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Doctoral Thesis

Lithium isotopes in the Earth and terrestrial planets

Author(s): Magna, Tomáš

Publication Date: 2006

Permanent Link: https://doi.org/10.3929/ethz-a-005383304

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Doctoral thesis ETH No. 16981

Lithium isotopes in the Earth and terrestrial

PLANETS

A dissertation submitted to the

Swiss federal Institute of Technology Zürich

for the degree of

Doctor of Natural Sciences

presented by

Tomas Magna

Mgr., Charles University in Prague, Czech Republic

born March 3, 1977

citizen of the Czech Republic

Accepted on the recommendation of

Prof Christoph Heinrich, ETH Zürich, examiner

Dr. Uwe Wiechert, Freie Universität Berlin, Germany, co-examiner

Prof. Alex N. Halliday, University of Oxford, UK, co-examiner

Prof. Timothy L. Grove, Massachusetts Institute of Technology, USA, co-examiner

Dr. Tim Elliott, University of Bristol, UK, co-examiner

Zürich 2006

Frontcover: A view of the Apollo 11 lunar module "Eagle" as it returned from the surface of

the Moon to dock with the command module "Columbia". A smooth mare area is visible on

the Moon below and a half-illuminated Earth hangs over the horizon. The picture was taken

by command module pilot Michael Collins just before docking on 21 July 1969.

Credit: NASA (http://www.nasa.gov)

Acknowledgement

This work would not be possible without an extreme effort of two prominent

personalities to whom I am more than grateful for their support, enthusiasm and criticism.

Alex Halliday was always well disposed and full of ideas and possible directions of the PhD

project. I also wish to thank him for giving me the opportunity to enter the great team of

geoscientists housed at ETH. Uwe Wiechert was the best choice I could ever get for a

supervisor. He pushed me further in my work and forced me to think about every single piece

of what I was doing, saying or writing. He always had time to talk, discuss and even spend his

time together with me in the lab during many long evenings; I really appreciate that. This I

found of much use for the future and I am indebted to Uwe for such an approach in

supervising my PhD. I hope he did not get mad after correcting first drafts of individual

chapters of the Thesis. In this place it is also appropriate to acknowledge Chris Ballentine

who fortunately for me, failed in trying to change my mind to noble gases.

In this place it is worth to thank Christoph Heinrich for overtaking the responsibility

after Alex Halliday which in turn lead to finishing the Thesis. Tim Elliott and Tim Grove arc

acknowledged for taking a part as co-referees and also for their help and advice that I was

often asking for at various stages of my work. I also thank Fin Stuart for his effort and

invaluable input during the work on Iceland and Jan Mayen manuscript. I am grateful to

Torsten Vennemann for language corrections of "Kurzfassung" and Francois Bussy, Denise

Bussien, Mike Cosca, Othmar Müntener and Torsten Vennemann for comments on first

versions of my presentation.

From the beginning to the end of my stay at ETH, Britt Meyer was helping me in

various things not always necessarily related to the academic and without her help I could

hardly live as easy as I used to. Also, Valentina Müller-Weckerle often provided help during

my stay at ETH.

I am indebted to Felix Oberli for introduction into the multi-collector mass spectrometry

and thorough and detailed explanations of every single aspect of the large Nul700 mass

spectrometer and for always a proper troubleshooting (including weekends). Martin Meier

provided much help in the lab with acid purification and material support. I am grateful to

Heiri Baur for time that he dedicated me on various occasions while explaining me different

3

and always complex things (primarily in noble gas lab but subsequently in other aspects too).

Machinery and electronics unit (Urs Menet, Andreas Süsli, Bruno Rutsche and Donat

Niederer) is a well-coordinated team that provided much help by manufacturing and/or

repairing plenty of necessary things. My lab work was in part possible due to effort of Marie-

Theres Bär, Irene Ivanov-Bucher and Ramon Aubert who provided invaluable help in making

columns, preparing samples and introducing me into the mineral separations, respectively.

While many people appeared on various occasions in my vicinity during the period of

my doctoral study at ETH, it would be a heroic effort to provide all names. Hence, I will pick

up individual people from this endless list. Darrell Harrison was a superb guy all the time and

helped me many times with variety of things and problems, easy or complicated (complicated

in most cases, of course). Rainer Wieler was paying attention to my frequent questions and

always tried to find a solution. He also introduced me into complex problematic of lunar

research. Special thanks go to Bettina Zimmermann for helping me with many things. Martin

Frank, Marcus Gutjahr and Uwe Wiechert shared the lab with me and as such, they provided

a lot of help or suggestions. The early ETH days in the noble gas lab were smoother due to

assistance of Veronika Heber and Nadia Vogel (and again, of course, Heiri Baur). The

productive team of PhD students and post-docs in IGMR was a fantastic group and I thank all

of them for sharing their time, contacts, ideas and other things without which it would not be

such a joy and pleasure to do a doctorate study at ETH Zürich.

Last, but not least and definitely the most important for me was the patience and ever¬

lasting support of my family. In particular, the last time was variegated by the birth of my

daughter Eliska in May 2005, the light of my life and sense of my being. This Thesis belongs

to Jifina and my family in a great part.

4

Contents

Abstract 7

Kurzfassung 9

Chapter 1: Introduction 11

1.1 Lithium isotopes in the global geochemical cycle 12

1.2 Physiochemical properties of lithium 13

1.3 Artificially isotopically enriched lithium 13

1.4 Objectives of this study 16

Chapter 2: Low-blank isotope ratio measurement of small samples of

lithium using multiple-collector ICPMS 17

2.1 Abstract 18

2.2 Introduction 18

2.3 Experimental 19

2.3.1 L-SVECpreparation 20

2.3.2 Sample preparation 20

2.3.3 Chromatography 21

2.3.4 Mass spectrometry 22

2.4 Results and discussion 23

2.4.1 Separation oflithium 23

2.4.2 Instrumental mass fractionation 26

2.4.3 L-SVEC measurements 30

2.4.4 International reference rocks 31

2.4.5 Other reference materials and rocks 33

2.5 Conclusions 35

Acknowledgement 35

Chapter 3: Lithium isotope fractionation in the southern Cascadia

subduction zone 37

3.1 Abstract 38

3.2 Introduction 38

3.3 Geology of the region 40

3.4 Analytical techniques 40

3.5 Results 42

3.6 Discussion 45

3.6.1 High-alumina olivine tholeiites 45

3.6.2 Basaltic andésites andprimitive magnesian andésites 47

3.6.3 Andésites and dacites 52

3.7 Li isotope fractionation in the southern Cascadia zone 54

5

3.8 Comparison with other arcs 57

3.9 Conclusions 60

Acknowledgement 61

Supplementary figure 62

Chapter 4: Combined Li and He isotopes in Tceland and Jan Mayenlavas and the nature of the North Atlantic mantle 63

4.1 Abstract 64

4.2 Introduction 65

4.3 Samples and geologic background 66

4.4 Analytical techniques 69

4.5 Results 70

4.6 Discussion 74

4.6.1 Lithium isotope composition of the North Atlantic mantle constrained fromIcelandic basalts andJan Mayen ankaramites 74

4.6.2 Li-He isotope correlation in Icelandic basalts 77

4.6.3 Implicationsfor the origin ofJan Mayen basalts from Li isotopes 82

4.7 Conclusions 85

Acknowledgement 86

Chapter 5: New constraints on the lithium isotope compositions of the

Moon and terrestrial planets 87

5.1 Abstract 88

5.2 Introduction 88

5.3 Methods and samples 90

5.4 Results 92

5.5 Discussion 97

5.5.1 Lunar Li and the isotope effects ofexposure °7

5.5.2 Lithium isotope composition ofthe Moon 101

5.5.3 Lithium isotope composition ofthe terrestrial mantle 106

5.5.4 Lithium isotope composition ofMars and Vesta 109

5.5.5 Lithium isotope composition ofchondrites Ill

5.6 Conclusions - inner solar system Li isotope compositions 1] 1

Acknowledgement 112

Chapter 6: Summary and outlook 113

6.1 Summary and outlook 114

References 119

Curriculum Vitae 137

6

Abstract

Lithium isotopes display very large relative mass difference between the two naturally

occurring isotopes, 6Li and 7Li, which reaches -17%. It is the largest difference among non¬

gaseous elements. Due to this, lithium is prone to extensive isotope fractionation in nature.

