3 Al2Si3O12 + 3 MgCO3
shifts to the right with increasing pressure, the carbonate composition changes.
The phase relation in the present study indicates the following geological implications. (1) If
the carbonate sedimentary rocks within a relatively ’Hot’ oceanic plate subducts, the Ca-enricheddolomitic melt is produced at ∼120 km. Gorring and Kay (2000) reported that the calcite-bearing
peridotite xenolith experienced carbonatitic metasomatism in Southern Patagonia. This xenolith camefrom back-arc basin related with the subducting hot oceanic plate. We considered that this carbonatitic
melt was occurred in a slab, and then changed its composition to Ca-rich to react with wall-rock(peridotite) during ascending (Dalton and Wood, 1993). (2) If the carbonate sedimentary rocks withina relatively ’Cold’ oceanic plate subducts, the carbonate changes in order calcite, dolomite, aragonite+ magnesite with increasing depth. Aragonite within eclogite seems to decrease or disappear becauseof the reaction with garnet, and finally magnesite is stable in the upper mantle.
References
Bohlen, S. R. and Boettcher, A. L. (1982) The quartz-coesite transformation: a pressure determinationand the effects of other components. J. Geophys. Res., 87, 7073-7078.
Dalton and Wood (1993) The composition of primary carbonate melt and their evolution throughwallrock reaction in the mantle. Earth Planet. Sci. Let., 119, 511-525.
Gorring, M. L. and Kay, S. M. (2000) Carbonatite metasomatized peridotite xenoliths from southern Patagonia: implications for lithospheric processes and Neogene plateau magmatism. Contrib.
Mineral. Petrol., 140, 55-72.

Dehydration of subducting slabs and water storage in the mantle

125

Peacock, S. M., Rushmer, T. and Thompson, A. B. (1994) Partial melting of subducting oceanic crust.
Earth Planet. Sci. Let., 121, 227-244.
Schertl, H. P. and Okay, A. I. (1994) A coesite inclusion in dolomite in Dabie Shan, China: petrological and rheological significance. Eur. J. Mineral., 6, 995-1000.
Shirasaka, M., Takahashi, E., Nishihara, Y., Matsukage, K. and Kikegawa, T. (2001) In situ X-ray
observation of the reaction dolomite = aragonite + magnesite at 900-1300 K. American Mineralogist, submitted

Session 4

126
4.12

An Experimental Study of the Effect of Water on the 410-km Discontinuity
Joseph R. Smyth1∗ , Daniel J. Frost2

1 Department of Geological Sciences, University of Colorado, Boulder, CO 80309 USA; and2

Bayersiches Geoinstitut, Universitä Bayreuth, Bayreuth, D94550 Germany.
∗<joseph.smyth@colorado.edu>

The H content of the Earth is one of the most poorly constrained compositional variables forthe planet. The nominally anhydrous olivine and spinelloid phases thought to compose the bulk of
the upper mantle and transition zone may contain many times the amount of H and O that reside in
the hydrosphere. The seismic discontinuities at depths of 410 and 660 kilometers, which bound theTransition Zone, are generally thought to be due to phase changes in an olivine-dominated mantle
composition. The discontinuity at 410 kilometers corresponds to the olivine-wadsleyite transition
with an increase in both density and S-wave velocity of about five percent. Wood (1995) suggested
that the presence of H2O in the transition zone would extend the stability of wadsleyite to lower
pressures and broaden the two-phase region. To test this hypothesis, we have conducted a series ofsynthesis experiments in the multi-anvil press on hydrous and anhydrous peridotite compositions andcharacterized the products by electron microprobe and single-crystal X-ray diffraction.
Three preliminary experiments were conducted in a hydrous olivine system. Starting material wasformulated from powdered natural Fo90 olivine, brucite, quartz, and wtite to achieve a compositionof Fo90 plus 3 wt percent H2 O with a slight excess of silica. Three additional experiments wereconducted in a hydrous peridotite system. The starting material was formulated from powdered natural
Fo90 olivine, plus synthetic brucite, quartz, pyrope and wtite to achieve a composition of Fo90 plus
3 wt percent H2 O with a sufficient excess of silica and alumina to give approximately 10 normative
percent each of pyroxene and pyrope-almandine garnet. The results of our synthesis experiments
are consistent with the prediction of Wood (1995) that the presence of H2 O extends the stability ofwadsleyite to 0.6 to 1.0 GPa lower pressure and would necessarily broaden the two-phase loop to 18
or as much as 30 km.
We have characterized the olivine and wadsleyite products by single-crystal X-ray diffraction.
Kohlstedt et al. (1996) reported that olivine can accept 2000 ppm by weight H2 O at pressures inexcess of 10GPa. We have refined the atom position parameters and unit cells of five hydrous andanhydrous olivines. We observe a significant expansion of the unit cell volume and the M2 polyhedral
volume in the hydrous samples indicating that H incorporation occurs in association with M2 cation
vacancies. The volume expansion is large enough to have a significant effect on seismic velocities.
Trivalent cations and H are strongly partitioned to wadsleyite. We have refined the atom position
parameters and unit cells of six hydrous and anhydrous wadsleyites. Those that contain the largest
proportions of trivalent cations show the greatest deviations from orthorhombic symmetry. Althougholivine cannot accept nearly as much H as wadsleyite, it has a larger volume of hydration than wadsleyite.
In all of our hydrous runs containing both olivine and wadsleyite, there appears a sharp boundary
between the regions of olivine and wadsleyite. The texture of the run thus does not appear to bea simple chemical equilibrium, but rather a diffusion-controlled boundary. Hydrogen is known todiffuse very rapidly in these materials (Mackwell and Kohlstedt, 1990). This raises the possibility
that diffusion of H might control the texture and may affect the sharpness of the boundary in thenatural system.
Hydrous wadsleyite is still about five percent more dense than anhydrous olivine. In a hypothetical
two-phase region consisting of olivine and wadsleyite plus lesser amounts of garnet and clinopyroxene
extending over a depth 20 km in a hydrous system, gravitational equilibrium may be approached by
rapid diffusion of H. This would enrich the lower parts of the two-phase region in wadsleyite andthe upper part in olivine, thus sharpening the boundary. This mechanism could easily sharpen theDehydration of subducting slabs and water storage in the mantle

127

apparent boundary to four kilometers or less as indicated in some seismic studies (Benz and Vidale
1993; Yamazaki and Hirahara, 1994; Vidale et al., 1995). Because diffusion of H is considerably more
rapid than convection it might cause the transition zone to serve as a trap for H in the Earth.
In summary, our experimental results are consistent with the prediction of Wood (1995) that thepresence of H would move the olivine-wadsleyite transition to 0.6 to 1.0 GPa lower pressure andwould expand the two phase region to 20 km or more. Although olivine can accept only up to about
2000 ppm by weight H2 O whereas wadsleyite can accept 30,000 ppm, olivine has the larger volume
of hydration. Both hydrous olivine and hydrous wadsleyite are expected to be significantly slower
than their anhydrous equivalents, but in both the hydrous and anhydrous systems wadsleyite is still
about five percent more dense than olivine. Because H diffusion is rapid relative to convection it ispossible that H diffusion at the boundary may reduce the two-phase transition interval to approachgravitational equilibrium. So the 410-km discontinuity in a hydrous system is likely to appear sharper,
shallower, and slower than in an anhydrous system.
References
Benz, H. M. and Vidale, J. E. (1993) Nature 365, 147.
Kohlstedt, D.L.,et al.(1996) Contrib. Mineral. Petrol., 123, 345.
Mackwell, S. J. and Kohlstedt, D.L. (199) J. Geophys. Res. 95, 5079.
Neele, F., (1996) Geophys. Res. Lett. 23, 419-422.
Vidale, J.E., et al. (1995) Geophys. Res. Lett. 22, 2557.
Wood, B.J. (1995) Science 268, 74.
Yamazaki, A. and Hirahara, K. (1994) Geophys. Res. Lett. 21, 1811.

Session 4

128
4.13

The effect of water on the high-pressure phase boundaries in the system
Mg2 SiO4-Fe2 SiO4

Toru Inoue§∗ , Yuji Higo†, Takayuki Ueda†, Yasutika Tanimoto† and Tetsuo Irifune§§ Geodynamics Research Center, Ehime Univ., Matsuyama 790-8577, Japan,†

Dept. Earth Sciences, Ehime Univ., Matsuyama 790-8577, Japan∗<inoue@sci.ehime-u.ac.jp>

1. Introduction
H2 O is an important volatile material in the Earth, which may be transported to the Earth’s deep
mantle by subducting slab [e.g. Irifune et al., 1998; Ono, 1998; Okamoto and Maruyama, 1999].
Recent experiments clarified that wadsleyite(β) and ringwoodite(γ), which are high pressure polymorphs of olivine(α) and the major high pressure phases constituting the transition zone (410-660km), can accommodate significant amounts of H2O (∼2-3 wt%) in their crystal structures [Inoue etal., 1995; Kohlstedt et al., 1996; Inoue et al., 1998; Yusa et al., 2000; Ohtani et al., 2000]. Such H2 O
should affect the nature of the phase transformation of olivine polymorphs, as it changes the densities
and the thermodynamic parameters of the anhydrous forms.
We have conducted high pressure experiments to clarify the effect of H2 O on the α-β-γ andpostspinel phase boundaries. On the basis of these experimental data, we discuss the nature of theseismic discontinuities in the mantle transition zone.
2. Experimental procedure
High pressure experiments were conducted using an MA8-type high pressure apparatus in Ehime
University. The pressure calibration was done on the basis of the anhydrous experiments of α-β-γphase transitions [Katsura and Ito. 1989; Akaogi et al., 1989], and the postspinel transition [Irifune etal.,1998]. Temperature was measured with a W97Re3-W75Re25 thermocouple, whose hot junction
was placed in the center of the furnace assembly. The fluctuation of temperature throughout the run
was kept within ±5 ◦C in each condition, but no pressure correction was made for the emf of thethermocouple.
Three starting materials (Mg2 SiO4, (Mg0.9 Fe0.1 )2 SiO4 and (Mg0.8 Fe0.2 )2SiO4) with 1 wt% H2 O
were prepared in theα-β-γ phase transitions; these were made as mechanical mixtures of the oxides(SiO2 and MgO), hydroxide (Mg(OH)2) and Fe2SiO4 in adequate proportions. For anhydrous starting
material, we used the mixture of forsterite and fayalite. In the postspinel transition, three hydrous
starting materials (Mg2 SiO4 + 1, 2 and 3 wt% H2 O) were prepared.
In most experiments, two capsules containing hydrous and anhydrous starting materials wereplaced in one cell assembly to directly compare the difference between the two materials under thesame pressure and temperature condition. In some experiments, two starting materials with different
H2 O contents were also used for comparison.
In high-pressure and high-temperature experiments, pressure was applied first to the target ram
loads and then the temperature was kept for mainly 30 minutes, and then the run was quenched byturning off the electric power.
The polished section of the recovered run product was prepared to identify the phases present andto measure their chemical compositions. A wave-dispersive electron probe microanalyser (EPMA)
and a microfocused X-ray diffractometer were used for these purposes.
3. Results and discusssion

Dehydration of subducting slabs and water storage in the mantle

129

In the α-β-γ phase transitions, the experiments were conducted in a pressure range from 11 to
21 GPa at a constant temperature of 1400 ◦C in the system Mg2 SiO4-Fe2SiO4 with 1 wt% H2 O. The
phase boundary between α and β in hydrous condition shifted to low pressure side comparing withthe anhydrous boundary; this result is consistent with that reported by Chen et al. (in preparation).
While in the β-γ transition, the boundary shifted to high pressure side comparing with the anhydrous
boundary.
In the postspinel transition, the experiments were conducted in a pressure range from 21 to 22
GPa at a constant temperature of 1600 ◦ C in the system Mg2 SiO4-H2O 1, 2 and 3 wt%. The boundaryof postspinel transformation shifted to high pressure side comparing with the anhydrous boundary.
The interpretation of the above results can be made by the partition behavior of H2 O between
α-β-γ and perovskite(Pv)+periclase(Pc), i.e. (H2 O in α) < (H2 O in β) > (H2O in γ) > (H2O in Pv
and Pc). Details of the present results are discussed in our presentation.
References
Akaogi, M., E. Ito, and A. Navrotsky, J. Geophys. Res., 94, 15671-15685, 1989.
Chen, J., T. Inoue, H. Yurimoto, and D.J. Weidner, submitted to Science.
Inoue, T., H. Yurimoto, and Y. Kudoh, Geophys. Res. Lett., 22, 117-120, 1995.
Inoue, T., D.J. Weidner, P.A. Northrup, and J.B. Parise, Earth Planet. Sci. Lett., 160, 107-113, 1998.
Irifune, T, N. Kubo, M. Isshiki, and Y. Yamazaki, Geophys. Res. Lett., 25, 203-206, 1998.
Irifune, T., N. Nishiyama, K. Kuroda, T. Inoue, M. Isshiki, W. Utsumi, K. Funakoshi, S. Urakawa, T.
Uchida, T. Katsura, and O. Ohtaka, Science, 279, 1698-1700, 1998.
Katsura, T. and E. Ito, J. Geophys. Res., 94, 15663-15670, 1989.
Kohlstedt, D.L., H. Keppler, and D.C. Rubie, Contib. Mineral. Petrol., 123, 345-357, 1996.
Ohtani, E, H. Mizobata, and H. Yurimoto, Phys. Chem. Minerals, 27, 533-544, 2000.
Okamoto, K., and S. Maruyama, Am. Mineral., 84, 362-373, 1999.
Ono, S., J. Geophys. Res., 103, 18253-18267, 1998.
Yusa, H., T. Inoue, and Y. Ohishi, Geophys. Res. Lett., 27, 413-416, 2000.

Session 4

130
4.14

Is the transition zone a water reservoir ?
G.C.M Richard∗ , M. Monnereau

Laboratoire de Dynamique Terrestre et Planétaire, UMR 5562, Observatoire Midi-Pyrénées, 14 Avenue
Edouard Belin, 31400 Toulouse, France∗
email: Guillaume.Richard@cnes.fr

Water is present everywhere in the Earth’s mantle today, albeit sometimes at the parts per million(p.p.m) level. Earth’s mantle total water content is estimated to be close to actual hydrosphere’s watercontent. Numerous studies have been devoted to determination of water solubility in mantle rocks [1,2, 3]. They all show strong solubility variation from one mineral phase to another. Principally, waterpartitioning have made the transition zone a probable water-trap of the Earth’s mantle. Nevertheless,water distribution within the mantle is still debated to date.
Here we have investigated the effects of mantle dynamics on water distribution in a 2D cartesianbox. The model takes into account water partitioning between mantle transition zone and upper andlower mantle (i.e between resp. Olivine-Spinel and Spinel-PostSpinel ). That leads, in absence of
advection to an equilibrium state with 100 ppm, 1000 ppm, 10 ppm respectively in the upper mantle,
transition zone and lower mantle. We have modelled mantle temperature field using depth-dependent
viscosity and plate like surface conditions. Water injection at trench has also been simulated.
Numerical experiments suggest that diffusivity of water has to be very high, at least two orders ofmagnitude higher than the one experimentally determined [1] to influence significantly water distribution in Earth’s mantle (Fig. 1). Infact, the diffusion process is not efficient enougth to balance themixing due to mantle dynamics and to keep water into the transition zone. We find that the distributionof water should be quite homogeneous in the whole mantle. Therefore, water amount deeper than thetransition zone would be in excess according to the lower mantle rocks solubility. This adresses thestability of free water in the lower mantle and its mobility by percolation process.

References[1] J. Ingrin, H. Skogby, Hydrogen in nominally anhydrous upper mantle minerals: Concentrationlevels and implications, Eur. J. Mineral. 12 (2000) 543-570.
[2] D. L. Kohlstedt, H. Keppler, D. C. Rubie, Solubility of water in the alpha, beta and gamma phasesof (Mg,Fe)2SiO4, Contributions to mineralogy and petrology 123(4) (1996) 345-357.
[3] N. Bolfan-Casanova, H. Keppler, D. C. Rubie, Water partitioning between nominally anhydrous
mminerals in the MgO-SiO2-H2 O system up to 24 GPa: implications for the distribution of waterin the Earth’s mantle, Earth Planet. Sci. Lett. 182 (2000) 209-221.

Dehydration of subducting slabs and water storage in the mantle-6

K = 10131

K = 10

50

70

90

110

-8

130

H2O

Stream function

Temperature

Fig. 1 Effect of water diffusivity on Earth’s mantle water distribution. Stream (central box) created by bottom
heating (bottom box) is applied to an homogeneous water content mantle. Snapshots of the water content in
ppm after convergence is supplied for different water diffusivities (top boxes). Concentration variation in thehigh diffusivity case K = 10−6 m2 ·s−1 (left hand side) reach a factor two, but stay two low to be discernable inthis range in the low diffusivity case K = 10−8 m2 ·s−1 (right hand side)Session 4

132
4.15

Solubility of water in lower mantle minerals

Motohiko Murakami∗, Kei Hirose, Hisayoshi Yurimoto, Satoru Nakashima∗

Tokyo Institute of Technology

The maximum solubilities of water in the representative lower mantle minerals, MgSiO3-rich perovskite, magnesiowustite, and CaSiO3 -rich perovskite, were determined using secondary ion mass
spectrometry. They were synthesized in the peridotitic compositions by the multianvil apparatus at
23.5 GPa under water-saturated conditions. We found that Mg-perovskite and magnesiowustite dissolved about 0.2 wt% H2 O and Ca-perovskite dissolved about 0.4 wt% H2O, respectively. The OH
absorption bands in Mg-perovskite and magnesiowustite were also confirmed using microscopic FTIR, and the water content estimated from these OH absorption bands is consistent with that fromSIMS analysis. These results suggest that lower mantle in itself can storage considerable amounts ofwater, and thus could provide essential implications for the physical properties of the Earth’s lower
mantle.

Dehydration of subducting slabs and water storage in the mantle
4.16

133

Electrical conductivity of hydrous materials at high pressures and high temperatures.
Tomoo Katsura1 , Kiyoshi Fujita2, and Keishi Shinoda3
1: Institute for Study of the Earth’s Interior, Okayama University
2: Faculty of Sciences, Kobe University
3: Faculty of Sciences, Osaka City University

Hydrogen plays important roles in electrical conduction in the Earth. For example, electrical conduction in the upper crust is essentially controlled by fluid phases interstitially present in rocks. It is
possible that hydrogen is important for electrical conduction also in the mantle and lower crust. From
this reason, we are measuring electrical conductivity of hydrous materials at high pressures. Here, wereport results of measurement for brucite and granulite.
Electrical conductivity was measured using sinusoidal signal with 10 mHz frequency and 10 Vp−pamplitude. Reference resistance from 104 to 106 Ω was connected to the sample in a series. The sample impedance was calculated from reference resistance and voltages on the reference resistance andwhole circuit. Effect of capacitance was reduced by assuming that capacitance is connected with sample resistance in parallel. High pressures and temperatures are generated using 6-8 type and DIA-typeapparatuses. To avoid decomposition of hydrous minerals or escape of a fluid phase, measurement has
to be conducted at relatively low temperatures, typically up to 1000 K, in which minerals and rocks
usually have very low electrical conductivity. To obtain high insulation resistance, high-pressure cells
were composed of boron nitride, magnesia, thermocouples, and metallic heater system, resulting in
ca. 10 GΩ insulation resistance up to 1000 K.
Brucite contains Mg, O, and H, but not transition metals. Therefore, hydrogen ion (proton) should
be a carrier of electrical conduction in brucite. Pressure and temperature conditions of measurement
were 3, 6, 9 and 12 GPa and from 300 to 700 K, respectively. In each pressure step, measurement was
carried out in the heating and cooling cycles. The activation energy is decreased with elevated pressure: from 0.99 eV at 3 GPa to 0.75 eV at 12 GPa. This is probably because distance from hydrogento the neighboring site becomes shorter with compression, On the contrary, the pre-exponential term
is nearly constant of 5 × 103 S/m, indicating no change of population of the electrical carrier (proton)
Another measurement was conducted for garnet-cordielite-biotite-granulite from Hidaka metamorphic belt, Hokkaido, Japan. It is considered that this granulite was formed at 1000-1100 K and
1 GPa. It was ground into powder, loaded in a Pt capsule, and sintered at 1025 K and 1 GPa for 1
hour. The sintered sample with 2.5 mm diameter and 1.1 mm thickness was set in a pressure cell forelectrical conductivity measurement and compressed to 1 GPa. First we heated to 500 K to purge
water in a pressure medium, and decreased temperature to ca. 300 K. After that, we conducted measurements in two heating cycles: heating from 300 K to 900 K and cooling to 300K, and heating from
300K to 1000 K and cooling to 1000 K. The conductivity was decreased by one order of magnitude
at low temperatures by the heating from 300 to 900 K, which is probably due to annealing of thesample. The conductivities measured at in the cooling from 900 K to 300 K and the heating from 300
to 1000 K are identical, indicating that conductivity values measured in these paths are essential to
the granulite. The activation energy and pre-exponential term of this intrinsic conductivity are 0.39
eV and 8 × 10−3 S/m, respectively. The conductivity varies by more than 3 orders of magnitude in
the first cooling, first heating, and second heating paths in the present temperature range. However,after reaching 1000 K, conductivity does not varies with cooling, is almost constant of 1 × 10−4 S/m.
This high and constant conductivity could be attributed to a fluid phase formed by dehydration.

October 5 (Friday) 09:00 — 12:00

Session 5

Magmatic differentiation and the evolution of the
mantle and core

Keynote: Eiichi Takahashi, Tokyo Institute of Technology
Chair: Bernie Wood, University of Bristol

135

Magmatic differentiation and the evolution of the mantle and core
5.1

137

Melt Generation in the Hawaiian Plume

E. Takahashi1 , K. Nakajima1 , R. Takeguchi1 , K. Shinozaki1, Ren Zhang-Young1, R. Tanaka2 and E.
Nakamura2
1: Magma Factory, Earth & Planet Sci., Tokyo Inst. of Tech., 2-12-1 Ookayama Meguroku, Tokyo 152, Japan
2: ISEI, Okayama Univ., Misasa Japan

Genesis of magmas in large hot spot is very important to understand the temperature, chemicalcomposition of the hot mantle plume. Compared with the magmatism under the mid oceanic ridges
and those in the subduction zones, magma genesis related to mantle plumes is poorly understood. Inorder to study the origin and dynamics of mantle plumes, we have carried out both filed works andexperimental work on several volcanoes in Hawaii which is the most magma roductive hot spot on
our planet.
1. Reconstruction of Koolau volcano
Among Hawaiian shield volcanoes, Koolau shows very silica-enriched trend (SiO2=53-55 wt% at
MgO=7-5 wt%, see Fig.1). The Koolau tholeiites are also peculiar in that they show enriched isotope
signatures in Sr, Nd and Pb (Roden & Frey, 1994). Based on co-vraition in major elements and isotopes among Hawaiian shield volcanoes, Hauri (1996) proposed that the high-silica Koolau tholeiite
may contain ca. 10% of dacitic melt derived from former oceanic crusts entrained in the plume. In order to constrain the primary magma composition for Koolau volcano, Takeguchi & Takahashi (2001)
studied 250 rock specimen collected from subaerial part of Koolau which covers wider stratigraphic
units than previous works (Frey et al.1994; Jackson et al., 1999). In addition, Takahashi & Shinozaki(2001) studied submarine portion of the volcano (ca. 200 rock samples recovered by JAMSTEC submersibles) where much of its deep interior is exposed by the giant Nuuanu landslide (Moore et al.,
1989, AGU Monograph on Hawaii, 2001) A sharp transition in lava chemistry and isotope ratios wasobserved at a low stratigraphic horizon of Koolau volcano (first noted by Jackson et al., 1999) where
high-silica Koolau type lava changes to mildly silica-enriched normal Hawaiian tholeiite (similar to
Mauna Loa and Kilauea). This transition in lava chemistry accompanies with large changes in Sr, Nd
and Pb isotope ratios from EM-1 like Koolau values to those of modern Mauna Loa and Kilauea.
We have shown for the first time that a Hawaiian shield volcano recorded shift in tholeiite magmacomposition from one to the other and the silica-rich Koolau tholeiite represents only the final growthstage (Makapuu stage) of the volcano (Shinozaki et al., 2001; Tanaka et al., 2001).
2. Primary Koolau magma composition
Picrites are occasionally found in Hawaiian shield volcanoes as well as in Koolau. Implicationof the picrites, however, depends on the researchers. Some argues that olivine-rich picrites represent primary magma composition of the Hawaiian hot spot. Presence of kink banding in may olivine
crystals in picrite (e.g., Heltz, 1987; Garcia, 1996), however, casts some doubt for the simple highpressure phenocryst hypothesis. Kink banding is common in mantle peridotite with subsolidus deformation. Many of olivine crystals in mantle peridotite, however, dose not show kink banding. Presenceof olivine xenocrysts with kink banding implies that there may be more olivine xenocrysts in the
Hawaiian picrite which dose not show kink banding.
In order to discriminate olivine xenocrysts, we carried out dislocation mapping analysis for four
picrite specimens from Koolau volcano (Takeguchi & Takahashi, 2001). Polished rock chips were
heated at 800◦C for 1 hr in air for oxidation, and then the dislocation density was calculated by point
counting technique using petrologic microscope. At least three areas were counted within an olivine
crystal and the results were averaged. In some picrite, olivine crystals were divided into two distinct
types by the dislocation density; type-A, with dislocation desity 106 -109 /cm2; type C, 100 -103 /cm2.
The type-A crystals are enriched in MgO and NiO than type C olivine. Judging from the bulk chemical

138

Session 5composition of the picrite, the type C crystals can be in equilibrium with the host magma but the
type-A crystals are difficult. In some picrite, olivine crystal with marginal dislocation density occurs;
type-B ( 104-106 /cm2 ). Type-B crystals may also be xenocrystic origin but have been annealed in
the host magma for longer time. Based on our dislocation mapping it was found that olivine crystals
with Fo>85% are xenocrystic origin in Koolau volcano. Conduit zone of the gigantic Hawaiian shield
volcanoes penetrate more than 70km thick Pacific plate. These highly strained olivine xenocrysts may
originate from the conduit zone where high temperature deformation of the rocks take place due tosuccessive intrusion of hot basaltic magma.
After carefully eliminating the effect of olivine xenocrysts, we concluded that primitive magmaof Koolau volcano has a basaltic andesite composition (SiO2=53 wt%, MgO=7 wt%). Picrites arepresent in Koolau but they are mixtures of this primitive melt and olivine xenocrysts (Takeguchi &Takahashi, 2001).
3. Experiments to understand the melting process in the Hawaiian Plume
In order to understand the condition of melting and the source rocks for the various primary
Hawaiian magmas, we carried out melting experiments with MORB/peridotite and Archean-basalt/peridotite
sandwich starting materials (Takahashi & Nakajima, 2001). It was found that Koolau type primitive
melts can be produced by extensive partial melting of basaltic source rock alone at 2.5-3.0 GPa and1350-1400◦C; conditions slightly below the dry solidus of mantle peridotite. At conditions on or
above the peridotite solidus, melts become saturated with olivine and are similar to picritic magma in
Kilauea. Accordingly, we propose that the high-silica Koolau primary magma are produced by almost
entirely from the recycled oceanic crust (cf. ca.10% by Hauri, 1996). Changing isotope compositions
found in Koolau volcano may indicate variety of mixing ratio between eclogite derived melt and theambient peridotite derived melt.
In the basalt/peridotite sandwich melting experiments, melt proportion within the basalt domain
decreases as temperature increase near the peridotite solidus. This is due to the enhanced chemical
reaction between the basalt domain and the ambient peridotite. Due to this effect, bulk melt proportion in the sandwich starting material records minimum at a temperature near the peridotite solidus.
Accordingly, melts produced in the peridotite at temperatures above the peridotite solidus are moreenriched in incompatible elements K, Ti and LREE than those produced at lower temperatures in thebasalt domain. This might explains the difference in trace element characters of Koolau and Kilauea.
Based on the melting experiments and the field observation, we propose a model that the Hawaiian plume had a potential mantle temperature (PMT) of only 1350-1400◦C in the Koolau stage (ca2.5 Ma). This PMT is much lower than the estimate for the modern Hawaiian plume by Watson &McKenzie (1991, PMT=1558◦C) assuming homogeneous peridotite source. Our experiments showthat only slight increase in PMT (10-50deg) will shift the Koolau type melts to the parental Mauna
Loa and Kilauea type melts (SiO2=49, MgO=14) in the plume head.
4. Phase relation of eclogite and its effect on bulk density
The important role of recycled old oceanic crust in magma genesis in large hot spots (particularlythat in Hawaii has become clear in our study. The mismatch in size of the tomographic image ofthe mantle plume and their magma production rate has been discussed by Zhao et al (2001) andthey proposed that it could be explained by the different amount of eclogite entrainments inside theplume. Distribution of the once subducted eclogite blocks in the deep mantle and their entrainment
mechanism, therefore, is essentially important to understand the mantle plume related magmatism.
According to Irifune & Ringwood (1987, 1993) eclogite of MORB composition consists of garnetand omphasite at pressures below 15 GPa and becomes almost monomineralic majorite rock at higher
pressures. This is reconfirmed in our experiments. The proportion of pyroxene/garnet, however, wasfound to change significantly with temperature at given pressure. The amount of garnet in Irifune &Ringwood’s experiments were found to be underestimated (as much as 20-50%) due to metastable

Magmatic differentiation and the evolution of the mantle and core139

growth of aluminous pyroxenes. The amount of coesite and stishovite in eclogite was underestimatedbecause of the same reason. Increase in the amount of garnet and stishovite increase the bulk densityof eclogite dramatically.
Due to the combined effect of higher amount of garnet and stishovite, calculated zero-pressurebulk density of MORB eclogite equilibrated at 10 GPa and 1200◦C increased from 3.6 g/cc (Irifune &Ringwood, 1987) to 3.75 g/cc (Nishihara, Aoki and Takahashi, in preparation). The density difference
between the eclogite and fertile peridotite (KLB-1) is 0.1-0.3 g/cc at temperatures of 1200-1800◦C in
the upper mantle (<400km depth) whereas it decreases to less than 0.1 g/cc in the mantle transitionzone (400-700km depth) at temperatures above 1800◦C.
Equation of state of majoritic garnet synthesized from natural abyssal basalt was determined up
to 16GPa and 1000◦C by in situ X-ray diffraction using a single-stage, cubic multi-anvil apparatus
MAX-90 and a double stage SPEED apparatus. The MORB majorite was synthesized from a starting
material at 19 GPa and 1900◦C using 6-8 type multi-anvil apparatus. The bulk modulus for the MORB
majorite was obtained by the third-order Birch-Murnaghan equation of state. The obtained value(160.9 GPa) is much smaller than MORB-majorite (226.2GPa) by Faust & Knittle (1995) but isconsistent with majorite in Mg-Al-Si system by Yagi et al (1987) (Nishihara et al., 2001)
E.Takahashi and K.Nakajima (2001) Melting Process in the Hawaiian plume: An experimental study
In Evolution of Hawaiian Volcanoes: recent progress in deep underwater research, Amer. Geophys. Union, Monograph, (in press)
K. Shinozaki, Zhong-Yuan Ren and E. Takahashi (2001) Geochemical and Petrological Characteristics of Nuuanu and Wailau landslide blocks, In Evolution of Hawaiian Volcanoes: recent progressin deep underwater research, Amer. Geophys. Union, Monograph, (in press).
R.Tanaka, E.Nakamura and E.Takahashi (2001) Geochemical Characteristics of Koolau volcano andNuuanu and Wailau landslide deposits, Hawaii, In Evolution of Hawaiian Volcanoes: recent
progress in deep underwater research, Amer. Geophys. Union, Monograph, (in press).
D.Zhao, F.Ochi and E.Takahashi (2001) Seismic images of mantle plumes under hotspot regions, In
Evolution of Hawaiian Volcanoes: recent progress in deep underwater research, Amer. Geophys.
Union, Monograph, (in press).
R.Takeguchi and E.Takahashi (2001) Primary magma composition a the final growth stage of Koolau
volcano, Hawaii. Submitted to Geochimica Cosmochimica Acta.