Special attention has been paid to surface processes and those in the hydrosphere, e.g.

weathering and seawater alteration, where this fractionation may reach values as large as

30%o. Pristine mantle rocks have been studied only very recently with the development of

improved analytical procedures and plasma source mass spectrometry techniques. However,

the mantle represents the largest portion of the Earth and thus, it should have a well-defined

Li isotope composition. This dissertation was aimed at partly filling the gap resulting from a

lack of Li isotope data for mantle rocks together with volcanic rocks from subduction zones

and ocean island basalts. Additional new data were obtained for a variety of meteorites and

lunar rocks.

A new rapid method of chemical isolation and purification of Li has been developed in

this dissertation for precise and accurate determination of Li isotopes. This newly developed

separation technique was combined with measurements of Li isotopes on a large-sector MC-

ICPMS. Mass spectrometry measurements involved sample bracketing with standards to

correct for instrumental mass bias. Long-term reproducibility of isotope measurements (±

0.3%o, 2a) has been significantly improved by a factor of 2-3 compared to the classical TIMS

methods. Li isotope homogeneity of the international Li standard material has been verified.

A number of international reference rocks have been characterised in terms of Li isotopes

that could further improve an inter-laboratory comparison.

Primitive olivine tholeiites and basaltic andésites from the southern Cascadia

subduction zone show a range of S7Li values that correlate inversely with distance from the

arc trench. Olivine tholeiites provide evidence for heterogeneous mantle beneath Mt. Shasta

and Medicine Lake. On the contrary, basaltic andésites and primitive magnesian andésites

show remarkably diverse Li isotope patterns below the two studied volcanoes. This has been

attributed to distinct values of S7Li of the fluids involved in the generation of basaltic

andésites, with fluids depleted in 7Li interacting with Medicine Lake lavas. This likely

7

reflects continuous Li isotope fractionation during dehydration of the subducted slab with

preferential removal of Li into the fluid phase at shallow depths.

Suites of picritic basalts from Hengill, Iceland, and ankaramitic basalts from Jan Mayen

support the view of a homogeneous upper mantle (ö7Li - +4%o), irrespective of the degree of

trace element enrichment. Inverse and linear co-variation of Li and He isotopes in Icelandic

lavas is not reflected in other trace element or isotope ratios. It is suggested that three mantle

components provide an explanation for the observed trends, contrary to previously proposed

implementation of altered Icelandic crust. At least one of these components carries a recycled

component as evidenced by positive Eu anomaly and super-chondritic Sr/Nd. Jan Mayen

lavas seem to accommodate a signature of enriched mantle (EM2).

Samples from terrestrial planetary bodies (chondrites, eucrites, Martian meteorites,

lunar basalts and terrestrial mantle peridotites) have significant variations in 57Li values with

predominantly light Li in chondrites. All other planetary objects display nearly identical Li

isotope compositions indistinguishable from the assumed terrestrial upper mantle at ô Li of

about +4%o. Lunar rocks span a substantial range of 6%o with extremely high 87Li for the

lunar crust which is likely caused by equilibrium isotope fractionation during plagioclase

segregation from the early lunar magma ocean. Variations of 8 Li among lunar basalts and

glasses can be explained by melting of early- and late-stage cumulates with variable amounts

of major minerals with isotopically different Li. Lack of Li isotope fractionation among inner

solar system planetary bodies relative to lighter Li in chondrites could provide evidence for

small but significant isotope fractionation during planetary formation processes.

8

Kurzfassung

Lithiumisotopen zeigen sehr grosse relative Massendifferenz zwischen den zwei

natürlich vorkommenden Isotopen, f'Li und 7Li, die bis -17% reicht. Es ist die grösste

Differenz zwischen Elementen, die nicht als Gase vorkommen. Als Konsequenz ist Lithium

anfällig zur extensiven Isotopenfraktionierung in der Natur. Besondere Aufmerksamkeit

wurde den Oberflächeprozessen und Prozessen in der Hydrosphäre gewidmet, z.B.

Verwitterung und Alteration in der Gegenwart von Meerwasser, wobei diese Fraktionierung

bis zu 30%o reichen kann. Die ursprünglichen Mantelgesteine wurden mittels einer neu

entwickelten und verbesserten analytischen und massenspektrometischen Methode

untersucht. Da der Mantel das grösste Reservoir für Li der Erde darstellt, stand eine

Charakterisierung seiner Li-Isotopie im Vordergrund dieser neuen Messungen. Um somit ein

besseres Verständnis über die Li-Isotopie des Mantels zu bekommen, wurden insbesondere

Lithiumisotopendaten für Mantelgcsteinc, vulkanische Gesteine aus Subduktionszonen und

ozeanische Inselbasalte ermittelt. Zusätzlich, wurden Daten für verschiedene Meteoriten und

Mond-Gesteine ermittelt.

Eine neue und schnelle Methode für die chemische Trennung und Aufbereitung des

Lithiums ist als Teil dieser Dissertation entwickelt worden, um eine genauere und präzisere

Bestimmung der Li-Isotopie als bisher zu ermöglichen. Nach der neuen

Aufbereitungsmethode wurden die Messungen der Li-Isotopie an einem MC-1CPMS mit

hoher Auflösung mittels einem gezielten Standard-Probe-Standard Messverfahren analysiert,

um den instrumentalen Massenbias deutlich zu verringern. Die langfristige

Reproduzierbarkeit der Isotopenmessungen (+ 0.3%o, 2a) ist zwei- bis dreifach verbessert

worden, im vergleich zu den klassischen TIMS Methoden. Isotopenhomogenität des

internationalen Lithiumstandards ist nachgewiesen worden. Eine Reihe von internationalen

Referenzgesteinen wurde ebenfalls bezüglich der Lithiumisotopen charakterisiert, was einen

Vergleich der Daten zwischen verschiedenen Labors vereinfachen soll.

Olivin-führende Tholeiite und Basalt-Andcsite aus der Süd-Cascadia Subduktionszone

zeigen grosse Variationen von ö7Li, die negativ mit dem Abstand vom Graben korrelieren.

Die Olivin-Tholeiite weisen auf einen heterogenen Mantel unter Mt. Shasta und Medicine

Lake. Im Kontrast dazu, haben die Basalt-Andesite und Magnesium-Andesite

9

bemerkenswerterweise sehr unterschiedliche Profile der Lithiumisotopen unter den zwei

Vulkanen. Diese Unterschiede können verschiedenen 57Li Werten der Fluide die sich an der

Generation der Basalt-Andesite beteiligt haben, zugerechnet werden. Vorwiegend leichte

Lithiumisotopie der Fluide wurden für Proben von Medicine Lake gemessen. Vermutlich

reflektiert das eine kontinuierliche Lithiumisotopenfraktionierung während Dehydration der

subduzierten Platte zusammen mit bevorzugter Abgabe von 7Li in die fluide Phase bei

niedrigen Subduktionstiefen.

Proben der pikritischen Basalte von Hengill, Island, und den Basalten von Jan Mayen,

unterstützen einen homogenen oberen Mantel (ô7Li ~ +4%o), unabhängig vom Grad der

Spurenelementenanreicherung. Die inverse und lineare Korrelation der Lithium- und

Heliumisotopen in den Lavas von Island, wird nicht von anderen Spurenelementen- oder

Isotopenverhältnissen widerspiegelt. Es ist wahrscheinlich, dass drei verschiedene

Mantclkomponenten die beobachteten Verhältnisse erklären können. Mindestens eine dieser

Komponenten beinhaltet rezykliertes Material, welches durch eine positive

Europiumanomalie und über-chondritischen Sr/Nd belegt wird. Basalte von Jan Mayen

deuten vermutlich auf die Komponente eines angereicherten Mantels (EM2).