Session 5

140
5.2

Mantle composition and conditions of core formation
B.J. Wood
University of Bristol

A major, defining event in the history of the earth was the separation, within about 80 Ma of accretion of its iron-rich core. The profound chemical fractionation associated with this event segregatedinto the core most of the earth’s Fe and ’siderophile’ elements (Ni, Co, W, Re etc) together with about
10% of one or more ’light’ elements. This resulted in characteristic patterns of depletion in the silicate earth of siderophile (removed to the core) and volatile (possibly incompletely accreted) elements
relative to primitive chondritic material The ’light’ element, in reality a mixture of light elements,
is probably dominated by S, with Si and C as potentially important minor components .The patternof siderophile element depletion in the silicate earth cannot be explained by low pressure equilibrium between metallic core and a silicate melt of mantle composition because siderophile element
abundances in the mantle are too high. This observation has led to two competing models of coreformation. In one model the earth accreted heterogeneously and began to melt when it was about 25%of the current size due to the energies of impact . Core segregation would begin and continue near thesurface of the accreting planet and metal-silicate fractionation patterns be established at low pressure(about 1 Gpa) .In order to explain the high contents of siderophile elements in the mantle conditions
would have had to become more oxidising as accretion progressed. Most of the mantle abundancesof the highly siderophile elements would then have been added as a ’late veneer’ of oxidised material
after core formation. The alternative model relies on the observation that the metal-silicate partition
coefficients for some siderophile elements (e.g Ni,Co, Re) decrease with increasing pressure suchthat there is a theoretical pressure-temperature condition of about 30 GPa and 2500C where the current mantle abundances of these elements could be explained by single-stage equilibrium between
core and mantle. In this model it is postulated that descending droplets of liquid metal equilibrated
with the silicate at the base of a 750 km deep magma and established most of the current pattern ofsiderophile element depletion in a single event. Distinguishing between the two models is difficult,
but we believe that weakly siderophile elements such as V, Cr, Si and Nb provide the best constraints.
The silicate earth is substantially depleted in V and Cr relative to CI chondrites, depletions frequently ascribed to dissolution in the core. In contrast, refractory 鼠 ithophileelements such as Ca,
Sc and the rare earth elements (REE) are known to be present in the same proportions in the silicate
earth as in chondritic meteorites. Nb has always been treated as lithophile and refractory, yet at 25
GPa and plausible temperatures for core segregation, Nb rather than being lithophile, exhibits almost
identical siderophile character to V and Cr. Thus V, Cr and Nb have approximately the same ratios of(core concentration)/ (silicate earth concentration). This means either that the depletions of the upper
mantle in V and Cr were not due to core segregation or that Nb is also depleted in the upper mantle.
Geochemical data suggest that a slight Nb depletion is plausible. In either case strongly reducing conditions would lead to too great a depletion of the mantle in Nb, so a large part of core formation undersuch conditions can be excluded. Si is known to dissolve readily in Fe at low pressures and has been
proposed as being a major part of the light element component of the core (e.g All 鑒 re et. al. Earth
Planet. Sci. Lett. 134, 1995). Si remains substantially more lithophile than Nb under all conditions wehave investigated (2.5-25 GPa, 1750-2300 C), however. This means that Si cannot be part of the lightelement inventory of the core unless extreme conditions of temperature and pressure were involved(40 GPa, 3600K or higher).

Magmatic differentiation and the evolution of the mantle and core
5.3

141

Boundary layer fractionation constrained by differential information from theKutsugata lava flow, Rishiri Volcano, Japan
Takeshi Kuritani
Institute for Study of the Earth’s Interior, Okayama University<kuritani@misasa.okayama-u.ac.jp>

Detailed investigation of an erupted magma with limited compositional diversity provides instantaneous information on incremental magma differentiation processes in a magma chamber. Kutsugata
lava, a suitable target of such study, is a Quaternary alkali basalt (51.3-53.2 wt% in SiO2) from Rishiri
Volcano, northern Japan. Despite the narrow range in the whole rock compositional variation, chemical and modal compositions of mineral phases crystallized in the magma chamber vary systematicallywith the whole rock composition. In the North lava (51.3-51.9 wt% in SiO2) the less differentiatedportion of the Kutsugata lava, most crystals which include low-Ni olivine and plagioclase were derived from the mushy boundary layer. The main part of the magma body was principally aphyric (<0.5 vol% crystals). Estimated chemical compositions of fractionated mineral phases during differentiation of the magma coincide with the observed compositions of low-Ni olivine and plagioclase
crystals. This indicates that the main magma was differentiated by separation of crystals grown in themush zone. The low-density interstitial melt is suggested to have been extracted from the floor mushzone with average crystallinity of > 30 vol% by such mechanisms as compaction and compositional
convection. This fractionated melt was mixed with the overlying main magma, causing differentiationof the Kutsugata magma. The average temperature of the extracted melt is 1010◦C, significantly lower
than 1100◦ C estimated for the main magma. A quantitative model of magmatic differentiation, which
includes thermal and compositional evolution of a mushy boundary layer, can successfully reproduce
the observed compositional trends of the North lava.

Session 5

142
5.4

Constraints on Porosity in Partially Molten Regions in the Upper Mantle from
Permeability Measurements and U-series Modeling
Ulrich Faul
Research School of Earth Sciences, The Australian National University, Canberra

The amount of melt present (porosity) in regions of the upper mantle undergoing melting depends on the physical properties of solid matrix and melt. A critical parameter is the permeabilityof the matrix, which is determined by the geometry of the intergranular melt distribution. While
experimental observations indicate that basaltic melt in a peridotitic matrix is interconnected at all
melt fractions, this does not automatically equate to high permeability and hence low residual porosity. High permeability is predicted for self-similar models such as cylinders on a cubic grid and theisotropic equilibrium model consisting of triple junction tubules at the edges of tetrakahedral grains.
Permeability measurements on texturally equilibrated calcite and quartz (Wark and Watson, EPSL,
164, 591-605,1998).show that natural systems deviate significantly from the ideal models, with lower
than predicted permeabilities at low porosities. SEM and TEM observations of the melt distribution
in partially molten dunite indicate the existence of wetted two-grain boundaries, which become more
numerous as the melt fraction increases. Modeling of the evolution of permeability with increasing
porosity (melt fraction) suggests that the permeability will be very low until the wetted two-grainboundaries become abundant enough to interconnect, at which point the permeability increases byorders of magnitude.
The porosity in a 1D melting column can be calculated from a fluid dynamics model based onconservation of mass and Darcy’s law (e.g. Spiegelman and Elliot, EPSL, 118, 1-20, 1993). Besides
permeability, the density difference between matrix and melt and the viscosity of the melt influence
the mobility of the melt and hence porosity. Figure 1 shows the porosity distribution in an 85km
deep melting column for different permeability models and measurements. Only if the permeability
in partially molten upper mantle were equal to that of cylinders on a cubic grid (and hence the melt
geometry would have to be that of cylinders on a cubic grid) would the porosity in the melting column
be substantially below 1%, as inferred from a range of geochemical models. With Wark and Watson’s
measured permeabilities the porosity quickly exceeds 1%.

Fig 1. Porosity distribution calculated for a 1D melting column with a solid matrix upwelling velocity of 3
cm/year, density difference melt-matrix of 500kg/m3 and a melt viscosity of 1 Pas. Melt production begins at
100km depth and ends at 25 km depth, melt productivity is constant at 0.4%/km. The different permeability
models are indicated by dashed lines (cylinders on a cubic grid), dotted (isotropic equilibrium model) andsolid (measured, Wark and Watson). The porosities calculated with the measured permeabilities indicate thatmelt is present at the percent level in partially molten upper mantle.

Magmatic differentiation and the evolution of the mantle and core143

The calculated porosities allow calculation of U-decay series disequilibria with the equilibrium
porous flow model developed by Spiegelman and Elliot (1993) and Spiegelman (Geochem. Geophys.
Geosyst. 1, 1999 6C 000030, 2000). The differences between the permeability models above arelarger than possible variations in permeability expected due to variations in phase proportions indifferent parts of the upper manlte. The same permeability model should therefore be used to calculate
isotopic excesses due to melting at mid-ocean ridges and ocean islands. Observed isotopic excesses in
MORBs are often quite large, while OIBs generally have small excesses except for Th in relatively low
degree melts. To reproduce the mid-range of measured isotopic disequilibria at mid-ocean ridges withmoderate spreading rates, the permeability of peridotite has to be that of the cylinder model, which is
the most permeable model that can be physically realized. Maximum porosities of the order of 0.1%
necessary to satisfy high Ra excesses in depleted MORBs imply substantially higher permeability.
In contrast, when used for modeling OIBs the cylinder model predicts excesses that are too high forthe low degree melts. With the threshold model mid-range excesses for the longer lived isotopes Th
and Pa are calculated for MORBs when a small amount of volatile-rich melt with low viscosity anddensity is present below the anhydrous peridotite solidus, and melt accumulation is rapid. Similarly,
observed excesses at Hawaii can be reproduced when the different melt properties are taken into
account. Because of the near ’instantaneous’ decay of Ra at porosities above 0.1% on the time scaleof physically realistic melt transport models, Ra excesses are likely to be due to processes close to theeruption site of the magmas.

144

Session 5

5.5

Evidence for slab tear and the related magmatism beneath the northwestern El Salvador,
Central America
Ryoji Tanaka and Eizo Nakamura
PML, ISEI, Okayama Univ., Misasa, Japan

Large felsic magmatism, volcano-plutonic complexes, back-arc side alkaline volcanoes, and highsulfidation type mineralization are often distributed at regular intervals along volcanic arc systems.
In order to explain such a geographical distribution, existence of large-scale tectonic segmentations
has been proposed for some volcanic arcs. Tectonic segments and segment boundaries are defined by
the offset of along-arc volcanic alignment or seismic discontinuity. It is very critical in the genesis ofarc magmas whether the subducting slab is torn or not, because heat and material exchange between
supra and sub-slab asthenospheric mantle could occur through the torn “slab window”.
A cinder cone field in NW El Salvador has been geochemically studied in order to clarify themagma genesis related to the unique segment boundary in Central America. This area was chosen
for the following reasons: (1) hypocenters deeper than 120km depth are not observed in this area,(2) Quaternary volcanoes exist successively in the across-arc direction towards the back-arc side, and(3) there is a seismic offset of the Wadati-Benioff zone beneath the area. Volcanoes are divided into
four groups according to distance from the volcanic front toward the back-arc side; VF (volcanic
front), BVF (behind volcanic front)-A, BVF-B, and BVF-C. The trace element ratios and Sr, Nd, and
Pb isotopic ratios discriminate two across-arc geochemical trends. The two trends for each isotopic
ratio cross at BVF-A lavas when plotted against the distance from the volcanic front. The B/Nb and
the Pb and Nd isotopic ratios indicate that Trend I (from VF to BVF-A) can be attributed to thesuccessive depletion of a slab derived fluid component in the magma source toward the back-arc
side. On the other hand, Trend II (from BVF-A to BVF-C) shows the high Zr/Hf, La/Yb and thepositive correlation between Zr/Hf and Pb isotopic ratios. Lavas in Trend II have higher SiO2, Ni, Cr
contents for a given value of MgO contents than those in Trend I. Among the BVF-A lavas, the mostmagnesian and nepheline-normative lava from a cinder cone shows the lowest 87 Sr/86 Sr, Nd , Zr/Nb
and the highest Zr/Hf and Pb isotopic ratios. The primary magma of this lava is inferred to be picrite.
This primary magma composition indicates the extremely high H2 O fugacity and the existence ofrefractory peridotite.
These data suggest the existence of slab tear in this region. The slab tear would induce to release the fluid component from the subducting hydrated harzburgite (serpentinite) and to melt the
subducting oceanic crust. These effects can be a trigger of alkali basalt and magnesium-rich andesite
magmatism behind the volcanic front in this region.

Magmatic differentiation and the evolution of the mantle and core
5.6

145

Geochemical evolution of Koolau Volcano, Hawaii
Ryoji Tanaka1 , Eizo Nakamura1, Eiichi Takahashi2

1

PML, ISEI, Okayama Univ., and 2 Dept. Earth Planet. Sci., Tokyo Inst. Tech.

Koolau shield is composed of older Main shield stage and younger Makapuu stage lavas. So-called
Koolau-type lavas, as recognized geochemically enriched-endmember component in the Hawaiian
shield, are restricted to only the Makapuu stage. In both the subaerial shield and northwestern submarine flank, Main shield stage lavas are buried by Makapuu stage lavas. Only the uppermost partof Main shield stage lavas are exposed in the lowermost subaerial sequences on the cliff, and the
majority of the Main shield stage lavas were obtained from the Nuuanu landslide blocks offshore
north to the Oahu Island. Main shield stage lavas are characterized by lower Ba/Nb, La/Nb, Pb/Nd,
Sr/Nb, Nd/Nb, Zr/Nb, and 87 Sr/86 Sr and higher 143 Nd/144 Nd and 206 Pb/204Pb, which resemble Loaor Kea-type compositions, than those of Makapuu stage lavas. Sr-Nd-Pb isotope systematics suggestthat Koolau shield lavas were generated from at least three endmember components, such as Enriched
Makapuu (EMK), Depleted Makapuu (DMK), and Kea (KEA) components. Both DMK and KEA are
similarly depleted in Sr and Nd isotopic compositions (87 Sr/86 Sr < 0.7035, 143 Nd/144Nd > 0.5130),but the 206 Pb/204Pb of KEA (> 18.80) is distinctively higher than DMK (< 17.78). On the other
hand, EMK is isotopically enriched in 87 Sr/86Sr (> 0.70427) and 143 Nd/144 Nd (< 0.51265) comparedwith DMK and KEA. The 207 Pb/204Pb and 208 Pb/204Pb of EMK (> 15.45 and > 37.8) are higher
than those of DMK (< 15.37 and < 37.2, respectively). Higher SiO2, Al2O3 /CaO and incompatible
trace elements than other endmembers also characterize EMK. The proportion of these endmember
components successively changed from earlier KEA-dominated with an extremely small proportionof EMK, to later EMK-dominated. These observations reveal that the Koolau shield stage lavas spannearly the entire range of the isotopic compositions reported previously for shield stage lavas fromall Hawaiian islands, and reveal a secular change in the amount of plume component involved in thegeneration of Koolau shield magmas. This suggests that it is inadequate to model the evolution andgeochemical structure of Hawaiian mantle based only on geochemical investigation of near-surface
lavas as is typically done for oceanic island basalts.

146

Session 5

5.7

A possible discontinuous behavior of melt percolation in partially molten mantle and its
implication to compositional layering in orogenic lherzolites.
Masaaki Obata∗ , Eiichi Takazawa∗ Department of Geology and Mineralogy, Kyoto University, Kita-shirakawa, Kyoyto 606-8502 Japan

Permeability is an important parameter of partially molten rocks that controls the readiness of the
migration of partial melts along the grain boundaries. Theoretically the permeability is a function ofporosity (or melt fraction) and geometry of the pores (microstructure) of the partially molten rocks.
In the upper mantle, high-temperature conditions, textural equilibrium is believed to be maintained
or to be closely approached, with the microstructure being controlled by the balancing of the surface
energies between the melt and the solids. Several models have been proposed for the permeability
function on the basis of percolation experiments and theoretical consideration for a simple and ideallized geometries. All the previous models, however, assumes one-to-one relationships between thepermeability and the porosity for a given grain size in the matrix. This may be valid only when themelt is not in motion and when the melt pressure is equal to the solid pressure. However, in a dynamical situation where melt is in motion with respect to the solid framework, the melt pressure may
deviate from the solid pressure. This inequality of the melt and solid pressures may affect the geometry of the pore and thus the microstructure and thereby affect the permeability even if the melt fraction
is kept constant. Quantitatively, in a dynamical situation, a small unbalance in pressure that is generated between the melt and solid may be adjusted through a local recrystallization by the variation of
the average curvature between the melt solid interface, which in turn affect the melt connectivity andthus the permeability function. We point out a possibility that a discontinuous topological change may
occur at some critical melt fraction that causes a sudden jump of permeability function. Such a discontinuous change of permeability may cause an abrupt motion of melt migration. Abrupt change ofcomposition observed in some orogenic lherzolite mass may be the products of such abrupt behaviorof melt percolation in the upper mantle.
We will present a geological example for the products of such a process from a large orogenic
lherzolite mass, the Horoman peridotite, Hokkaido, Japan. In the Horoman peridotite, in a 100 metermeasured section harzburgite, lherzolite and plagioclase lherzolite occur in this sequence forming alayered structure (Takazawa et al, 2001). Important features are (1) that these rocks define a single
colinear trend in compositional space in terms of whole-rock major elements composition and (2) that
there are several compositional gaps between the rocks types. The observed compositional collinearity is best explained by the partial melting and melt extraction model (Obata and Nagahara, 1987;
Takazawa et al, 2001) but the compositional jump cannot be explained by any conventional model.
We interpret the observed compositional gaps is a manifestation of a sudden jump of the permeability
function mentioned above.

Magmatic differentiation and the evolution of the mantle and core
5.8

147

The role of water in generation of komatiite magma
Yuki Asahara, Eiji Ohtani
Tohoku University

The question whether komatiites were formed under the wet or dry conditions is fundamental forunderstanding the early thermal history of the Earth. Komatiites are high magnesian volcanic rocks
and they were thought to have been formed by partial melting of the dry mantle at high pressuresand temperatures [1]. Recently, some workers suggested a possibility that komatiites magmas weregenerated under hydrous conditions based on the evidence from natural samples [2,3].
Here we reports the results of melting experiments with various H2 O contents at high pressuresand temperatures, and we evaluate possible roles for water in the generation of komatiites magmas.
Hydrous phase relations of primitive mantle compositions in the CaO-MgO-Al2O3 -SiO2-H2 O
system with 1, 2, and 5 wt.% H2 O have been investigated at 4-8 GPa. Compositions of aluminum
depleted komatiites are consistent with the liquid formed by more than 30% melting of the primitive
mantle containing H2 O up to 5 wt.% above 8 GPa. Aluminum undepleted komatiite can be formedby more than 50 wt.% melting of dry or wet primitive peridotite with H2 O up to 2 wt.% at around
4-6 GPa. Variation in chemistry of komatiites might be explained by various depths of partial melting
in the primitive mantle under various hydrous conditions, and the variation may reveal dehydration
history of Earth’s mantle.
References[1] Herzberg, C. (1992), Depth and degree of melting komatiite., J. Geophys. Res. 97, 4521 - 4540.
[2] Parman et al. (1997), Emplacement conditions of komatiites magmas from the 3.49 Ga komati
formation, Barberton Greenstone Belt, South Africa. Earth planet. Sci. Lett. 150, 303 - 323.
[3] Shimizu et al. (2001), Cr-spinel, an excellent micro-container for retaining primitive melts - implications for hydrous plume origin for komatiites, Earth planet. Sci. Lett. 189, 177 - 188.

Session 5

148
5.9

Determination of trace element partitioning between majorite and silicate melt in theprimitive mantle composition at pressure of 15 GPa

Eizo Nakamura1, Toshihiro Suzuki2, Katsura Kobayashi1 , Akio Makishima1 and Masaki Akaogi2
1 The Pheasant Memorial Laboratory for Geochemistry & Cosmochemistry (PML), Institute for Study of the

Earth’s Interior, Okayama university at Misasa, Tottori 682-0193, Japan 2 Department of Chemistry,
Gakushuin University, Mejiro, Toshima-ku, Tokyo 171-8588, Japan

Majorite, a garnet solid solution containing pyroxene component, is considered as one of theimportant constituent minerals in the transition zone of the earth’s mantle. Thus it has been proposedthat majorite played an important role in the chemical fractionation deep in the earth’s mantle based on
the high-pressure partitioning experiments. However, most of the previous experiments were carriedout using starting materials doped by considerably large amounts of small number of key elements(wt. % level available for EPMA) for “trace elements”, and the run durations were relatively shortas < 10 minutes, making it difficult to apply simply the results to the natural system in the mantle.
In order to overwhelm these problems, we undertook melting experiments using a starting material
similar to primitive mantle composition both in major and trace elements of McDonough and Sun(1995), and changing run duration from 10 to 120 minutes.
The starting material named PML-PM-1 was prepared by mixing Horoman plagioclase lherzolite
and Komuro andesite with approximately 9:1 resulting in primitive mantle composition except forslight depletion of HFSE and slight enrichment of alkaline elements due to the island arc signature of
Komuro andesite. Melting experiments were performed under conditions of 15 GPa and 2400◦ C using
a 6-8 type multi-anvil high pressure apparatus installed at Gakushuin University. Details of the experiments are described in Suzuki et al. (2000). Experimental charges were analyzed by a combination of
SEM-EDX and ion microprobe (Cameca ims 5f) for major elements and trace elements, respectively.
Quantitative analyses of 17 trace elements in majorite and melt with ion microprobe were calibratedusing standard materials such as mantle-derived homogeneous clinopyroxene or garnet, and fused
MORB glasses using a piston cylinder, respectively, which were precisely characterized by ICP-MSand ion microprobe. The concentration levels of trace elements in the standard materials are similar to
those of majorites and melts except for La, reducing matrix effect in the ion probe analyses and thus
capable of measuring sub-ppb concentration level of 10 µm spot in diameter. It was difficult to totally
avoid a large contamination of La from the surrounding LaCrO3 tube, which was used as a thermal
insulator, even though the graphite capsule was further enclosed by Re and MgO cylinder and discs.
However, the contamination of La was consequently reduced to several tens ppm by this treatment,
which is more than two order of magnitude smaller than those in the pervious method.
The D values (Concentration in Majorite/Concentration in Melt) of REE smoothly decrease withincreasing ionic radius. D of La very well fits the extrapolation of middle to light REE values in spiteof extraordinary large contamination of La. However, the results are largely affected by the run duration. That is, the D values of LREE become smaller with increasing run duration (D=0.07 to 0.003 for
La in 10 to 120 minutes runs, respectively), and the LREE/HREE ratios nearly plateau at the run duration more than 60 minutes. Furthermore, the extent of time effect to D values systematically increases
with increasing ionic radius. When compared to the previous experiments and one of the present
experiment carried out run duration shorter than 10 minutes, D values of extremely incompatible elements such as La and Nb in 60 to 120 minutes runs are more than one order of magnitude smaller.
These observations clearly indicate that the D values of trace elements between majorite and melt obtained with shorter experiments (< 10 min.) are out of equilibration so that most of the trace element
partitioning experiments especially for incompatible elements must be re-examined. This is supportedby the imaging observation of secondary ions performed by high-sensitivity ion microprobe, Cameca
ims 1270. That is, the La distribution in majorite grains formed in the shorter run duration show acompositional gradient with lower concentration in the core than the rim, but relatively uniform lower

Magmatic differentiation and the evolution of the mantle and core

149

concentration in the longer runs exceeding 60 minutes. The D values of LREE obtained for majorite
are essentially similar to those for garnet formed under condition of 5 Gpa using JB2 as a starting
material, basalt from Oshima Island, except for Eu negative anomaly with garnet. However, D valuesof HREE for garnet are significantly larger than those of majorit exceeding analytical errors. Thismight be resulted from incorporation of pyroxene component into majorite with increasing pressure.
If this is the case, the role of majorite garnet for the HREE fractionation in the deep in the mantle
decreases with further increasing pressure.
References
McDonough, W.F. and Sun, S-s., 1995. The composition of the Earth. Chem. Geol, 120, 223-253.
Suzuki, T., Akaogi, M. and Nakamura, E., 2000. Partitioning of major elements between garnetstructured minerals and silicate melt at pressure of 3-15 GPa. Phys. Earth Planet. Inter. 120,
79-92.

Session 5

150
5.10

Melting of peridotite under lower mantle conditions and its implication todifferentiation of a deep magma ocean
E. Ito, A. Kubo, T. Katsura, and M. J. Walter

Institute for Study of the Earth Interior, Okayama University, Misasa Tottori-ken 682-0193, Japan

It is generally accepted that the Earth would have been covered by a deep magma ocean in an earlystage of its history. Possible heat sources for such extensive melting are (1) the released gravitationalenergy of accreting planetesimals, (2) that associated with segregation of the central metallic corefrom the uniform primordial earth material, and (3) a giant impact which might trigger formation of
the moon. It is highly likely that some crystal differentiation proceeded in the magma ocean resulting in large scale chemical stratification of the mantle. Melting experiments on mantle material up topressures corresponding to the deep magma ocean may provide important information to verify the
hypothesis. Especially important is proper identification of the liquidus phases and element partitioning between the liquidus phase and melt. In the present study, we have carried out melting experimenton peridotite up to 33 GPa, and propose a working hypothesis for crystal fractionation in deep magma
ocean.
High-pressure experiments were performed using a split-sphere type of high-pressure apparatus
equipped with sintered diamond (SD) anvils [1]. The SD cubes of 14 mm edge length with truncatedcorner of 1.5 or 2.0 mm were employed together with an octahedral pressure medium of magnesia
mixed with 5%Cr2 O3. Powdered peridotite (KR4003 [2]) sample was put directly into a cylindrical
Re heater of 0.7 mm(outer diameter) × 3.8 mm (length), which was set in the octahedron with a
LaCrO3 sleeve. Experimental pressure was determined based on the calibration curve constructed atroom temperature using several fixed points: ZnS (15.6 GPa), GaP (22 GPa), and Zr (33 GPa). Effectof sample heating on the generated pressure was minimized by using a sleeve of 0.7 mm thikness, as
described elsewhere [1]. The sample was heated to ca. 2500◦C for 2-3 min at the prescribed load, andthen quenched by shutting off the electric power to the heater. The recovered products were made topolished sections, which were examined by electron microscopy and then analyzed by the electronprobe micro analyzer.
The melt was identified under the scanning electron microscope by the presence of dendritic texture. Most of the melt concentrated in the central, high temperature region of the charge, which would
have been caused by the dissolution of solid phase components at high temperature and precipitationof those components at lower temperature as a consequence of the saturation gradient due to the steep
temperature gradient through the charge (1-2 ◦ C/µm). This makes location of the liquidus very clear
in the charge, whereas the solidus is defined more ambiguously in terms of grain size and the presenceof psuedomorphs of the starting material. In the super solidus region, phase identification was carriedout based on morphology and chemical composition. The temperature gradient in the charge makes
it possible to establish the crystallization sequence of different phases with decreasing temperaturefrom the liqudus.
Ferropericlase (Fp) is the liquidus phase up to about 30 GPa, which is successively followed bymagnesian perovskite (Mg-Pv) and calcium perovskite (Ca-Pv). Both Fp and Mg-Pv coexist along theliquidus at 31 GPa, indicating multiple saturation of these phases. At 33 GPa, however, the front of Fp
grains moves back from the liquidus to the slightly lower temperature region and Mg-Pv becomes theliquidus phase. Ca-Pv crystallizes at a fairly lower temperature than Fp and Mg-Pv at pressures up to
ca. 29 GPa. However the crystallization temperatures of Fp and Ca-Pv become closer with increasingpressure, and the former might be only a few degrees higher than the latter at 33 GPa. The change inliquidus phase in peridotite found in the present study indicates that the composition of the initial melt
formed by melting in peridotitic mantle deeper than 900 km should be more basic (or more deficientin silica) than the peridotitic composition.

Magmatic differentiation and the evolution of the mantle and core151

Remarkable variation in composition was not observed over the central “melt pools” in the run
products, which have composition very close to the initial peridotitic. However, both Mg-Pv and Fp
on the liquidus have definitely higher Mg/(Mg+Fe) ratios than those located close to the solidus.
Partition coefficients between Mg-Pv, Fp, and melt for Si, Mg, Ca, Fe, and Ni do not show prominentpressure dependence compared with those determined at lower pressure [3]. However, Al, Cr, and Ti
become “incompatible” in Mg-Pv at 33 GPa. It should be also noted that alkaline (K and Na) and rareearth (La) elements strongly prefer Ca-Pv rather than melt.
Using the compositions of liquidus Mg-Pv, Fp, and Ca-Pv at 33 GPa, we have examined differentiation of these phases from a C1 bulk mantle composition [4] in terms of mass balance regression
with respect to Mg, Si, Al, Fe, and Ca. The calculation indicates that preferential subtraction of 34%
Mg-Pv and 3% Ca-Pv yields a residual melt close to peridotitic composition, involvement of Fp being
negligible. The fractionated perovskites would pile up in the bottom of the mantle to form a layer
attaining to a depth ca. 1600 km. Although the perovskite layer may have subsequently interactedwith the upper materials, it could stably lie in the lower half of the mantle for a long term because of
its intrinsically high density. The layer, on the other hand, is characterized by the high capability of
Ca-Pv to accommodate alkaline and rare earth elements as an enriched and heat-producing reservoir[5].
References[1] Ito, E., A. Kubo, T. Katsura, M. Akaogi, and T. Fujita (1998) Geophys. Res. Lett., 25, 821-824.
[2] Xue, X., H. Baadsgaard, A. J. Irving, and C. M. Scarfe (1990) J. Geophys. Res., 95, 15879-15891.
[3] McFarlene, E. a., M. J. Drake, and D. C. Drake (1994) Geochi. Cosmochi. Acta, 58, 5161-5172.
[4] Ohtani, E., T. Kato, and H. Sawamoto (1986) Nature, 322, 352-353.
[5] Kellogg, L. H., B. H. Hager, and R. D. van der Hilst (1999) Science, 283, 1881-1884.

Session 5

152
5.11

Core formation in a reduced magma ocean: new constraints from W, P, Ni and Co
M. Walter
Institute for Study of the Earth’s Interior, Okayama Univ., Misasa, Tottori-ken 682-01, Japan;
walter@misasa.okayama-u.ac.jp

Models for equilibrium core formation in a magma ocean are based on mantle siderophile element(SE) abundances, estimates of bulk Earth SE abundances and a knowledge of metal/silicate partition
coefficients, Dmet/sil . Equilibrium models must account for the effects of pressure (P), temperature(T), silicate and metal melt compositions, and oxygen fugacity (fO2) on Dmet/sil . The most comprehensive model to date is by Righter and Drake (1999: RD99)1, in which all available partitioning data
for several SE’s are empirically fitted to an equation of the form:ln Dmet/sil = alnf O2 + b/T + cP/T + d(nbo/t) + eln(1-Xs) + fln(1-Xc)+ g(5)Based on regressions for five moderately siderophile elements (P, W, Co, Ni and Mo), RD99
concluded that metal-silicate equilibrium may have occurred in a hydrous, magma-ocean at ∼ 27
GPa and 2200K. Final equilibration may have occurred when metal ponded at a solid, perovskite-richmagma ocean floor, and this model has arguably become the standard equilibrium magma-ocean coreformation model. In the RD99 model, fO2 is determined to be about IW-0.4, yielding an FeO content
of more than 20 wt% in the residual bulk silicate Earth, more than a factor of two greater than present
estimates for the upper mantle. Thus, the RD99 model implicitly requires metal/silicate equilibriumin an oxidized magma ocean, with the mantle having become subsequently more reduced.
Using equation 1 to fit partitioning data can be fraught with danger because the valence state of
the SE must be constant over the entire range of conditions being regressed, and lnD must be linearly
dependent on intensive variables such as P and T. Here, we examine new partitioning data for W,
which show that valence state is not constant and the pressure effect is non-linear. We also present
new partitioning data for P that pressure causes P to become considerably more lithophile, an effect
not found previously for this element. We combine our new data for W and P with literature data for
Ni and Co to test magma-ocean core-formation models.
The dependence of Dmet/sil on fO2 can be formulated from thermodynamics as:log D = -x/4 • logfO2 + constant(6)

where the slope (x/4) is a function of the valence state (x) of the metal oxide. This has important
implications for modeling core formation because fO2 is an unconstrained parameter in bulk Earthmodels, and Dmet/sil can be highly dependent on fO2 especially for high valence cations. Thus, it iscritical to know the valence of each SE over a wide range of P and T. We have conducted more than
70 experiments to determine Dmet/sil for W over a range of conditions from 1 to 18 GPa, 2100 - 3000
K, IW-1 to IW-2.5, and in komatiitic to peridotitic melt compositions. Results for W are summarized
as follows: 1) At 1 and 2 GPa, regressions clearly indicate that W occurs in a 5+ oxidation state in
the silicate. 2) At 6 - 18 GPa regressions show that W occurs in 4+ oxidation state. 3) The effect ofpressure on DW is highly non-linear, with DW increasing up to about 5 GPa, but decreasing monotonically at higher P. This change if pressure effect is probably related to the change in valence. 4) DW
is highly dependent on silicate melt composition. We have conducted 16 experiments to determine
Dmet/sil for P over a range of conditions from 2 to 18 GPa, 2100 - 2700 K, IW-1.5 to IW-2.5, andin komatiitic to peridotitic melt compositions. Results are summarized as follows: 1) P apparently
occurs in a 4+ valence state over the entire range of conditions. 2) T causes an increase in DP whereaspressure causes a uniform decrease in DP. 3) DP is moderately dependent on silicate composition. In
addition, partitioning data for Ni and Co consistently indicate that these elements, like Fe, exist in a
2+ valence state over a wide range of P and T at fO2 below the IW buffer2.