Die Meteoritproben (Chondrite, Eukrite, von Mars stammende Meteorite, Mond-Basalte

und terrestrische Mantelperidotite) zeigen massgebliche Variationen in 57Li Werten, wobei

die Chondrite vorwiegend leichte Lithiumisotopenverhältnisse haben. Alle anderen Proben

zeigen fast identische Lithiumisotopenverhältnisse, identisch mit denen die für den

terrestrischen oberen Mantel angenommen werden (8 Li ~ 14%o). Mond-Gesteine erstrecken

sich wesentlich über 6%o mit extrem schweren Werten für die Kruste des Mondes, die

voraussichtlich durch Gleichgewichtsisotopenfraktionierung während einer Trennung des

Plagioklas vom frühen Magmaozean des Mondes verursacht ist. Unterschiede zwischen

Mond-Basalten und Gläsern sind wahrscheinlich auf Schmelzprozesse mit unterschiedlichen

Mengen von Früh- und Spätstadiumkumulaten der Hauptminerale zurück zu führen, wobei

diese Minerale unterschiedliche Lithiumisotopenverhältnisse haben. Abwesenheit einer

Isotopenfraktionicrung von Lithium zwischen den inneren Planeten im Vergleich zu einem

geringen Unterschied dieser zu den Chondriten, deutet auf geringe aber doch messbare

Isotopenfraktionierung während der Bildung der Planeten.

10

Chapter 1 Introduction

Chapter 1

Introduction

11

Chapter 1 Introdut. tion

1.1 Lithium isotopes in the global geochemical cycle

This PhD thesis is dedicated to the investigation of lithium (Li) isotopes in volcanic

arcs, ocean islands, mantle rocks and samples from the moon and other planetary bodies. The

large mass difference of-16% between 6Li and 7Li causes large isotope fractionation in

global geochemical cycles and makes Li isotopes an interesting geochemical tracer (e.g.

(ELLIOTT et al., 2004; Tomascak, 2004). The total variation of S7Li among terrestrial

materials is as large as 70%o (see Fig. 1.1). Rivers carry heavy Li to the oceans (Huh et al.,

1998, 2001; ViGiER et al., 2002; Pistiner and Henderson, 2003; Pogge von Strandmann

et al., 2006) and residual weathered products may develop extremely light 57Li (Rudnick et

al., 2004). The upper continental crust tends to have low ô Li compared to mantle and

averages at about ö7Li ~ 0%o (Teng et al., 2004) although our current knowledge of Li in the

continental crust is based on only two systematic studies (BRYANT et al., 2004; TENG et al.,

2004) and sparse measurements of single granitic rocks (JAMES and PALMER, 2000;

TOMASCAK et al., 2003; RUDNICK et al., 2004). Considerable diversity in the Li isotope

compositions of marine sediments has been illustrated recently (Chan et al., 2006). The ö7Li

of global marine sediments varies widely (-4 to +15%o) and reflects lithology, provenance and

diagenetic processes. There is no relation between distribution of Li isotopes and arc-trench

location. Marine sediments yield a range of 57Li (Chan et al., 2006) that closely resembles

that reported for altered oceanic crust (Chan et al., 1992, 2002a; Bouman et al., 2004). This

complicates a direct fingerprinting of the slab components in arc volcanics.

The most puzzling Li isotope results to date are perhaps related to subduction zone

volcanism. It has been shown that Li abundances correlate with fluid signatures (e.g.

DoRfc-NDORF et al., 2000) but Li isotope ratios are not correlated (Chapter 3). At least in two

volcanic arcs Li isotopes change systematically with increasing distance form the arc trench

(see Chapter 3 and Moriguti and Nakamura, 1998b). This is clear evidence that Li isotopes

fractionate during subduction of oceanic lithosphère into the mantle. Experimental studies

provide evidence that Li may be efficiently retained in the slab by several mineral species,

e.g. staurolite, serpentine, phengite or Ti-clinohumite (Woodland et al., 2002; Zack et al.,

2003; SCAMBELLURI et al., 2004; Wunder et al., 2006) but the exact mechanism and

response of Li isotopes to progressively changing pressures and temperatures, i.e. mineralogy,

is poorly constrained (MARSCHALL et al., 2006, 2007). It is further unclear whether

dehydration and metamorphism of subducted oceanic lithosphère generates distinct mantle

12

Chapter I Introduction

reservoirs because eclogite samples seem to preferentially carry isolopically light Li to the

lower mantle. Lithium isotope compositions from ferropericlase inclusions are light compared

to oceanic basalts and upper mantle rocks (Seitz et al., 2006a). This suggests that isotopically

light Li is recycled into the deep mantle mainly by eclogites and/or dehydrated serpentinites.

This is, however, not supported by data from ocean island basalts, which are predominantly

composed of isotopically heavy Li.

1.2 Physiochemical properties of lithium

Lithium has its origin in the Greek name for stone: Xrcr]oa (lithos). It was discovered

by J.A. Arfvedson in 1817 (ten years after potassium and sodium were isolated by Sir

Humphry Davy) and isolated as a metal using electrolysis one year later by Sir Humphry

Davy. Lithium is the lightest metal of the alkali group and the lightest of all solids of the

periodic table with a density of just 0.543 g.cm" .Its atomic weight is 6.941 + 0.002. In

metallic form Li forms a cubic lattice. It has the highest melting and boiling temperature

(180.5 and 1342°C, respectively) of all alkali metals. The first ionisation energy of Li (5.39

eV) is also higher compared with other elements of this group. Some chemical and physical

properties of Li are caused by the small ionic radius of the Li'

ion (~76 pm) which is similar

to that of Mg+

(72 pm) allowing Li to substitute for Mg. The two naturally existing isotopes

of lithium, Li and Li, were unambiguously distinguished in the early '20s (DEMPSTER,

1921) and further investigated by Aston (1927, 1932) who refined the abundances and

atomic weight of Li.

Lithium is a lithophile element. Although in the mantle Li is only slightly incompatible

(Ryan and LANGMUIR, 1987), in granitic rocks it can accumulate during the late stages of

magmatic differentiation. Pegmatites have generally high Li abundances in percent levels

(e.g. Fabre et al., 2002; Teng et al., 2006b) in which Li often forms its own minerals like

spodumene or petalite.

1.3 Artificially isotopically enriched lithium

Lithium is an element of interest also in fields outside geochemistry, e.g. medical

sciences (Schou, 1988), astrophysics and astronomy or the nuclear industry (Olive and

13


"T t vt-i r-T r-r | t

i seawatör

rivers

marine

carbons

1 lakes ^^^H

hinn*»nii?^^^^

îtes^^^^ii

É^H shales & loess

^^^iï-. pendotite S ovroxetute

Ml arc lavas

OIB

Hi altered MORS

fresh MORB

< •

chondrites

( i i i i i i i i i . r . i . , i i . , ,

-20 -10 0 10 20 30 40 50

Ô7LiL-SVEC (%o)

Fig. 1.1: Lithium isotope compositions of various terrestrial materials and chondrites (after Tomascak, 2004)

Light grey vertical line represents reference material (L-SVEC: SRM8545), dark grey vertical line represents Li

isotope composition ofpristine, non-metasomatic upper mantle (à Li = 4.1 ± 0.4%o; see Chapter 5).

Schramm, 1992; Ritzenhoff et al., 1997; Suzuki et al., 2000; Montalban and Rebolo,

2002). The nuclear industry has been using f'Li as a moderator because of the large thermal

neutron cross-section. The nuclear industry, therefore, processed large quantities of natural Li

to produce 6Li-enriched lithium. The residual Li is extremely enriched in 7Li. These 7Li-rich

materials are commercially distributed and eventually re-occur in e.g. reference solutions

and/or as impurities in some chemical reagents (Ql et al., 1997). This can lead to severe

analytical problems for concentration measurements, e.g. when mass spectrometry is used for

trace element studies. Another problem rises from métallurgie use of L12B4O7 to decrease

melting temperatures of substances, e.g. for XRF analysis of geological samples. When these

reagents are made from industrially processed Li, even minute quantities of Li2B4C»7, e.g.

contamination by dust particles, may dramatically alter Li isotope ratios of natural samples.