Magmatic differentiation and the evolution of the mantle and core153

We have tested equilibrium core formation in a magma-ocean using our new partitioning data for
W and P, the latest literature data for Ni and Co, and a bulk-Earth chemical model3 . First, in a test of
the model conditions of RD99 (27 GPa, 2250 K, IW-0.4, nbo/t=2.7), we find that Ni and Co depletions
can be successfully matched within about a factor of 2. However, according to our new data, at theRD99 conditions W would be about a factor of 30 and P about a factor of 700 too lithophile to accountfor mantle depletions! Even though our modelling is simplified because it does not presently take into
account the effects of S or C in the metal, it is extremely unlikely that the effects of these minor
components can account for the severe mismatches in predicted and observed depletions.
Using a trial and error approach, a set of conditions is located that can account for the observed
depletions of Ni and Co within a factor of 2, and W and P within a factor of about 3. The conditions
are: 40 GPa, 4000 K, nbo/t=2.8 (peridotite), and fO2 of IW-2.6. A similar P - T condition has been
determined based on data for Ni and Co in isolation2. This result must be considered with caution,
however, because the effects of pressure and temperature for all elements have been extrapolated
considerably outside the range of the data, requiring further experiments for verification.
Taken at face value this new model indicates core-mantle equilibration in a magma ocean at a
depth of about 1200 kms. Such a depth is not too dissimilar to a recent tomographic estimate for thedepth of topography in the lower mantle (relic magma ocean floor?)4. Further, the relatively low fO2
predicted in our model would leave a mantle with only about 4 wt% FeO, or about half the estimatedvalue of the present upper mantle, so the mantle would have become subsequently oxidized. The low
fO2 predicted in our model is required in order to make W and P sufficiently siderophile, counteracting
the effect of pressure on these elements. Corroborating evidence for a reduced Earth during coreformation comes from recent partitioning data, which show that mantle depletions in V and Nb may
be due to sequestering of these elements into the core if conditions were highly reduced (i.e. ∼ IW-3)5. Further, at such reduced conditions considerable Si can dissolve into the core, accounting at leastin part for the light element in the core as well as the sub-chondritic Si/Mg ratio of the upper mantle.
[1] Righter, K. and Drake, M., EPSL, 174, 383, 1999.
[2] Li, J. and Agee, C., GCA, 65, 1821, 2001.
[3] McDonough, W., Chemical Geology, 120, 223, 1995.
[4] van der Hilst, R. and Karason, H., Science, 283, 1885, 1999.
[5] Wade, J., and Wood, B., Nature, 2001.

154

Session 5

5.12

Mechanisms of metal transport during formation of the earth’s core: constraints fromthe kinetics of metal-silicate reactions

D.C. Rubie1 , C. Holzapfel1, J. Reid1 , S. Fortenfant1, H.J. Melosh2 , B.T. Poe1 , D.J. Frost1, K.
Righter2
1 Bayerisches Geoinstitut, Universitaet Bayreuth, D-95440 Bayreuth, Germany
2

Lunar and Planetary Laboratory, University of Arizona, Tucson, AZ 85721, USA

The concentrations of siderophile elements in the Earth’s mantle are believed to have resulted
from reactions between liquid metal and silicates/oxides during the early history of the Earth. Basedon measured metal-silicate partition coefficients, siderophile element concentrations are too large to
be explained by chemical equilibrium during the separation of liquid metal from silicates at low
pressures. However, partition coefficients for elements such as Ni and Co decrease with increasingpressure at an oxygen fugacity just below the IW buffer and reach appropriate values at 25-30 GPa.
It has therefore been proposed that the concentrations of such elements in the mantle are the result of
metal-silicate equilibration at the base of a magma ocean 700-800 km deep (e.g. Li and Agee, 1996;
Righter et al., 1997). According to this model, liquid metal ponds at the base of a silicate magma
ocean and, after equilibrating chemically, descends as diapirs to form the Earth’s core. In the case thata magma ocean stage did not exist on the Earth (e.g. Jones and Palme, 2000), siderophile element
concentrations must have resulted from reaction between liquid metal and crystalline silicates/oxides.
The role of reaction kinetics in controlling the partitioning of siderophile elements during coreformation has been discussed briefly by Stevenson (1990) and Karato and Murthy (1997). We areinvestigating kinetics between liquid metal and crystalline mantle phases and between liquid metal
and liquid silicate, both experimentally and theoretically. Multianvil experiments on liquid metalmagnesiowüstite at 10 GPa and 2200 K show that equilibration rates are controlled by the diffusionof siderophile elements in the crystalline phase. Diffusion data (e.g. Fe-Mg interdiffusion) for mantle
minerals can therefore be used to model rates of metal-silicate reactions during core formation. Highpressure diffusion data for olivine, wadsleyite, magnesiowustite and (Mg,Fe)SiO3 perovskite showthat rates of reaction must vary strongly with depth. This is because diffusivities differ strongly between these phases: Fe-Mg diffusion is relatively fast in wadsleyite whereas in perovskite it is several
orders of magnitude slower than in the other phases. Considering likely rates of grain-scale metal percolation, migration of a large fraction of the liquid metal to the proto-core will occur with very limited
chemical exchange with crystalline mantle. Only the last small fraction of metal will percolate slowly
enough to allow equilibrium to be reached.
For modeling the kinetics of equilibration between liquid metal and liquid silicate, experimental
measurements of viscosity and diffusivity are being performed. Model reaction rates are controlled
by chemical diffusion through a silicate liquid boundary layer adjacent to the liquid metal. In thecase of reaction between a layer of segregated liquid metal and overlying silicate liquid at the baseof a deep magma ocean, the convection velocity of the magma ocean controls both the equilibration
rate and the rate at which the magma ocean cools. Results indicate that time scales of chemical
equilibration are several orders of magnitude longer than the time scales of cooling and crystallizationof the magma ocean. In an alternative model, falling liquid metal droplets and silicate melt react in aconvecting magma ocean. In this case, the droplet size and settling velocity are critical parameters. Fora typical silicate liquid viscosity at 2600 K, the stable droplet diameter is estimated to be ∼1 cm with asettling velocity of ∼0.5 m/s. Using such parameters, liquid metal droplets are predicted to equilibratechemically after settling a distance of <200 m. These models indicate that the concentrations ofmoderately siderophile elements in the mantle are most likely the result of chemical interaction of asilicate magma ocean with small metal droplets but not with a layer of liquid metal that has alreadysegregated at the base of the magma ocean.

Magmatic differentiation and the evolution of the mantle and core

155

References
Jones JH, Palme H (2000) Geochemical constraints on the origin of the Earth and Moon. In: Originof the Earth and Moon, Canup, RM, Righter K (eds), Univ. Arizona Press, Tucson, 197-216.
Karato S, Murthy VR (1997) Core formation and chemical equilibration in the Earth - I. Physicalconsiderations. Phys. Earth Plan. Int. 100, 61-79.
Li J, Agee CB (1996) Geochemistry of mantle-core differentiation at high pressure. Nature 381, 686689.
Righter K, Drake MJ, Yaxley G (1997) Prediction of siderophile element/silicate partition coefficientsto 20 GPa and 2800◦C: the effects of pressure, temperature, oxygen fugacity, and silicate andmetallic melt compositions. Phys. Earth Planet Int. 100, 115-134.
Stevenson, DJ (1990) Fluid dynamics of core formation. In: Origin of the Earth, Newsom HE, JonesJH (eds), Oxford Univ., New York, 231-249.

Session 5

156
5.13

Effect of Light Elements on Viscosity of Liquid Iron-Alloy

H. Terasaki1, T. Kato1 , S. Urakawa2, K. Funakoshi3, A. Suzuki4,
K. Sato5 , M. Hasegawa2
1 Gescience Inst., Univ. of Tsukuba, 2 Dept. of Earth Sc., Okayama Univ., 3 JASRI,
4 Inst. of Mineral. Petrol. Economic geol., Tohoku Univ., 5 Foundation for Advancement of Int. Sc.

Introduction
The viscosity of liquid iron-alloy is one of the most fundamental physical properties to investigate
the convectional behavior of the Earth’s outer core. The outer core consists mainly of liquid Fe-Ni
alloy with a significant amount of the light elements such as S, O, C, Si, and H [1]. Dissolved lightelements must influence the physical properties of liquid iron under pressure. We have reported theviscosity of liquid Fe [2] and Fe-FeS [3] up to 7 GPa. In this study, we measure the viscosity of theliquid in the system of Fe-Fe3C using X-ray radiography falling sphere method and investigate theeffect of sulfur and carbon on the viscosity of liquid iron-alloy.
Experimental Procedure
High pressure and temperature experiments were performed by MA-8 multianvil press (SPEED1500) installed at BL04B1, SPring-8, Japan. Starting material was a mixture of Fe (4N) and Fe3 C(3N) powder. Experimental pressure was determined from the Birch-Murnaghan equation of state forh-BN and MgO [4], which were used as the sample capsule and X-ray window, respectively. Temperature was monitored just above the capsule inside the heater by the W3%Re-W25%Re thermocouple.
Detailed cell description was shown in ref. 3. Radiography was observed by the high-speed exposure
CCD camera (C4880-80-14A) with resolution of about 5 µm and with the exposure time of 11-21
ms. Images were recorded with computer hard disk at the rate of 30 ms. We used Pt and Pt13%Rh
spheres with diameter of 100-150 µm, as viscosity markers. A double layered sample assembly was
adopted to avoid influence of the partial melting of the sample during the viscosity measurement. The
sphere was set in the silicate layer, which overlay the Fe- Fe3C layer in a sample capsule [3]. We usedeither albite (NaAlSi3O8 ) powder as a material in the silicate layer. These spheres were imaged as
the dark shadows due to its high X-ray absorption coefficient. Settling velocity of the marker sphere
was obtained from the continuous radiographic images (Fig.1), and the viscosity coefficients were
calculated using Stokes’ equation with wall and end correction.

Magmatic differentiation and the evolution of the mantle and core

157

Results and Discussion
The viscosity coefficients of the liquid Fe-Fe3C are 0.004-0.018 Pa s at 3 GPa, 1605-1923 K.
Viscosity coefficient of eutectic Fe-Fe3C is smaller than that of liquid Fe and Fe3C end member.
It seems to take minimum value at the eutectic point (C=4 wt%). This tendency is similar to that ofliquid Fe-FeS alloy as shown in Fig.2. In case of the Fe-FeS system, it can be interpreted as the changeof near neighbor atomic pair from Fe-Fe pair to Fe-S pair at higher sulfur content [5]. Since it may
be explained by the similar structural change of the liquid Fe-Fe3C, the structural analysis of liquid
Fe-Fe3C is required. Pressure effect of viscosity can be estimated from present high pressure viscosity
data together with the data measured at ambient pressure. Activation volume (pressure dependence)of liquid iron is larger than that of liquid Fe-FeS and Fe-Fe3C.
In the results of radiographic observation, a large deformation of the boundary between the sampleiron-alloy and trapper silicate was found during sample melting in the system of Fe and Fe-Fe3C.
However, such deformation was not observed in the system of Fe-FeS. This phenomenon indicatesthat surface tension of liquid iron-alloy decreases with increasing sulfur content. On the other hand,addition of carbon shows little influence on the surface tension.
References[1] J.P. Poirier, Phys. Earth Planet. Int. 85, 319-337 (1994).
[2] H. Terasaki et al., Submitted[3] H. Terasaki et al., Earth Planet. Sci. Lett. 190, 93 (2001).
[4] J. C. Jamieson et al., in High-Pressure Research in Geophysics, S. Akimoto, M. H. Manghani, Eds.
(Center for Academic Publishing, Tokyo, 1982), pp. 27-48.
[5] S. Urakawa et al., Rev. High Pressure Sci. Technol. 7, 286-288 (1998).

Session 5

158
5.14

Critical melt fraction for percolation of core-forming materials
Takashi Yoshino, Michael J. Walter and Tomoo KatsuraInstitute for Study of the Earth’s Interior, Okayama University

The formation of an iron core and silicate mantle was a major differentiation event in the early history of the Earth. In a homogeneously accreting Earth silicates and metal alloys are initially intimately
mixed, so core formation requires a mechanism for separating and mobilizing the iron alloy. For thisreason, there has been interest in the wetting properties of liquid iron or alloy melts in silicate mineralmatrix. Previous experimental studies concerned with the wetting behavior of alloy melts were basedon the determination of dihedral angles under variable conditions. Most of the results show large dihedral angles of over 90 deg for Fe metal in olivine-rich matrix, suggesting a non-wetting propertyof iron-rich melts. Although it seems reasonable from statistical angle measurements to conclude thatiron-rich metallic melts cannot have segregated from an olivine-rich silicate mantle, a certain amountof liquid metal will connect along grain boundary when the dihedral angle is above the critical valueof 60 deg. In this case, an excess of melt over the critical melt fraction (CMF) can create permeability
and segregate from the silicate matrix leaving residual melts trapped at grain boundaries and as inclusions in silicates. The stranded melts may influence the absolute and relative abundance of siderophileelements in the Earth’s mantle. Therefore, it is important to determine the CMF for melt permeability
in 3 dimension.
In order to assess the CMF in metal-silicate system and the feasibility of percolation as a coreformation process, we performed experiments over a range of metal/alloy mass fraction up to 0.7
in San Carlos olivine + sulfide or San Carlos olivine + iron mixtures. Usage of a starting material
composed of San Carlos olivine including variable proportions of synthetic sulfide and metal alloy
allows examinations of the CMF with respect to a silicate mineralogy analogous to Earth’s upper
mantle. Sample materials are packed into graphite capsule. Experiments were performed using apiston cylinder apparatus at 1573-1673K and 1.5GPa. Electrical conductivity of a part of the retrieved
sample was measured by digital voltmeter to access metal/sulfide connectivity under room condition.
From the above analysis, the lower bound of the CMF was determined.
Run products of olivine-iron mixture are composed of olivine, magnesiowustite (mw) and ironand run products of olivine-sulfide mixture are composed of olivine and sulfide. In both systems,
electrical conductivity was lost when percolation drainage had reduced the alloy volume to ∼20%,
indicating a lack of interconnectivity (Fig. 1). This threshold value corresponds to an interconnection
threshold for a nearly 150 deg dihedral angle, based on a prediction for body-centered cubic packingof plane-faced tetrakaidecahedra (von Bargen and Waff, 1986). However, the previously reported
dihedral angles at the junction of metallosulfide melt with silicate are generally smaller, in the rangeof 65-115 deg. In addition, the percolation threshold does not change significantly with the sulfur
content in the melt, although the dihedral angles are very sensitive to the amount of S in the melt.
These inconsistencies may indicate that the system did not establish textural equilibrium during the
3day run duration. Figure 1 shows power law correlation between volume fraction and resistivity. The
slope of ∼-7 in the olivine+mw+iron system is slightly larger than that in system including sulfide(∼-5). Difference in slope may reflect effect of magnesiowustite on bulk resistivity.
The results indicate a critical condition at which the metal segregation process in silicate solid
matrix switch from grain boundary percolation to gravitational sinking of molten iron droplets. When
metallosulfide fraction is over ∼20 vol.%, interconnected molten metal will be able to segregate byporous flow. These considerations may provide new insights into the core formation process in theearly history of the Earth and other planetary bodies, and yield important clues for understanding theabsolute and relative abundances of siderophile elements in the Earth’s mantle.

Magmatic differentiation and the evolution of the mantle and core159

Session 5

160
5.15

Geomagnetic field model for the last 5 my and conditions at core-mantle boundary
Tadahiro Hatakeyama and Masaru Kono
Institute for Study of the Earth’s Interior, Okayama University, Japan

We have proposed a geomagnetic field model [1] of stable magnetic periods for the last 5 million
years which consists of time-averaged field (TAF) and paleosecular variation (PSV). In this model the
TAF and PSV was calculated with considering the nonlinear relationship between model parameters
and observations derived from paleomagnetic studies.
The determined TAF models both of normal and reverse geomagnetic polarities show that thegeocenteric axial quadrupole term, g20 , is significantly larger than other nondipole terms and has same
sign as the geocentric axial dipole, g10 . This feature indicates that the geomagnetic equator is shifted
northward from the geographic equator and that “the VGP (virtual geomagnetic pole) far-side effect”
seen in the paleomagnetic is caused mainly by g20 component in both polarity periods instead of
the higher order terms, which is consistent with previous zonal TAF models [2]. When we consider
the symmetry with respect to the field (B) in the magnetic induction equations, it is likely that thisphenomenon is due to the boundary conditions of the fluid outer core such as heat flows.
On the other hand, the geocentric axial octupole term, g30 , was also thought as a remarkable component in a previous TAF model [3]. However, this term is not excellent in our TAF model because
the PSV effect [4] on the mean direction affects mainly this and the effect seems to mislead us intooverestimating “apparent” g30 component when it is ignored in the TAF model.
In this study, we will compare the characteristics of the TAF and PSV models derived from paleomagnetic data to the results of the geodynamo simulation with using heterogeneous boundary
conditions of heat flux [5] and discuss the role of the condition at the core-mantle boundary region.
References[1] Hatakeyama T. and M. Kono, “A self-consistent time-averaged geomagnetic field and paleosecularvariation model for the last 5 My”, Abst. IAGA-IASPEI joint scientific assembly, 73, 2001
[2] Wilson, R. L., “Permanent aspects of the Earth’s non-dipole magnetic field over upper tertiary
times”, Geophys. J. R. astr. Soc., 19, 417–437, 1970.
[3] McElhinny, M. W., P. W. McFadden and R. T. Merrill, “The time-averaged paleomagnetic field
0–5 Ma”, J. Geophys. Res., 101, 25007-25027, 1996
[4] Hatakeyama T. and M. Kono, “Shift of the mean magnetic field values: Effect of scatter due tosecular variation”, Earth Planets Space, 53, 31–44, 2001
[5] Glatzmaier, G. A., R. S. Coe, L. Hongre and P. H. Roberts, “The role fo the Earth’s mantle incontrolling the frequency of geomagnetic reversals”, Nature, 401, 885–890, 1999

October 5 (Friday) 13:30 – 16:30

Session 6

Future Perspectives

Keynote: Dave Mao, Geophysical Laboratory, Carnegie Institution
Chair: Dave Rubie, Bayerisches Geoinstitut, University of Beyreuth

161

Future Perspectives
6.3

163

Mantle convection as a tool for understanding the links among the various geologicalphenomena
S. Honda1 and M. Yoshida2
1 Hiroshima University <satoru@yesterday.geol.sci.hiroshima-u.ac.jp>
2 Tokyo University

A successful model of mantle convection must explain observed geological phenomena and, thus,
it will clarify their causal links based on physical and chemical principles. Current active research
field of mantle convection may be categorized as two major topics, that is, (1) a construction ofself-consistent model of mantle convection which shows a plate-like behavior (e.g., Tackley [2000])and (2) an understanding of the dynamics of deep interior of the Earth (e.g. Forte and Mitrovica[2001]). Both topics have serious difficulties to be overcome. Main difficulties of (1) are an insufficientknowledge and a troublesome numerical treatment of low temperature rheology and those of (2) are
the insufficient observational constraints on the deep mantle process. In this presentation, we show
our current research efforts related to these topics.
Major features of the plate tectonic movement may be the existence of (1) the rigid plate interior,(2) the transform faults and (3) the asymmetrical subduction. Consider 2D case, that is, ignore (2), therigid plate movement can be modeled relatively easily. The key point is the weakening mechanismof plate boundaries, which may be caused by a yielding, at least, as a first-order approximation.
However, the modeling of the third point turns out to be fairly difficult. This may be caused by thecoarse mesh, which cannot resolve the evolution of thin weak zone, and/or the absence of possible
asymmetric difference in physical properties around the subduction zones. To achieve the one-sided
subduction, we assume the lateral variation of the yield stress (Honda et al. [2000]), in which theyield stress of overlying lithosphere is higher than that of the subducting lithosphere. Generally, weobserve an inclined subduction of the cold lithosphere, although the long-existing “favorable” onesided subduction appears to be difficult. Further modeling efforts like the usage of the finer meshmay be required. We also observe that the lateral variation of yield stress controls the place of future
subduction.
A modeling the deep mantle process may be constrained by the present geophysical data, such as
the plate movement, gravity and seismic tomography, and the mixing rate inferred from the geochemical data. Recently, we have started the 3D modeling of present mantle convection hoping that we canconstrain the whole mantle process (Yoshida et al. [2001]). Plate movements are strongly controlled
by the lateral variation of viscosity in the lithosphere. Using the radially stratified viscosity except thelithosphere and the mantle density anomalies inferred from the subduction history, we found that the
weak plate margin changes both the plate velocity and geoid substantially as Zhong and Davies [1999]
also pointed out. This suggests a trade-off between the rheology of the lithosphere and the viscosity
stratification in the lower mantle in explaining the geophysical data. Since the viscosity distribution
in the lower mantle is a key to understand the nature of the whole mantle flow (Forte et al. [2001]),
the rheology of the lithosphere is also important to clarify the view of mantle-wide process.
References
Forte, A. M., and J. X. Mitrovica, Deep-mantle high-viscosity flow and thermochemical structure
inferred from seismic and geodynamic data, Nature, 410, 1049-1056, 2001.
Honda, S., T. Nakakuki, Y. Tatsumi and T. Eguchi, A simple model of mantle convection including apast history of yielding, Geophys. Res. Lett., 27, 1559-1562, 2000.
Tackley, P. J., Self-consistent generation of tectonic plates in time-dependent, three-dimensional mantle convection simulations 1. Pseudoplastic yielding, Geochemistry Geophysics Geosystem, 1,
2000.

164

Session 6

Yoshida, M., S. Honda, M. Kido and Y. Iwase, Numerical simulation for the prediction of the platemotions: Effects of lateral viscosity variations in the lithosphere, Earth Planets Space, 53, 709721, 2001.
Zhong, S., and G. F. Davies, Effects of plate and slab viscosities on the geoid, Earth Planet. Sci. Lett.,
170, 487-496, 1999.

Future Perspectives
6.4

165
High-pressure crystallography in mineral physics

Takamitsu Yamanaka∗ , A. Yoshiasa, O. Otaka and N. Nagai
Department of Earth and Space Science Graduate School of Science Osaka University 1-1 Machikaneyama
Toyonaka, Osaka 560-0043 Japan

1．Purpose of the high-pressure crystallography
Numerous studies of bulk modulus and equation state (EOS) have been repeatedly carried out for
geophysical interests. A dynamical process is one of the significant projects in high-pressure crystallography. Many aspects of structure changes under compression, such as phase transformation, lattice
deformation, cation ordering, decomposition, amorphization and solid reactions. Phase transitions under compression due to lattice instability, electronic state change and magnetic spin ordering can beelucidated by X-ray diffractometry, absorption and resonance using diamond anvil cell together withother X-ray optical measurements. Developments of many types of detectors and synchrotron radiation accelerate kinetic studies and dynamical studies of those changes. Hydrostaticitiy gives an effecton the phase transformation, under shear stress, which does not necessarily gives a thermodynamically stable phase. Nonhydrostaticity of compression brings metastable states or intermediate states,
because they are formed in the dynamical process. Besides the stress-strain relation, the crystallite
size is an effective parameter of phase transition under compression. Metastable or unstable phases
have been formed by the Oswald rule.
Many phase transformations and the phase stabilities have been undertaken under static highpressure conditions during the last twenty years, Our previous studies related the phase transformation
are shown in Table 1.
Table 1. Phase transformation under compression
lattice change
electronic state change
spin-lattice interaction
no coordination change
charge disproportion
Jahn-Teller transition
3+
5+
4+
lattice type change
Fe + Fe ↔ 2Fe
molecular dissociation Au1+ + Au3+ ↔ 2Au2+
cation ordor ↔ disorderband overlap
high-low spin transition
polyhedral joint
density of state
Mott transition
coordination changes↔, f↔ orbital change spin-Perierls transition
Martensitic transition
orbital rearrange
2. Development using synchrotron radiation for structure study
Since synchrotron radiation at Photon Factory (PF) at at Tsukuba and SPring-8 in Japan has beenutilized, structure study as well as phase study becomes more fashion under high pressure by theextensive studies using DAC or mutianvil apparatus. For example in SPring-8, seven beam lines areavailable or designed for high-pressure studies by means of XAFS, 7 axes diffractometer for structure
analysis, laser heating, low temperature, IR and Mösbauer systems. High-pressure studies often encountered several difficulties for in situ observation. Significant progress in in situ diffraction studiesof earth’s interior under compression and high temperature has been made using synchrotron radiation facilities.
3. Time-resolved observation for kinetic study
The transition rate can be determined by profile fitting method. In order to establish the kinetic
study of the pressure-induced transformation present devises performed an in situ and time dependent measurement of diffraction intensity with high accuracy. Time-resolved and in situ diffraction
studies using synchrotron radiation give a brilliant light to the dynamical observation of pressureinduced structure changes. Then this system offers more reliable phase identification, lattice constants

Session 6

166

and diffraction intensity. Consequently structure information under pressure becomes noticeably improved.
Time resolved diffraction study of the phase transition becomes possible by combination of astrong source with quick photon counting system, because the time dependent structure analysis requires the repeating diffraction measurement with a short time interval.
4. Variety of structure transition with nonhydrostaticity
Phase transformation under compression is not necessary to produce a thermodynamically stablephase. From the viewpoint of statistical process, macroscopic time and space average, of structuretransition under pressure, virial theorem proposes the vision of the structure stability under the equilibrium condition between summation of interatomic potential and surrounding pressure Pext aroundthe crystal.
Pext =

1   ∂Ψij
NkB T
−

(−

V
3V
∂rij

V:volume N:number of particles Ψij :interatomic potential rij :distance between i and j particle

Ostwald step rule indicates that transformation path is not only one. Metastable phase having asmall energy barrier, i.e. activation energy lower than that of stable phase is possibly to turn up during
the transformation process. Especially anisotropic stress easily accelerates several steps of metastablephases. The shear stress Pij induced from the nonhydrostatic compression brings a lattice deformation and gliding of slabs such as martensic transition in the random fashion. The phase transition isprobable by the non-diffusion transformation mechanism under stress field.
MD calculation reproduces the pressure-induced transformations under any shear stress conditionand has an anisotropic property in the structure. Born criterion gives the phase stability from themutual relation in elastic constants Cij . MD calculation suggests a mechanism of the pressure-induced
modification, which is a precursor atomic movement on the transition process.
5. Charge density distribution analysis.
Crystal structure analyses under nonambient conditions have been significant subjects not only
for geosciences but also for material science. Crystal structure and physical properties under extreme
conditions have been elucidated by single-crystal diffraction studies. Combined with heating system and DAC, high-temperature and high-pressure measurements have attracted an interest of X-ray
crystallography. Synchrotron radiation facilities have accelerated the high-pressure structure study,
because of their great advantages for x-ray diffraction studies at nonambient conditions.
Single-crystal structure refinements under high-pressure encountered many difficulties such as
nonhydrostaticity, large blind region due to pressure cell, large X-ray absorption and limitation ofcompression. Our new system solved these difficulties and made it possible to discuss the electron
density distribution under pressures up to 50GPa.
Valence electron density distribution as a function of pressure gives a brilliant implication to understand the structure stability and phase transformation mechanism. Localization of valence electron
inducing ionic character rather than covalent and less population of bonding electron with elevatingpressure up to 30GPa were found by charge density analysis of stishovite. Electron density of stateobtained from molecular orbital calculation accords with the above results. Besides the above electron
density distribution under pressure, effective charge and dipole moment give insight into the crystal
chemistry and bonding character under compression.

Future Perspectives
6.5

167

Dynamo process in the core under varying energy flow conditions
Masaru Kono

Institute for Study of the Earth’s Interior, Okayama University, Misasa, Tottori-ken 682-0193, Japan. email:mkono@misasa.okayama-u.ac.jp

In the outer core of the Earth, thermal or compositional convection takes place. As the electricalconductivity of the fluid core is high enough, the convection induces electric current and thus dynamo
action takes place. In the last five years or so, simulation of realistic dynamos became possible. These
dynamo models are three-dimensional and fully nonlinear, and do not make any arbitrary assumptions. The parameter range of the models, especially the Ekman number and the Rayleigh number,is not quite that of the Earth (Figure 1). A direct comparison of geomagnetic and paleomagnetic
observations with simulation results is therefore not quite satisfactory. In spite of these constraints,
comparisons have been made between observations and model calculations, and in many cases very
interesting results were obtained. Thus it seems that the results from model studies provide us useful
information for considering the real state of the Earth’s core.

1019
1018
Earth
1017

KS95
GR95

1016
1015

RaM

GR96

106

KB97
KAK97

1014

Rayleigh Number Ra

1013

COG98
SK99

104

GBT99
1012
1011

KMH99
Benchmark

102

1010
109

100

108
107

10-2

106
105
104

Rac by Jones et al. (2000)
Pr = 101
Pr = 100
Pr = 10-1
Pr = 10-2

103
102
101
10-10

10-9

10-8

10-7

10-6

10-5

10-4

10-3

10-2

10-1

Ekman Number E

Figure 1. Relation between the Ekman number E and Rayleigh number Ra in various MHD dynamo models.
Thick lines indicate the theoretical critical Rayleigh numbers, and thin lines show the modified Rayleighnumber RaM = E · Ra. The value for the Earth is based on an eddy viscosity of 1 m2 s−1 .

Session 6

168

Through the model studies, it became clear that the Ekman number (which signifies the Coriolis
force in comparison with the viscous force) plays the most important role (Figure 2). Also importantis the mode of energy transfer. The flow of energy is controled by the thermal structure of the mantle
just above the core-mantle boundary (CMB). Because the thermal diffusion time is very long in thecore, while any temperature difference at the CMB will be equilibrated on the core side in very shorttime, the temperature heterogeneity in the deep mantle will determine the heat flow from the Earth’s
core. As the lower mantle near the CMB is very heterogeneous (such as D ), the heat flow from thecore must also be quite inhomogeneous. I will discuss about the possibility of looking for changing
energy flow conditions based on these findings.

(a) Benchmark(b) SK99(c) GR96

EquatorCore surface

Figure 2. Snapshots of the flow in the equatorial plane (left) and in the Ekman layer below the CMB (right,
r/rc = 0.98R 0) in three typical dynamo models. Benchmark is a slow rotator, SK99 (Sakuraba and Kono,
1999) is an intermediate rotator, and GR96 (Glatzmaier and Roberts, 1966) is a fast rotator. Color shade andcontours give the flux perpendicular to the plane under consideration (Vz on the left and Vr on the right).
Arrows show the velocity component in the plane (Vs and Vφ on the left and Vθ and Vφ on the right).