This is illustrated in Fig. 1.2 for several Cameroon Line basalts presumably contaminated by

Li2B4C>7 powder and discussed next in more detail.

14

Chapter I Introduction

1000

~ 100o

_l

10

0

a7Li = 4.90±1.35%o

10 15

Li (ppm)

20

lc

tsEEta

IS

25

Fig I 2 Cameroon Line samples partly contaminated with industrially processed Li fhe most extreme à Li

value of 490.3%o has been obtained for basalt AN7 Three other samples display oLi values ^ISSXo. Some

samples, however, pose typical values, i e within the range offresh oceanic basalts Note that Li abundances in

contaminated samples are increased by less than a factor of 5 compared to basalts showing no clear signature

of artificially indue ed Li Dashed line average oj non-contaminated samples, dotted lines 2SD Note the

logarithmic scale for ifLi

Common terrestrial basaltic rocks have 57Li values between ~+l%o and ~+6%o. These

comprise fresh mid-ocean ridge basalts (MORBs), ocean island basalts (OIBs) and island arc

basalts (IABs). Deviations are ascribed to secondary processes, e.g. weathering or interactions

with seawater. However, these changes do not exceed a few percent; the maximum ô Li

reaches ~+44%0 as recorded in pore fluids of marine sediments (ZHANG et al., 1998).

Moreover, seafloor alteration at ~2°C with an apparent fractionation factor a ~ 0.981 (Chan

et al., 1992) generates altered basalts with ô7Li of about +14%o which cannot explain the

observed Li enrichment in some of the Cameroon Line basalts. One possibility of producing

the extremely light Li isotope compositions could be a fractionation on ion-exchange columns

(MORIGUTi and Nakamura, 1998a; KOSLER et al., 2001) but the recovery of Li in Cameroon

Line basalts reached 100% within error. Therefore, analytical artifacts related to column

chemistry are not a viable explanation for high 87Li values for some Cameroon Line basalts.

More likely it is caused by contamination with artificially Li-enriched material picked up at

Edinburgh University where the samples were prepared for XRF analysis.

15


1.4 Objectives of this study

A major objective of this study was to establish a new rapid technique to separate Li

from silicate matrices for precise Li isotope measurements using MC-ICPMS. The analytical

method, mass spectrometry and data evaluation are described in Chapter 2. Another step

became necessary to separate Cr from Li when clinopyroxenes or fertile mantle xenoliths

were studied (see Chapter 5). This latter step of analytical treatment has been applied to

samples from Jan Mayen (Chapter 4) and to lunar rocks and meteorites (Chapter 5).

The following investigations were performed in this study:

a) Lithium isotope measurements of a suite of high-alumina olivine tholeiites and basaltic

andésites from Mt. Shasta and Medicine Lake volcanoes (northern California Cascades)

have been performed to study the geochemical behaviour of Li in subduction zones.

Similar studies were performed at other arc locations. However, the southern Cascades

are distinct as they consist mainly of basaltic lavas. We provide evidence for isotope

fractionation of Li during subduction. (Chapter 3).

b) Lithium isotopes have been used to identify recycled components in the source of

Icelandic lavas from Hengill and Jan Mayen Island lavas. For Hengill lavas a trend

between Li and He isotopes has been found which most likely reflects a minimum of three

distinct components in the plume head and adjacent mantle. Jan Mayen ankaramitic

basalts have a homogeneous isotope composition of 57Li = 4.2 ± 0.6%o within error

identical with that of the bulk silicate earth (Chapter 4).

c) Mineral separates from primitive unmetasomatised peridotites were investigated to

constrain the Li isotope composition of the terrestrial upper mantle. Data for lunar rocks

have revealed systematic Li isotope variations related to the crystallisation of the lunar

magma ocean. The Li isotope compositions of other terrestrial planetary bodies of the

inner solar system have been constrained (Chapter 5).

16

Chapter 2 Analytical procedures

Chapter 2

Low-blank isotope ratio measurement of small

samples of lithium using multiple-collector

ICPMS

published as: Magna T., Wiechert U.H. and Halliday A.N. (2004) Low-blank isotope ratio

measurement of small samples of lithium using multiple-collector ICPMS. International

Journal ofMass Spectrometry 239 (1), 67-76

17

Chapter 2 A nalytical procedure s

2.1 Abstract

A new method is presented for separation of lithium from silicate rocks and high

precision MC-ICPMS analysis. A relatively small (3.57 meq) resin volume is able to separate

lithium from all silicate rocks in a single step using only 16 ml of nitric acid mixed with

methanol. Some advantages of the method are high sample throughput, low blanks and

elution parameters that are insensitive to lithology. Elution schemes are presented for a range

of igneous rocks and minerals. ô7Li values for standards from the U.S. Geological Survey

(USGS) for BCR-1, BHVO-2 and AGV-2 are 2.4%o, 4.6%0 and 7.9%c, respectively. Reference

materials from the Geological Survey of Japan (GSJ) JB-2 and JR-2 give 57Li values of 4.7%e

and 3.8%c, respectively. The data for most reference rocks reproduce to within better than

0.5%o. This column method can also be used for the direct separation of high-field-strength

elements (Ti, Zr, Hf, Nb, Ta) from silicate rocks.

2.2 Introduction

Lithium (Li) isotope geochemistry has developed rapidly in recent years. It was first

used effectively following the development of a borate technique for measurement by thermal

ionization mass spectrometry (TIMS; CHAN, 1987). However, it has been further enhanced by

the advent of multiple-collector inductively coupled plasma mass spectrometry (MC-ICPMS;

Tomascak et al., 1999a). The ionic radius of Li+ (~ 0.59Â) is similar to that of Mg2+ (~

0.57A). Therefore, lithium can substitute for magnesium in olivine, enstatite and diopside

(Shitz and Woodland, 2000). This substitution behaviour contrasts with that of the large

alkali ions (K, Rb and Cs) and means that Li behaves like a moderately incompatible element

during partial melting of mantle rocks. In aqueous solutions lithium is strongly hydrated.

These chemical properties make Li and its isotopes interesting for the study of hydrothermal

processes (Chan et al., 1992), continental weathering rates (Zhang et al„ 1998; Huh et al.,

2001), and as a tracer of subducted oceanic crust in the mantle (Moriguti and Nakamura,

1998b; Tomascak et al., 1999b).

Lithium has two stable isotopes, 6Li and 7Li, with a large relative mass difference of

-17%. Mass-dependent lithium isotope fractionation has been known since Taylor and

Urey (1938) observed isotope fractionation of 25% as they percolated a lithium solution

through a zeolite column. Early attempts to measure the isotope composition of lithium in

18


geological materials were inconclusive and not precise enough to resolve any isotope

variations (SMALES and Webstkr, 1958). The first accurate and precise measurements of

lithium isotopes were performed using TIMS (Chan et al., 1993; 2002a; Sahoo and

Masuda, 1995; You and Chan, 1996; Moriguti and Nakamura, 1998b; Chan and

Kastner, 2000; Jamks and Palmkr, 2000; Chan and Frky, 2003). However the TIMS

technique suffers from a highly instable instrumental fractionation and requires Li that is

virtually free of matrix (You and Chan, 1996; Bicklh et al„ 2000). Therefore, complicated

ion exchange procedures have to be applied to obtain very pure Li solutions (MORIGUTI and

NAKAMURA, 1998a). Quadrupole ICPMS is able to measure isotope ratios on very small

amounts of lithium but is limited to lower precision compared with MC-ICPMS (KOSLER et

al„ 2001; 2003). Other techniques that have been used for the measurement of lithium

isotopes include atomic absorption spectrometry (Chapman and Dalk, 1976; Mkikr, 1982),

secondary ionization mass spectrometry (Chaussidon and Robert, 1998) and laser-excited

atomic fluorescence spectroscopy (Smith et al., 1998). The advent of multiple-collector

ICPMS offers the opportunity for small amounts of lithium to be analysed to high precision

(Tomascak and Tera, 1997; Tomascak et al., 1999a; Nishio and Nakai, 2002).