Future Perspectives
6.6

169

Comprehensive geochemical analyses of extremely small amounts of terrestrial andextraterrestrial materials for sample-return missions

Eizo Nakamura, Akio Makishima, Takuya Moriguti, Katsura Kobayashi, Chie Sakaguchi, Tetsuya
Yokoyama, Ryoji Tanaka, Takeshi Kuritani & Hiroyuki Takei
The Pheasant Memorial Laboratory for Geochemistry and Cosmochemistry (PML), Institute for Study of the
Earth’s Interior, Okayama University at Misasa, Tottori 682-0193, Japan

We participated in the analytical competition related to the sample-return mission, MUSES-C of
the ISAS to be launched in November 2002 and concluded in June 2006. In this competition, we
determined as many major and trace element abundances and isotopes as possible, for both two powdered samples provided by ISAS and chondrules separated from Allende meteorite as an additional
demonstration to show the analytical capabilities for small fragments.
100 mg of each the samples were split into two fractions of approximately 30 and 70 mg forthe determinations of elemental abundances and the isotope analyses, respectively. The bulk concentrations of 55 elements in each 30 mg sample were analyzed using both quadruple-type ICP-MS andsector-type ICP-MS with analytical uncertainties better than 10 % (1?). Boron, Pb, Li, Rb, Sr, Sm and
Nd were successively separated from each 70 mg of sample using some nearly developed integrated
multi-ion exchange column chemistry, and then these isotopes were determined by TIMS with analytical errors similar to those obtained in our laboratory for terrestrial samples. Strontium, Nd, Re and
Os isotopic compositions were also analyzed on remaining aliquots of the sample solutions used fortrace element analyses by ICP-MS. The results obtained in this study indicate that samples 1C and2C are probably ordinary chondrite and the Allende meteorite, respectively. Extreme W, Ta and Nbenrichments in sample 1C were probably caused by contamination during sample preparation.
Chondrules were separated from Allende meteorite and then sliced into three pieces using a dicing
saw. The center parts of the slices were used to determine mineral phases by compositional mapping,
and to obtain major and trace element compositions (30 elements) by a combination of optical microscopy, SEM-EDX and ion microprobe. The outer slices, for which sample sizes were < 1 mg, werechemically treated to obtain the bulk trace element and isotope compositions of chondrules. For theseslices, 24 trace elements and Sr and Nd isotopic compositions were determined with extremely small
analytical uncertainties using ICP-MS and TIMS, respectively.
The analytical competition, using extremely small amounts of test samples has allowed a demonstration of the analytical capabilities of our geochemical laboratory, the PML, that could be applied
to the samples to be returned from an asteroid by the MUSES-C mission. Most of the mineralogicaland chemical data can be obtained for extremely small amounts of samples and with high spatial
resolution. Japanese analytical capabilities will be further improved by repeating this type of competition and evaluation, and by incorporating other analytical techniques such as light element isotope
analyses and age dating using ion and laser microprobes. Such improvements, aiming to analyze
samples with extremely small sample sizes should also make it possible to obtain comprehensive mineralogical and geochemical analyses with high spatial resolution for precious samples returned from
Mars.

170

Kurayoshi Symposium

6.7

Zircon age dating by dynamic multi-collection method using high resolution secondary
ion mass spectrometer
Katsura Kobayashi, Tomohiro Usui and Eizo Nakamura

Pheasant Memorial Laboratory, Institute for Study of the Earth’s Interior, Okayama University at Misasa

We report here development of zircon age dating method using a High Resolution Secondary Ion
Mass Spectrometer (HR-SIMS), Cameca ims-1270 with multi-collection system. The zircon datingby an ion microprobe has been mainly developed by SHRIMP group at Australian National University (e.g., Compston et al., 1983). Our method is totally different from their method in two features,i) simultaneous determination of Pb isotope composition by multi-collection system, ii) using of thehighly focused primary ion beam. The multi-collection method makes possible to cancel instrumental
fluctuations, mainly caused by the instability of the primary beam, and reduce the analytical duration. Our HR-SIMS equips 5 electron multipliers for detection of the secondary ion beam, and it
is available to collect 204 Pb+ , 206 Pb+ , 207 Pb+ and 208 Pb+ intensities, simultaneously. The configuration of detectors and the analytical condition are shown in Fig. 1. Using this condition, we obtained
207
Pb/206 Pb = 0.07557 ± 11, corresponding to 1073 ± 17 Ma (2σ, n=10), from Geostandard 91500
Zircon, which has 207 Pb/206 Pb = 0.07488 ± 1 and 1065 Ma by TIMS conventional analyses (Widenbeck et al., 1995). We also tried to apply the focused beam with (30 µm)2 pre-spattering substitute forKohler mode illumination beam, which typically provides over ∼30 µm spot diameter (by Cameca
ims-1270 at Swedish Museum of National History, Whitehouse et al., 1997). Squared pre-spattering
method effectively reduced contamination from surface Pb blank. The identical 207 Pb/206 Pb ratios tothose of the Kohler mode were obtained within the analytical errors from natural zircon in Sli Lanka
with 10 ∼ 1 nA range of the primary beam current, corresponding to 13 ∼ 3 µm in diameter (Fig.
2). This result indicates that our method is applicable for zircon dating of tiny sample (under 10 µm
diameter), which occurs commonly in natural samples, and for more precise analyses of zoned structure formed by episodes of growth and/or recrystallization. In order to support the analysis under highspatial resolution, we have also developed VISUAL STAGE system. This system provides identicalspatial coordinates on every image by an optical microscope, SEM, and SIMS. Using this system, we
carried on detailed descriptions by these apparatus before the zircon dating by the ion probe.
References
Widenbeck, M. et al. (1995). Three natural zircon standards for U-Th-Pb, Lu-Hf, trace element and
REE analysis, Geostand. Newslett, 19, 1-23.
Whitehouse, M. J. et al. (1997). Ion microprobe U-Pb zircon geochronology and correlation of Archaean gneisses from the Lewisian Complex of Gruinard Bay, northwestern Scotland, Geochim.
Cosmochim. Acta, 61, 4429-4438.

October 2 – 5, 2001

171

List of Participants
Aizawa, Yoshitaka
aizawa@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Akaogi, Masaki
masaki.akaogi@gakushuin.ac.jp
Department of Chemistry, Gakushuin University
Mejiro, Tokyo 171-8588, Japan
Alves, Sophie
salves@pheasant.misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Asahara, Yuki
asahara@ganko.tohoku.ac.jp
Institute of Mineralogy, Petrology and Economic Geology, Tohoku University
Aoba-ku, Sendai 980-8578, Japan
Campbell, Ian
Ian.Campbell@anu.edu.au
Research School of Earth Sciences, Australian National University
Mills Road, Canberra ACT 200, Australia
Castillo, Paterno
pcastillo@ucsd.edu
Scripps Institution of Oceanography, University of California, San Diego
9500 Gilman Drive, La Jolla, CA 92093-0212, USA
Chiba, Hitoshi
hchiba@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Faul, Ulrichuli.faul@anu.edu.au
Research School of Earth Sciences, Australian National University
Canberra ACT 200, Australia
Fei, Yingwei
fei@gl.ciw.edu
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Fujino, Kiyoshi
fujino@ep.sci.hokudai.ac.jp
Division of Earth and Planetary Sciences, Hokkaido University
Kita 10 Nishi 8, Kita-ku, Sapporo 060-0810, Japan
Fukao, Yoshio
fukao@eri.u-tokyo.ac.jp
Earthquake Research Institute, University of Tokyo
Yayoi, Bunkyo, Tokyo, 113-0032, Japan
Green, Harry, II
harry.green@ucr.edu
Institute of Geophysics and Planetary Physics, University of California, Riverside
University Ave., Riverside, CA 92521, USA
Hanyu, Takeshi
hanyu@eri.u-tokyo.ac.jp
Earthquake Research Institute, University of Tokyo
Yayoi 1-1-1, Bunkyo-ku, Tokyo 113-0032, Japan
Hart, Stanley
shart@whoi.edu
Department of Geology, Woods Hole Oceanographic Institution
Woods Hole, MA 02543, USA
Hatakeyama, Tadahiro
hatake@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Hemley, Russell J.
hemley@gl.ciw.edu
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Higo, Yuji
higo@sci.ehime-u.ac.jp
Depertment of Earth Sciences, Ehime University
Bunkyo-cho 2-5, Matsuyama 790-8577, Japan172

October 2 – 5, 2001
Hirao, Naohisa

173

hirao@ganko.tohoku.ac.jp
Faculty of Science, Tohoku University
Aoba, Aramaki, Aoba-ku, Sendai 980-8578, Japan
Hirato, Takafumi
hrt1@geo.titech.ac.jp
Department of Earth and Planetary Sciences, Tokyo Institute of Technology
O-okayama 2-12-1, Meguro, Tokyo 152-8551, Japan
Honda, Satoru
satoru@yesterday.geol.sci.hiroshima-u.ac.jp
Department of Earth and Planetary Systems Science, Hiroshima University
1-3-1 Kagami-yama, Higashi Hiroshima 739-8526, Japan
Iguchi, Tomohiro
tiguchi@geo.titech.ac.jp
Department of Earth and Planetary Sciences, Tokyo Institute of Technology
2-12-1 Ookayama, Meguro-ku, Tokyo 152-8551, Japan
Inoue, Toru
inoue@sci.ehime-u.ac.jp
Geodynamics Research Center, Ehime University
2-5 Bunkyo-cho, Matsuyama, Ehime 790-8577, Japan
Irifune, Tetsuo
irifune@dpc.ehime-u.ac.jp
Geodynamics Research Center, Ehime University
2-5 Bunkyo-cho, Matsuyama, Ehime 790-8577, Japan
Ito, Eiji
eiito@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Iwamori, Hikaru
hikaru@eps.s.u-tokyo.ac.jp
Department of Earth and Planetary Sciences, University of Tokyo
Hongo, Tokyo 113-0033, Japan
Kagi, Hiroyuki
kagi@eqchem.s.u-tokyo.ac.jp
Laboratory for Earthquake Chemistry, Graduate School of Science, University of Tokyo
Hongo 7-3-1, Bunkyo-ku, Tokyo 113-0033, Japan
Kanzaki, Masami
mkanzaki@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Kato, Mamoru
mkato@gaia.h.kyoto-u.ac.jp
IHS/Earth Science, Kyoto University
Yoshida-Nihonmatsu, Sakyo, Kyoto 606-8501, Japan
Katsura, Tomootkatsura@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Kawakatsu, Hitoshi
hitosi@eri.u-tokyo.ac.jp
Earthquake Research Institute, University of Tokyo
1-1-1 Yayoi, Bunkyo-ku, Tokyo 113-0032, Japan
Kawamoto, Tatsuhikokawamoto@bep.vgs.kyoto-u.ac.jp
Institute for Geothermal Sciences, Kyoto University
Noguchibaru, Beppu 874-0903, Japan
Kobayashi, Katsura
katsura@pheasant.misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Koga, Kenneth T.
ktkoga@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Kondo, Tadashi
tdskondo@mail.cc.tohoku.ac.jp
Faculty of Science, Tohoku University
Aoba, Aramaki, Aoba-ku, Sendai 980-8578, Japan
Kono, Masaru
mkono@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan

Kurayoshi Symposium

174
Kubo, Atsushi

akubo@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Kuritani, Takeshi
kuritani@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Kusakabe, Minoru
kusakabe@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Li, Jiejieli@gl.ciw.edu
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Mao, Ho-kwang (Dave)mao@gl.ciw.edu
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Masters, Guy
gmasters@ucsd.edu
IGPP, Scripps Institution of Oceanography, University of California, San Diego
9500 Gilman Drive, La Jolla, CA 92093-0225, USA
Matsui, Masanorim.matsui@sci.himeji-tech.ac.jp
Faculty of Science, Himeji Institute of Technology
Kouto, Kamigoricho, Akogun, Hyogo 678-1297, Japan
Matsumoto, Naokomatsu@pheasant.misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Mayama, Norihito
mayama1@mail7.dddd.ne.jp
Department of Earth Science, Okayama University
Tsushimannishizaka2-14-30, Okayama 700-0086, Japan
Mibe, Kenjimibe@magma.eri.u-tokyo.ac.jp
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Miyajima, Nobuyoshi
Nobuyoshi.Miyajima@uni-bayreuth.deBayerisches Geoinstitut Universität Bayreuth
Universitätstr.30, Bayreuth D-95440, Germany
Moreira, Manuelmoreira@ipgp.jussieu.fr
Geochimie-Cosmochimie, Institut de Physique du Globe de Paris
4 place Jussieu, Paris 75005, France
Moriguti, Takuya
moriguti@misasa.okayama-u.ac.jp
PML, Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Morishita, Tomoaki
moripta@kenroku.kanazawa-u.ac.jp
Earth Sciences, Kanazawa University
Kakuma, Kanazawa, Ishikawa 920-1192, Japan
Moriyama, Takeru
takeru@pheasant.misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Mysen, Bjorn
mysen@gl.ciw.edu
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Nagai, Takayanagai@ess.sci.osaka-u.ac.jp
Department of Earth and Space Science, Osaka University
1-1 Machikaneyama, Toyonaka, Osaka 560-0043, Japan
Nakamura, Eizo
eizonak@misasa.okayama-u.ac.jp
PML, Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan

October 2 – 5, 2001

175

Nakayama, Keisukeknakayam@geo.titech.ac.jp
Earth and Planetary Sciences, Tokyo Institute of Technology
Ookayama 2-12-1, Meguroku, Tokyo 152-8551, Japan
Nishihara, Yu
yuu@geo.titech.ac.jp
Magma Factory, Earth and Planetary Sciences, Tokyo Institute of Technology
2-12-1 Ookayama, Meguro-ku, Tokyo 152-8551, Japan
Nishikawa, Osamu
nisikawa@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Obata, Masaaki
obata@kueps.kyoto-u.ac.jp
Department of Geology and Mineralogy, Kyoto University
Kita-shirakawa, Kyoto 606-8502, Japan
Ohtaka, Osamu
ohtaka@ess.sci.osaka-u.ac.jp
Earth and Space Science, Osaka University
Machikaneyama 1-1, Toyonaka, Osaka 560-0043, Japan
Ohtani, Eiji
ohtani@mail.cc.tohoku.ac.jp
Institute of Mineralogy, Petrology, and Economic Geology, Tohoku University
Aza-Aoba, Aramaki, Aoba-ku, Sendai 980-8578, Japan
Ono, Shigeaki
sono@jamstec.go.jp
Institute for Frontier Research on Earth Evolution, Japan Marine Science and Technology
Center
2-15 Natsushima-cho, Yokosuka, Kanagawa-ken 237-0061, Japan
Osako, Masahiro
sako@kahaku.go.jp
Division of astronomy and geophysics, National Science Museum
3-23-1 Hyakunin-cho, Shinjuku-ku, Tokyo 169-0073, Japan
Presnall, Deanpresnall@gl.ciw.edu
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Prewitt, Charles T.
prewitt@gl.ciw.edu
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Rapp, Robert P.
rapp@sbmp04.ess.sunysb.edu
Department of Geosciences and Mineral Physics Institute, State University of New YorkESS Building, Stony Brook, New York 11794-2100, USA
Ren, Zhong-Yuan
ren@geo.titech.ac.jp
Earth and Planetary Sciences, Tokyo Institute of Technology
2-12-1 Ookayama, Meguro-ku, Tokyo 152-8551, Japan
Richard, Guillaume
Guillaume.Richard@cnes.fr
UMR 5562, Terrestrial and Planetary dynamics Laboratory, Observatoire Midi-Pyrenees
14, Avenue E. Belin, Toulouse 31400, France
Rose, Estelle F.
erose@pheasant.misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Rubie, David C.
dave.rubie@uni-bayreuth.deBayerisches Geoinstitut, Universität Bayreuth
Universitätstr. 30, Bayreuth D-95440, Germany
Sakaguchi, Chiecsaka@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Sakamoto, Daisuke
dai@esstop.ess.sci.osaka-u.ac.jp
Earth and Space Science, Graduate School of Science, Osaka University
Machikaneyama-1, Toyonaka, Osaka 560-0043, Japan