Most recently the procedures that have been developed to separate lithium from silicate

rocks use large cluant volumes (TOMASCAK et al., 1999a; NlSHIO and NAKAI, 2002;

Jkifcoate et al., 2004) and, therefore, are limited to relatively large sample amounts and are

sensitive to (Mg, Fe)-rich rock matrices (Chan et al., 1999, 2002b). Here we describe a rapid

and matrix-independent method for separation of lithium from very small silicate rock

samples. The capability of measuring Li isotopes using MC-ICPMS, with a high degree of

accuracy and precision is demonstrated with data for international reference rock standards.

2.3 Experimental

The natural abundances of 6Li and 7Li are 7.5% and 92.5%, respectively. That the

heavier isotope Li is more abundant than the lighter Li, has led some authors to write 8-

values as 5 Li. In this way positive Ô-valucs arc isotopically light and negative Ô-valucs arc

isotopically heavy. This is the opposite of how 6-notation is used for stable isotopes of

carbon, nitrogen, oxygen, sulphur and many other new isotope systems. We feel therefore it is

advisable to report data in the standard 5 Li-notation in agreement with the recommendations

(Coplen et al., 2002):

19


ö7Li(%0) =(7Li/6Li),

(7Li/6Li),.'sample

SVEC

x1000.

The lithium isotope composition is given in per mil relative to NIST SRM 8545 or L-SVEC

(7Li/6Li=12.02 ± 0.03; Flesch et al., 1973). Other standards are utilized sometimes, e.g.

seawater (Hokks and Sywali., 1997; Miu.ot et al., 2004) or IRM-016 (Ol et al., 1997). One

advantage of seawater as a standard for the lithium isotope scale is that seawater is a large

homogeneous reservoir. However, using seawater as a reference requires an additional

calibration step. Therefore, it is a potential source of errors and might cause additional

problems when data from different laboratories arc compared.

2.3.1 L-SVECpreparation

Approximately 100 mg of L-SVEC powder (L12CO3) from USGS was dissolved gently

in 2 ml of concentrated HNO3, then evaporated to dryness and taken up in 500 ml of 2%

HNO3. This produces a 40-ppm stock solution that can be further diluted to Li concentrations

suitable for high sensitive mass spectrometer analyses.

Galy et al. (2003) have reported isotopic heterogeneity in the magnesium international

reference material SRM 980. Therefore, comparisons between aliquots were made to confirm

isotopic homogeneity of L-SVEC standard material. For this purpose 50-ppb solutions of L-

SVEC aliquots were prepared from three international labs.

2.3.2 Sample preparation

All acids were twice distilled to reduce Li, Na and B blanks, methanol was distilled

once; this removes sodium from the acids and methanol which may have concentrations high

enough to cause matrix effects, i.e. shifts in isotopic ratios of samples relative to L-SVEC.

Indeed, the Li blank is also particularly crucial. Ql et al. (1997) reported highly anomalous

lithium isotope compositions (> I000%o) in some chemical reagents.

Approximately 100 mg of each rock powder was digested in Savillex® screw-top

beakers with 3 ml concentrated HF and 0.5 ml concentrated HNO3 for 24 hours on a hot plate

at ~130°C. After evaporation of the solution, the residue was re-dissolved and dried down

three times in concentrated HNO3 to remove fluorides completely. Finally, the sample was

dissolved and stored in 8 ml 6M HCl. An aliquot ('/sth of the solution, equivalent to -12.5 mg

of the original sample powder) from the 6M HCl solution was then evaporated and dissolved

20

Chapter 2 A nalytkal procedure s

in 0.5 ml of 0.67M HNO3 in methanol. This is made from a 30% v/v solution of methanol,

which dilutes concentrated HNO3 to the desired concentration and is hereafter referred to as

0.67M HN03/methanol. The final concentration of methanol in the solution is -27%. This

aliquot was used for Li separation by ion exchange chromatography.

2.3.3 Chromatography

The separation of lithium from other elements using mixtures of mineral acids and

organic media is based on the previous work of Sulcek and co-workers (SULCEK et al., 1965;

âuLCEK and RUBESKA, 1969). They established the technique employing a mixture of HCl

and methanol for separation of lithium from silicate rock matrices. The advantage of an HC1-

methanol mixture compared to pure HCl is the better separation of lithium from sodium, i.e.

methanol increases the difference of the partition coefficients between lithium and the other

alkali metals on cation resins. The extraction of lithium was further improved by using HNO3;

HCl may create Fe(III)-complexes that are not retained by cation resins and may therefore

clute Fe together with lithium (Strklow et al., 1974). Moreover, HNO3 increases the

distribution coefficient of Na but the behaviour of lithium is unchanged, i.e. Na is even more

strongly bound to the resin than with HCl. Such a methanol-nitric acid mixture has previously

been used by a number of research groups. For MC-ICPMS it was first used by Tomascak et

al. (1999a). Kosler et al. (2001) optimised the method for separation of small amounts of

lithium from carbonate rocks and shells of planktonic foraminifera. However, successful

separation of lithium from silicate rocks has been hampered by high eluant volumes and

matrix effects. The former may potentially cause high blanks that limited early studies of

lithium isotopes to samples with relatively high Li concentrations (e.g. basalts). Only in a

more recent study have these problems been overcome successfully (Jeffcoate et al., 2004)

although still large eluant volumes are used.

For this study Teflon columns with a volume of 2.1 ml were packed with BioRad® AG

50W-X8 (200-400 mesh) cation exchange resin. The columns are conical in shape with 6-mm

upper and 5-mm lower internal diameter. The resin is cleaned before the first use by repeated

rinsing with 6M HCl and de-ionized water, then with 3M HCl and de-ionized water. When

the columns are not used, they are stored in weakly acidified de-ionized water. Prior to

loading of a sample, the column is rinsed with 4 ml HCl and 4 ml de-ioni/>ed water and pre¬

conditioned with 2 ml 0.67M HNOi/methanol. The sample is loaded onto the resin in 0.5 ml

21

Chapter 2 Analytical pnx edures

of 0.67M HN03/methanol. Subsequently cations are eluted with IM HNO3 in methanol. This

is made from an 80% v/v solution of methanol, which dilutes concentrated HNO3 to the

desired concentration and is hereafter referred to as IM HNO^/mcthanol. The final

concentration of methanol is -15%. The eluate is then evaporated to dryness. Finally, samples

are dissolved in 2% HNO3 for analysis on the mass spectrometer. The time for passing 1 ml

through the columns is approximately 15 min. The procedural blank is usually less than 20 pg

Li.

The total capacity of the resin in 2.1-ml column is 3.57 meq (1.7 meq/ml). For 12.5 mg

of BHVO-2 the total capacity of major elements (total iron as Fe3+) is 0.283 meq. Silicon is

not included because it is fumed out by decomposition of the sample. Thus, the capacity of

the elements represents less than 10% of the total exchange capacity of the resin, which is in

the range for optimal separation of cations on this type of resin.

2.3.4 Mass spectrometry

Lithium isotope ratios were measured on Nu 1700 (Nu Instruments), a new large-

geometry high-resolution MC-ICPMS (HALLIDAY et al., 2002). The Nu 1700 is equipped with

3 electron multipliers and 16 Faraday cups with the outermost Faraday detectors L7 and H8

covering a mass dispersion of approximately 14% by neutral quadrupole setting conditions.

All Faraday cups are equipped with 10nQ resistors. We utilized the two outermost movable

Faraday cups L7 and H8 for the fiLi and 7Li beam, respectively. Sample solutions were

introduced through either desolvating spray nebulizer DSN-100 (Nu Instruments) or Aridus

Table 2.1 : Instrumental settings and conditions

RE power (W) 1400

Nebulizer pressure (psia) 49

Auxiliary gas flow rate (L/min) 0.9

Coolant gas flow rate (L/min) 13.0

Nebulizer micro-concentric

Spray chamber temperature (°C) 108

Membrane temperature (°C) 108

Sample uptake rate (uL/min) 80

Accelerator voltage (kV) 6

Resolution -700

Analyzer pressure (mbar) 2.3 x 10"9

Conditions for DSN-100, only argon is useda

1 psi = 0.06895 kg.cnï2

22


(Cetac) desolvator and ionized in argon plasma at 1400W. The other instrument parameters

are listed in Table 2.1.