176

Kurayoshi Symposium
Sanloup, Chrystele
sanloup@gl.ciw.edu
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Schmidt, Max W.
max.schmidt@erdw.ethz.ch
Institute for Mineralogy and Petrology, ETH
Sonneggstr. 10, Zürich 8092, Switzerland
Shirasaka, Mikisirasaka@geo.titech.ac.jp
Department of Earth and Planetary Scienses, Tokyo Institute of Technology
2-12-1 Ookayama, Meguro-ku, Tokyo 152-8551, Japan
Smyth, Joseph
joseph.smyth@colorado.edu
Department of Geological Sciences, University of Colorado
2200 Colorado Avenue, Boulder CO 80309-0399, USA
Song, Maoshungsong@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Stacey, Frank D.
f.stacey@cat.csiro.au
CSIRO, Exploration and Mining
Kenmore QLD 4069, Australia
Sueda, Yuichiro
sueday@sci.ehime-u.ac.jp
Department of Earth Sciences, Ehime University
Bunkyo-cho, Matsuyama, Ehime 790-8577, Japan
Takahashi, Eiichi
etakahas@geo.titech.ac.jp
Earth and Planetary Sciences, Tokyo Institute of Technology
2-12-1 Ookayama, Meguro-ku, Tokyo 152-8551, Japan
Takei, Hiroyuki
htakei@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Tanaka, Ryoji
ryoji@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Tange, Yoshinori
tan@issp.u-tokyo.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Tanimoto, Masaaki
tanimoto@pheasant.misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Tatsumi, Yoshiyuki
tatsumi@jamstec.go.jp
IFREE, JAMSTEC
Natsushima, Yokosuka, Kanagawa 237-0061, Japan
Terasaki, Hidenori
tera@luna.riko.tsukuba.ac.jp
Geoscience Institute, University of Tsukuba
Tennoudai 1-1-1, Tsukuba, Ibaraki 305-8571, Japan
Tsuchiya, Taku
takut@geo.titech.ac.jp
Earth and Planetary Science, Tokyo Institute of Technology
Ookayama 2-12-1, Meguro, Tokyo 152-8551, Japan
Tsujimura, Tomoyuki
tomoyuki@ganko.tohoku.ac.jp
Department of Mineralogy, Petrology and Economic Geology, Tohoku University
Aoba, SendaiMiyagi, 980-8578, JapanUrakawa, Satoru
afeg0320@cc.okayama-u.ac.jp
Department of Earth Science, Okayama University
3-1-1 Tsushima naka, Okayama700-8530, Japan
Usui, Tomohiro
tusui@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan

October 2 – 5, 2001

177

Van Orman, Jamesj.van-orman@gl.ciw.edu
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Van Westrenen, Willem
w.van-westrenen@gl.ciw.edu
Geophysical Laboratory, Carnegie Institution of Washington
5251 Broad Branch Rd., NW, Washington DC 20015, USA
Venkataramaiah, Ganesha Atni
ganmysore@hotmail.com
Department of Geology, Mysore University
Manasagangotri, Mysore, Karnataka 570 006, India
Walter, Michaelwalter@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Watada, Shingo
watada@eri.u-tokyo.ac.jp
Earthquake Research Institute, Tokyo University
1-1-1 Yayoi, Bunkyo-ku, Tokyo 113-0032, Japan
Wood, Bernard J.
b.j.wood@bris.ac.uk
Department of Earth Sciences, University of Bristol
Queens Road, Bristol, BS8 1RJ, U.K.
Xue, Xianyu
xianyu@dbmac1.misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Yagi, Takehikoyagi@issp.u-tokyo.ac.jp
Institute for Solid State Physics, University of Tokyo
5-1-5 Kashiwanoha, Kashiwa, Chiba 277-8581, Japan
Yamada, Hitoshi
hyamada@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Yamanaka, Takamitsu
b61400@center.osaka-u.ac.jp
Dept. Earth and Space Science, Osaka University
1-1 Machikaneyama, Toyonaka, Osaka 560-0043, Japan
Yamashita, Shigeru
shigeru@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Yamazaki, Daisuke
yamaz002@issp.u-tokyo.ac.jp
Institute for Solid State Physics, University of Tokyo
5-1-5 Kashiwanoha, Kashiwa, Chiba 277-8581, Japan
Yokoyama, Tetsuya
yokoyama@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Yoneda, Akira
yoneda@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Yoshino, Takashi
tyoshino@misasa.okayama-u.ac.jp
Institute for Study of the Earth’s Interior, Okayama University
827 Yamada, Misasa, Tottori 682-0193, Japan
Zhao, Dapeng
zhao@sci.ehime-u.ac.jp
Geodynamics Research Center, Ehime University
Bunkyo-cho 2-5, Matsuyama, Ehime 790-8577, Japan

Author Index
Kato, Mamoru . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5
Kato, T. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156
Katsura, Tomoo . . . . . . . . 61, 72, 77, 133, 150, 158
Kawakatsu, Hitoshi . . . . . . . . . . . . . . . . . . . . . . . . 5, 6
Kawamoto, T. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99
Kawamoto, Tatsuhiko . . . . . . . . . . . . . . . . . . . . . . 121
Kawamura, Katsuyuki . . . . . . . . . . . . . . . . . . . . . . 81
Kikegawa, Takumi . . . . . . . . . . . . . . . . . . . . . . . 79, 80
Kitakaze, A. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101
Kitamura, Masao . . . . . . . . . . . . . . . . . . . . . . . . . . . 24
Kittaka, Koichi . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78
Kobayashi, Katsura . . . 21, 22, 43, 46, 47, 90, 148,
169, 170
Koga, KT. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90
Kogiso, T. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13
Kondo, Tadashi . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80
Kono, Masaru . . . . . . . . . . . . . . . . . . . . . . . . . 160, 167
Kubo, Atsushi . . . . . . . . . . . . . . . . . . . . . . 61, 72, 150
Kudoh, Y. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101
Kuritani, Takeshi . . . . . . . . . . . 23, 43, 50, 141, 169

Aizawa, Yoshitaka . . . . . . . . . . . . . . . . . . . . . . . . . . 76
Akaogi, Masaki . . . . . . . . . . . . . . . . . . . . . . . . 63, 148Allègre, Claude J. . . . . . . . . . . . . . . . . . . . . . . . . . . . 57
Alves, Sophie . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57
Aoki, K. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104
Arai, Shoji . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 118
Asahara, Yuki . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147
Bebout, Gray E. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51
Bourdon, Bernard . . . . . . . . . . . . . . . . . . . . . . . . . . . 57
Campbell, Ian H.. . . . . . . . . . . . . . . . . . . . . . . . . . . . 15
Castillo, P.R.. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16
Chakraborty, S. . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88
Davies, Gareth R. . . . . . . . . . . . . . . . . . . . . . . . . . . . 17
Dunai, Tibor J. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17
Faul, Ulrich . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 142
Fortenfant, S. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154
Frost, Daniel J.. . . . . . . . . . . . . . 13, 88, 96, 126, 154
Fujihisa, H. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104
Fujino, Kiyoshi . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67
Fujita, Kiyoshi . . . . . . . . . . . . . . . . . . . . . . . . . . . . 133
Fukao, Yoshio . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3
Funakoshi, Ken-ichi . . . . . . . . . . . . . . . . . . . . . 71, 75
Funakoshi, Kenichi . . . . . . . . . . . . . . . . . 61, 72, 156

Langenhorst, Falko . . . . . . . . . . . . . . . . . . . . . . 88, 96
Laske, Gabi . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4
Makishima, Akio . 32, 35, 37, 39, 54, 55, 148, 169
Maruyama, Shigenori . . . . . . . . . . . . . . . . . . . . 30, 46
Masters, Guy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4
Matsui, Masanori . . . . . . . . . . . . . . . . . . . . . . . . . . . 66
Matsukage, Kyoko . . . . . . . . . . . . . . . . . . . . . 79, 121
Matsumoto, Naoko. . . . . . . . . . . . . . . . . . . . . . . . . . 24
Mayama, N. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77
Melosh, H.J. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154
Miyajima, Nobuyoshi . . . . . . . . . . . . . . . . . . . . 67, 96
Monnereau, M. . . . . . . . . . . . . . . . . . . . . . . . . . . . . 130
Moreira, Manuel . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18
Moriguti, Takuya . . . . . . . . . . . . . . . . 22, 25–27, 169
Morishita, Tomoaki . . . . . . . . . . . . . . . . . . . . . . . . 118
Murakami, Motohiko . . . . . . . . . . . . . . . . . . . . . . 132

Green, Harry W. . . . . . . . . . . . . . . . . . . . . . . . . . . . 114
Hanyu, Takeshi . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17
Hart, Stanley R.. . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11
Hasegawa, M. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156
Hasegawa, Masayuki . . . . . . . . . . . . . . . . . . . . . . . . 75
Hatakeyama, Tadahiro . . . . . . . . . . . . . . . . . . . . . 160
Helmstaedt, Herwart . . . . . . . . . . . . . . . . . . . . . . . . 46
Higo, Yuji . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71, 128
Hirao, Naohisa. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 80
Hirose, Kei . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132
Hirschmann, M. . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13
Holzapfel, Christian . . . . . . . . . . . . . . . . . . . . 88, 154
Honda, S. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163

Nagai, Takaya . . . . . . . . . . . . . . . . . . . . . . . . . . 78, 165
Naka, Hisanobu . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71
Nakajima, K. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 137
Nakamura, Eizo 21–23, 25–27, 29, 30, 32, 33, 35,
37, 39, 41, 43, 46, 47, 49–51, 53, 54, 55,
90, 137, 144, 145, 148, 169, 170
Nakanishi, Mayumi . . . . . . . . . . . . . . . . . . . . . . . . . 54
Nakano, Toshio. . . . . . . . . . . . . . . . . . . . . . . . . . 49, 53
Nakashima, Satoru . . . . . . . . . . . . . . . . . . . . . . . . . 132
Nakayama, Keisuke . . . . . . . . . . . . . . . . . . . . . . . . . 79
Nishihara, Yu . . . . . . . . . . . . . . . . . . . . . . . . . . 79, 103
Nishikawa, Osamu . . . . . . . . . . . . . . . . . . . 61, 72, 92

Iguchi, Tomohiro . . . . . . . . . . . . . . . . . . . . . . . 79, 103
Inoue, Toru . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71, 128
Irifune, Tetsuo . . . . . . . . . . . . . . . . . . . . . . . . . 71, 128
Ito, Eiji . . . . . . . . . . . . . . . . . . 61, 63, 69, 72, 83, 150
Iwamori, Hikaru . . . . . . . . . . . . . . . . . . . . . . . . . . . 117
Kagi, Hiroyuki . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 120
Kaneoka, Ichiro. . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17
Kanzaki, Masami . . . . . . . . . . . . . . . . . . . 93, 99–101

178

October 2 – 5, 2001
Nohda, Susumu . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17
Obata, Masaaki . . . . . . . . . . . . . . . . . . . . . . . . . . . . 146
Odawara, Reki . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67
Ogawa, Hisayuki . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67
Ohno, I. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77
Ohtaka, O. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95
Ohtani, Eiji . . . . . . . . . . . . . . . . . . . . . . . 80, 111, 147
Ono, S. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69
Osako, Masahiro . . . . . . . . . . . . . . . . . . . . . . . . . . . . 83
Ota, K. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95
Otaka, O. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 165
Poe, B.T. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154
Punongbayan, R.S. . . . . . . . . . . . . . . . . . . . . . . . . . 16
Rapp, Robert P. . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122
Reid, J.. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154
Ren, Zhang-Young . . . . . . . . . . . . . . . . . . . . . . . . . 137
Richard, G.C.M . . . . . . . . . . . . . . . . . . . . . . . . . . . 130
Righter, K. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 154
Rose, Estelle F.. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55
Rubie, Dave . . . . . . . . . . . . . . . . . . . . . . . . 88, 154, 96
Sakaguchi, Chie . . . . . . . . . . . . . . . . . . . . . . . . 29, 169
Sakamoto, D. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 95
Sakashita, M. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104
Sanehira, Takeshi . . . . . . . . . . . . . . . . . . . . . . . . . . . 71
Sasaki, Yohei . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67
Sata, Nagayoshi . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67
Sato, K. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156
Schiano, Pierre. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57
Schmidt, Max W. . . . . . . . . . . . . . . . . . . . . . . . . . . 109
Shimizu, Kenji . . . . . . . . . . . . . . . . . . . . . . . . . . 22, 30
Shimizu, N. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 122
Shinmei, Toru . . . . . . . . . . . . . . . . . . . . . . . . . . . 61, 72
Shinoda, Keishi . . . . . . . . . . . . . . . . . . . . . . . . . . . . 133
Shinozaki, K. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 137
Shirasaka, Miki . . . . . . . . . . . . . . . . . . . . . . . . 79, 124
Smyth, Joseph R. . . . . . . . . . . . . . . . . . . . . . . . . . . 126
Solidum, R.U. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16
Song, M. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104
Song, Maoshuang . . . . . . . . . . . . . . . . . . . . . . . . 61, 72
Stacey, Frank D. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7
Sueda, Yuichiro . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71
Suzuki, A. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156
Suzuki, I.. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77
Suzuki, Toshihiro . . . . . . . . . . . . . . . . . . . . . . . . . . 148
Takahashi, Eiichi . . . . . . . . . 79, 103, 124, 137, 145
Takazawa, Eiichi . . . . . . . . . . . . . . . . . . . . . . . . . . 146
Takeguchi, R. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 137
Takei, Hiroyuki . . . . . . . . . . . . . . . . . . . . . 32, 33, 169
Tamura, Akihiro . . . . . . . . . . . . . . . . . . . . . . . . . . . 118
Tanaka, Akira . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63

179
Tanaka, Ryoji . . . . . . . . . . . . 22, 137, 144, 145, 169
Tanimoto, Masaaki . . . . . . . . . . . . . . . . . . . . . . . . . . 50
Tanimoto, Yasutika . . . . . . . . . . . . . . . . . . . . . . . . 128
Tareen, J.A.K. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112
Terasaki, H. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 156
Tsuchiya, Taku . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 81
Tsujimura, T. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101
Ueda, Takayuki . . . . . . . . . . . . . . . . . . . . . . . . . 71, 128
Urakawa, Satoru . . . . . . . . . . . . . . . . . . . . . . . 75, 156
Usui, Tomohiro . . . . . . . . . . . . . . . . . . . . . 46, 47, 170
Uto, Kozo . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17
Utsumi, Wataru . . . . . . . . . . . . . . . . . . . . . . . . . . 71, 75
Venkatramiah, Ganesha Atni . . . . . . . . . . . . . . . 112
Walter, Michael J. . . . . . . . . . . . . . 90, 150, 152, 158
Wood, B.J. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 140
Xue, X. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 99–101
Yagi, Takehiko. . . . . . . . . . . . . . . . . . . . . . . . . . . 67, 96
Yamada, Hitoshi. . . . . . . . . . . . . . . . . . . . . . . . . 61, 72
Yamakawa, Junji . . . . . . . . . . . . . . . . . . . . . . . . . . . 75
Yamanaka, Takamitsu . . . . . . . . . . . . . . . 78, 95, 165
Yamashita, Shigeru . . . . . . . . . . . . . . . . . . . . . . . . 102
Yamawaki, H. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104
Yokoyama, Tetsuya . . . . 32, 35, 37, 39, 41, 43, 169
Yoneda, Akira . . . . . . . . . . . . . . . . . . . . . . . 77, 83, 86
Yoshiasa, A. . . . . . . . . . . . . . . . . . . . . . . . . . . . 95, 165
Yoshida, M. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163
Yoshikawa, Masako . . . . . . . . . . . . . . . . . . . . . . . . . 50
Yoshino, Takashi . . . . . . . . . . . . . . . . . . . . . . . . . . . 158
Yurimoto, Hisayoshi . . . . . . . . . . . . . . . . . . . . . . . 132
Zhao, Dapeng . . . . . . . . . . . . . . . . . . . . . . . . . . . 8, 116