For lithium measurements a "wide-angle" (WA) skimmer cone was used because this

increases sensitivity, compared with a standard cone, by a factor of 2 - 3. With this

configuration total ion beam intensities of (10 to 14) x 10"9 amps per ppm of Li in solution

were obtained at an uptake of ~ 80ul/min. Samples were run at a mass resolution of 700-800

(m/AM; AM is defined at 5% peak height). This is sufficient to resolve doubly charged C

and N on masses Li (required resolution 400) and Li (required resolution 490),

respectively. We scanned for 6LiH+ on mass 7Li at a resolution of -1400 (required resolution

1010) but no hydrides have been detected. Each measurement consists of 40 to 60 cycles of

10 seconds of integration resulting in a total integration time of 400 to 600 seconds. Our

standard procedure uses between 30 - 50 ng Li per analysis. However, a useful measurement

can still be obtained with ca. 5 ng Li at lower precision.

Concentration measurements of Li were performed on a quadrupole ICPMS (VG

Elemental PlasmaQuad 2+). All samples were diluted by 1000 to 3000 times and measured

relative to a L-SVEC standard solution. High dilution factors were chosen to keep the lithium

memory low. Matrix effects were corrected for using beryllium as an internal standard. The

least square regression coefficient of the calibration line was always between 0.98 and 1. The

reproducibility of the concentration measurements is ± 10% (2SD) based on multiple standard

analyses.

2.4 Results and discussion

2.4.1 Separation oflithium

The first 8 ml of IM HNOs/methanol removes Ti, Zr, Hf, Nb and Ta from the resin.

From 9 to 16 ml lithium is eluted. Sodium, the next element to elute after lithium, is not

present in the first 20 ml, i.e. an excellent separation of lithium and sodium is achieved. The

earliest breakthrough of Li is 7Li-enriched whereas the tail is enriched in 6Li (Moriguti and

NAKAMURA, 1998a; KoSLER et al., 2001). This means that 6Li is preferentially taken up by

AG 50W-X8 resin. This is because of a difference in the zero point energy and vibrational

frequencies of the two isotopes (Bighlkishn, 1965). The process must be essentially an

equilibrium fractionation between AG 50W-X8 resin and IM HN03/methanol rather than

23

Chapter 2 Analytical procedure s

diffusion, since the lighter isotope would elutc first in a kinetic process. However,

fractionation of lithium on AG 50W-X8 can be as high as 100%c for small-sized columns (see

Fig. 4 in Kosler et al„ 2001). Thus complete recovery of lithium is essential to avoid

chromatographic fractionation effects. The quantitative recovery of lithium is monitored by

120000

100000

80000

Fig, 2,1: Elution curve of lithium for silicate rocks. AC 50W-X8 ion exchange resin is usedfor: eclogite OK-1,

basalt JB-2, andésite AGV-2 and rhyolite JR-2, respectively. Note that all Li eluted within 16 ml independent

from the rock type. This is especially interesting for high-magnesium rocks. If any loss of Li is possible, then it

could happen due to longer tailing of lithium. This would shift lithium isotope composition towards heavier

values as previously well documented (MOKIGim and NAKAMURA, lWHa; Ko&LhR et ai, 2001).

concentration measurements prior to separation and thereafter. Samples for which less than

100% lithium (within analytical errors) was recovered have not been taken for isotope

analysis.

Chan et al. (2002b) reported a significant change of the elution behaviour of lithium

related to high magnesium and iron concentrations in volcanic rocks. In order to test whether

different rock matrices influence the elution parameters of our analytical procedure we have

analysed a series of samples with a range of MgO and SiÜ2 content: OK-l eclogite from

Oberkotzau, Germany (51% Si02 and >10% MgO; Schramm, 1993); the rock standards JB-

2, a basalt (53.3% Si02 and 4.6% MgO), AGV-2, an andésite (59.3% Si02 and 1.8 % MgO)

24


and JR-2, a rhyolite (76% Si02 and 0.04% MgO). Results of the column calibration are

shown in Fig. 2.1. In all these cases, lithium was quantitatively recovered between 9 and 16

ml.

Two mineral separates with completely different composition were also tested - an

olivine from Iherzolite 8520-9 (WIECHERT et al., 1997) and a spodumene from a zoned

pegmatite (Tin Mountain mine, Harney Peak, USA; SlRBHSCU and Nabhlek, 2003). Lithium

is a trace element in olivine (Skitz and Woodland, 2000; Paquin and Ai.thhrr, 2002;

PAQUIN et al., 2004). On the other hand, lithium in spodumene (nominally LiAlSi20h) is a

major element, thus, matrix is absent because Li itself is the major element. This makes such

minerals useful for testing whether the separation is suitable for these extreme compositions.

The results (Fig. 2.2) show that even such different compositions were not able to disturb the

separation scheme. Lithium is eluted from olivine within exactly the same span as from all

rocks mentioned above, starting with 9.ml and ending by 16.ml without concomitant elution

of magnesium. The only problem follows from the huge amount of Li in spodumene. Thus,

the elution window for Li has to be broadened until 19.ml is eluted because the Li tail is still

50000

40000 "-

w30000"

Q.O

20000

10000 +

0

Clo

8520-09 ol

-a—i)—a—if- a—P

18

500000

400000

300000 -"

200000 -"

100000

= spodumene

10 12 14 16 18

ml

20

Fig, 2,2: Elution of lithium for mineral separates, a) olivine from Iherzolite S520-9 and b) spodumene from

Harney Peak (USA). Note that elution of lithium is within exactly the same volume as for rocks. Opening the

windowfor spodumene is caused by large amount oflithium present in the sample.

25

Chapter 2 A nalytical procedure i

high. This test also shows that an excellent separation of large abundances of lithium from

other elements is achieved. Therefore, only unusual and dramatically high lithium

concentration in rocks or minerals may change lithium elution parameters, i.e. the method is

matrix independent for most silicate materials.

2.4.2 Instrumental mass fractionation

Nu 1700 can accommodate the mass dispersion of -15% suitable for simultaneous

detection of both lithium isotopes. However, all plasma source mass spectrometers suffer

from a mass bias effect (Freedman, 2002) that increases towards lighter masses. It depends

on cones, extraction lenses and many other parameters. To detect natural stable isotope

variations it is necessary to distinguish between instrumental and natural fractionation. For

stable isotopes in the lower mass range this is only possible by comparing samples with a

standard by using the so-called "sample-standard bracketing" technique (Tomascak el al.,

1999a). During this study 40-60 individual isotope ratios for each sample and standard run

were measured. The isotope ratio of the sample was calculated relative to the average of two

bracketing standards (Fig. 2.3). Typically three to four L-SVEC-normalizcd isotope ratios for

one sample were measured in one analytical session. The mass spectrometer extraction and

focusing potentials were tuned daily to yield the best signal stability. No tuning was

performed during the measurements because any changes of the instrumental setting would

change the mass bias. The precision of a single run varied between 0.06-0.15%o (2SE)

whereas errors calculated from averages of at least 4 sample runs give < 0.5%c (2SD).

To ensure that the Nu1700 was functioning correctly, a lithium solution prepared from

LiNOs (Merck) was measured before each analytical session. With the LiNO^ the accuracy

and precision of the mass spectrometer can be monitored, excluding any effects from column

chemistry. The isotope composition of the LiN03, measured over the past 12 months, is 0.91

± 0.15%c (2SD, n=15) relative to L-SVEC. This error also states our best possible long-term

reproducibility because this solution is not treated with the column chemistry procedure.

In order to evaluate possible background contributions to the analytical uncertainty of

the measurements, wc tested a DSN-100 (Nu Instruments) and Aridus (Cetac) desolvating

nebulizer in standard mode and time-resolved analysis (TRA) mode. A key difference of the

two analysis routines lies in how the zero correction is carried out. The standard mode

measures zeros by deflecting the ion beams whereas in TRA mode zeros arc measured with

26


14.34"-

14.32—

14.30"

_J14.28

"

-J

14.26"

14.24"

14.22"

14.20"

14.18 +

0123456789 10

run number

Fig. 2.3: Standard-sample-standard bracketing technique. Li/Li ratios of 4 single sample measurements

(squares) bracketed by 5 L-SVEC analyses (diamonds) are shown. Each symbol represents 40-60 single

measurements. The slight increase of the sample 7Lif'Li ratios is corrected by a similar increase of the ratio for

bracketing standard.

ion beams on cups, i.e. for the zero measurement a clean zero or blank solution is applied.

Typically zero measurements on mass 7 were in the range of several mV whereas 40 ppb Li

in solution gave between 5 and 7 volts (uptake rate 80 uj.tnin'1). The 7Li/6Li ratio of the

background has been detected at -12,6, while measured ratios of L-SVEC varied between

13.8 and 14.4 depending on cones, gas flow conditions and lens settings. To find the best

analysis procedure for lithium isotope measurements, L-SVEC solutions with concentrations

ranging from 10 to 45 ppb are compared with a 50-ppb L-SVEC standard solution. The

measured isotope ratios for these L-SVEC solutions vary by -3.5%e for the 10 ppb solution

relative to the 50-ppb solution for the DSN-100 when measured in standard mode (see Fig.

2.4). When the standard mode is employed together with the Aridus, lithium isotope ratios are

altered by -2.0%c (sec Fig. 2.4). When the TRA mode is applied, isotope ratios are

independent of sample concentrations (see Fig. 2.4).

We conclude that the mass bias effect on lithium isotopes is, within error, concentration

independent. It is likely that insufficient time elapses to wash out lithium between standard

and sample analysis. Between samples and standards it has been washed with 2% HNO3 for

120 seconds. Methods measuring zeros by deflecting ion beams cannot correct for any

Û Û

27


1.0

0.0'—-

E-1.0

0)Q.

u-2.0

1X1

>w

_l -3.0_lr-

to-4.0

-5.0

\ ^^^===^

ARIDUS (TRA; zero by blank solution)

ARIDUS (zero by ion beam deflection)

DSN-100 (zero by ion beam deflection)

_i i i i i i i i i i i i i i i i i

0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0

'sp/'50ppb

Fig. 2.4: Lithium, isotope ratio measurements ofL-SVEC solutions with different concentrations using an Aridus

(Cetac) and DSN-100 (Nu Instrument) desolvating nebulizerfor sample introduction. Plotted are isotope ratios

versus concentration of the sample solution relative to 50-ppb L-SVEC solution as total beam intensity (I) ratios

(I-SA/I-50ppb), Both Aridus and DSN-100 show a "concentration" effect when zero values are obtained by ion

beam deflection with the electrostatic analyser (ESA). In contrast, isotope ratios obtained by time-resolved

analysis do not show a concentration dependence of Li isotope ratio measurements (see text for discussion).

Symbols represent average values of at least 4 single measurements. Errors are 2SD.

lithium memory independent from the used desolvator. Better results are achieved by

measuring zeros using blank solution in TRA mode. Alternatively concentrations of sample

and standard solutions could be balanced within +10%. Then memory effects become

negligible independent of which method is used (Fig. 2.4).

Any matrix element in the sample solution may change the instrumental fractionation

(mass bias) and alter the isotope ratio of samples in a different way to the standard (e.g. Zhu

et al., 2002). Therefore, all solutions have been monitored for Na, Mg, Ca, Al and Fe. These

elements have been selected because of distribution coefficients similar to lithium in

analytical procedure or because they are major elements. In no case did we observe

significant amounts of Na, Mg or any other of these elements in the Li aliquots. In the worst

case the total fraction might make up 20% of the lithium. Such amounts would not have any

28


detectable effect on lithium isotope ratios. Despite not observing any significant amounts of

alkalis beside lithium, the effect of large concentrations of alkali elements in the final eluate

was tested. Pure L-SVEC solutions were individually doped with Na, Mg, Al, and a multi¬

element matrix consisting of Na, Mg, AI, Ca, K and Fe. Solutions with Na/Li, Mg/Li, and

Al/Li ratios of 1:1 and 5:1 and multi-element/Li solution with equal concentrations of all

elements were prepared and the isotope ratio of lithium was compared with pure L-SVEC

standard solution (Fig. 2.5). Sodium has no effect on lithium isotope ratio in either Na/Li

proportions. Magnesium causes slightly higher lithium isotope ratios but only at higher

concentrations (Mg/Li 5:1). Low magnesium concentrations (Mg/Li 1:1) have no detectable

effect. Aluminium shifts the Li isotope ratio toward lighter isotope ratios already at low Al/Li

(1:1). However, the effect is not concentration dependent because higher Al/Li (5:1)

fractionate lithium isotopes by exactly the same extent. The multi-element solution

(Na+Mg+Al+K+Ca+Fe/Li 1:1) shows no effect when compared with pure L-SVEC. This

suggests lithium isotope measurements on the Nu-1700 can be affected by matrix elements at

Fig. 2.5: The effects of matrix elements. L-SVEC solutions doped with Na (diamonds), Mg (circles), Al

(squares), and a multi-element (Na, Mg, AI, Ca, K and Fe) matrix (triangle) are compared with pure L-SVEC

solutions. Each symbol represents the average of4 single measurements. The grey field gives the reproducibility

of lithium isotope ratios for pure L-SVEC solutions. Open symbols represent element/lithium ratios of 1:1,

closed symbols 5:1. Error bars are 2SD.

29


only high concentrations (see also BRYANT et al., 2003b). However, for the data reported

here, matrix elements did not make up more than 20% of the lithium, and this precludes any

alteration of the isotope ratios by the matrix effects discussed above.

2.4.3 L-SVEC measurements

The international reference material for Li isotope analysis, L-SVEC, was prepared

from several kilograms of high-purity carbonate and its Li isotope composition measured.

This yielded an absolute ratio of 'Li^Li = 12.02 ± 0.03 (Flesch et al., 1973). However, such

a large error is greatly in excess of that needed for geological problems because lithium

isotope variations in nature arc never larger than several tens of per mil (sec TOMASCAK

(2004) and references therein). GALY et al. (2003) have shown that Mg international

reference material SRM 980 suffers from large amounts of heterogeneity in 2,;Mg/24Mg and

6Mg/24Mg. Therefore, we tested the homogeneity of Li L-SVEC. We measured three aliquots

of L-SVEC from three international labs (Univ. Maryland, USA; Open Univ., UK; Charles

Univ. in Prague, Czech Rep.) together with our own aliquot that has been used as a bracketing

solution. The results are shown in Fig. 2.6. All three aliquots were identical to our own within

1.00

0.75

E0.50

CD

30.25

O

>CO

0.00

_l -0.25

-0.50

-0.75

-1.00

Fig. 2.6: Comparison of several selected L-SVEC aliquots. It is clear that this reference material is

homogeneous in terms of lithium isotope composition. Open symbols represent individual runs; closed symbols

are the averages of the respective runs with 2SE error bars (diamonds - P. Tomascak, ifLi - 0.01 + 0.05%a;

circles - J. KoUer, Shi = 0.04 ±0.05%c; squares - R.II. James, SLi = 0.01 ±0.07%«; solid line is the average

ofour L-SVEC aliquot with dashed lines as 2SE error bars, Su = 0.003 ±0.099%o).

30


± 0.03%c. Thus, this standard would appear to be homogeneous and very suitable for high

precision isotope ratio mass spectrometry.

2.4.4 International reference rocks

The data for international reference rocks from this study are presented in Table 2.2.

The error of all averages is given as 2SD. Outliers are excluded from the calculations of

averages using 3a-confidence limits. There is no systematic correlation between measured

concentrations and isotope ratios of individual samples.

We chose a broad variety of reference rocks from basic to acidic compositions to test

whether our column chemistry is sensitive to different igneous rock types and to determine

precision and accuracy of the new lithium separation scheme on the 2.1ml columns and using

MC-ICPMS. Different volcanic rocks ranging from high MgO, low Si02 concentrations to

high Si02, low MgO chemistry were selected (see also section 2.3.1).

The 57Li value of 4.7 + 0.3%o (2SD) for JB-2 from this study (10 separate digestions of

JB-2) is in good agreement with previously published values 5.1 ± l.l%c (2SD; Tomascak et

al., 1999a) and 4.9 ± 0.7%c (2SD; Moriguti and Nakamura, 1998a) and within analytical

error of 4.3 ± 0.3%c (2SD; NiSHio and Nakai, 2002). However, this value is significantly

different from a 87Li of 6.8 ± 0.2%<? (2SD) reported by James and Palmer (2000). A

compilation of available data for JB-2 is given in Fig. 2.7.

For BHVO-2 a total of 9 averages gives a 87Li of 4.55 + 0.3%c (2SD). This is in

excellent agreement with S7Li of 4.5 ± 1.0%0 (2SD) reported by Zack et al. (2003). A 57Li

value of 5.3 + 0.2%c (2SE) has been obtained for BHVO-1 which is in the middle of the range

of published 67Li values from 5.0 to 5.8%« (James and Palmer, 2000; Bouman et al., 2002;

Chan and Frky, 2003). Although both BHVO-1 and BHVO-2 come from the same lava

flow, they are distinct by about 0.8%e in ö7Li and do not overlap within analytical error in our

study. This may be explicable by some small-scale heterogeneity. Heterogeneity for lead

isotopes has been reported for BHVO-2 (Woodhead and Hergt, 2000).

For JR-2 a Ô7Li of 3.8 ± 0.2%0 (2SE) has been measured which is within analytical error

the same as reported by JAMES and PALMER (2000; Ô7Li = 3.9 + 0.8%o; 2SD) and 3.9 + 0A%c

(2SD) reported by Chan and Frey (2003) using TIMS techniques. It is interesting to note that

this rhyolite still has a Ô Li similar to that for pristine mantle rocks (Chan, 2003) consistent

31


with no isotope fractionation during magmatic differentiation as reported by Tomascak et ai.

(1999b) for a sub-alkalic to alkalic basaltic rock suite.

Table 2.2 S7Li values for international reference rocks

Li (ppm)

measured

Li (ppm)

recomm.

87Li (%o) 2 SE n" Average (2SD) Other studies

57Li {%«) 2SD

AGV-1 12.Ü 12.0 6.74 0.20 3

BHVO-1 4.4 4.6 5.31 0.18 3 5.0b5.8e

5.21'

1.4

1.6

0.5

BCR-1 12.5 12.9 2.56

2.19

0.06

0.33

2

2 2.38 (0.52)

AGV-2 10.4

12.4

11.0 8.28

7.80

7.65

8.24

8.12

0.23

0.35

0.44

0.09

0.34

4

4

2

2

4

10.9 7.53 0.43 3 7.94 (0.64)

BHVO-2 4.8

5.0

4.4

5.0

4.9

5.0

4.7

5.0 4.50

4.84

4.36

4.65

4.46

4.52

4.40

4.58

0.28

0.08

0.36

0.22

0.14

0.31

0.16

4

1

2

3

3

3

2

4

4.62 0.13 3 4.55 (0.29) 4.51' 1.0

JB-2 8.4

7.8

8.2

7.8

7.8 4.93

4.90

4.69

4.66

0.33

0.21

0.04

0.28

3

2

4

3

4.44 0.07 4 6.8e 0.2

8.1 4.73

4.70

0.48 3

1

5.1"

4.7e

0.4

1.0

4.65 0.16 3 5.11 1.1

4.55 0.06 3 4.9s 0.7

8.4 4.73 0.12 2 4.70 (0.29) 4.2h 0.3

JR-2 78.6 79.2 3.84 0.18 3 3.9e

3.9d0.8

0.4

recomm. - recommended

a

number of individual measurements

hBouman et al. (2002)

c

James and Palmer (2000)dChan and Frey (2003)

eZacketal. (2003)fTomascak el al. (1999a)

8Moriguti and Nakamura (1998a)

1

Nisnio and Nakai (2002)

32


A 57Li of 7.9 ± 0.6%c (2SD) measured for AGV-2 is clearly above the ô7Li range of

fresh MORB mantle rocks (ZACK et al., 2003). The Li in AGV-1 and AGV-2 is slightly

different. The ô7Li value of AGV-1 (6.7 + 0.2%o, 2SE) is substantially lower.

8.0

7.0 -

Q. 6.0

o

>CO

5.0

4.0 -

3.0

James and Palmer (2000)

AA

4

4

Moriguti and Nakamura (1998a)A

This study

^_ _>

Tomascak et al. (1999a)

Chan and Frey (2003)

Zack et al. (2003)Nishio and Nakai (2002>

Fig. 2.7: Compilation of Su. values for GSJ reference material JB-2 from the literature and this study. The grey

bar shows + 2SDfor JB-2 obtainedfrom this study (0.29c/ot>).

2.4.5 Other reference materials and rocks

Lithium abundances and isotope compositions of several other materials are

summarized in Table 2.3.

The lithium isotope composition of NIST SRM 612, the nominally 50-ppm multi-

clement glass standard used for laser-ICPMS, is close to seawater (S7Li - 35.3 + 0.8%o, 2SD).

The Li content of ~33 ppm measured during this study is lower when compared to the

nominal value which reflects heterogeneity of this reference material as shown also by other

elements (see Pharce et al. (1997) and references therein). This may also be the source of the

poorer reproducibility of this reference material.

33


A 57Li = -8.3 ± 0.2%c (2SD) for the eclogite OK-1 is consistent with a MORB protolith

(Schramm, 1993) and fractionation of lithium isotopes related to subduction and dehydration

as proposed by Zack et al. (2003) for eclogitic rocks from Trescolmen, Switzerland.

Recently, even more negative 87Li values were reported for peridotitic rocks from far-east

Russia that have been metasomatised by an eclogitic melt (NlSHio et al., 2004).

Table 2.3: 87Li values for other rocks and standard materials

I a (ppm)measured

Li (ppm)recomm.

ö7Li (%c) 2 Sh Average 2 SI)

N 1ST 612

Eclogite OK-1

(Miinchbcrg, Germany)

8520-09 olivine (Mongolia)

Spodumene

(Harney Peak, USA)

Li nitrate

30.2

35.5

10.1

9.3

9.8

10.3

1.4

0.04"

43.17e 35.07 0.47 3

34.79 0.25 4

35.55 0.39 3

35.62 1 35.26 0.79

-8.30 0.08 2

-8.24 0.09 2

-8.16 0.11 4

-8.40 0.13 4 -8.27 0.20

3.60

7.23

0.36

0.15

0.88 0.11 4

0.98 0.03 2

0.87 0.17 4

1.04 0.30 7

0.88 0.11 4

1.03 0.23 5

0.91 0.14 4

0.97 0.11 4

0.93 0.14 2

0.92 0.25 4

0.88 0.32 4

0.93 0.23 4

0.84 0.15 3

0.77 0.16 3

0.83 0.09 3 0.91 0.15

recomm. - recommended

anumber of individual measurements

bConcentration matched to 40ppb of L-SVEC

'Pcarcectal. (1997)

The olivine fraction of peridotite 8520-09 has a low Li concentration of 1.4 ppm, which

is consistent with previous findings for garnet peridotites from Alpe Arami (Central Alps,

Switzerland; Paquin and Althkrr, 2002) and for mineral separates of peridotitic rocks from

Kenya, France, Australia, Italy and Germany (Skit/, and Woodland, 2000). Isotope analysis

34


of this olivine yielded 8 Li - 3.6 + 0.4%c (2SE). This is consistent with the conclusions, that

ô Li of unmetasomatised mantle is 4.2 ± 0.8%o (CHAN, 2003) similar to values for peridotites

reported by Chan et al. (2002b) and Brookhr et al. (2000). Spodumene from Tin Mountain,

Harney Peak, has 8 Li - 7.2 + 0.2%c (2SE). This is similar to lithium isotopes measured for I-

typc granites from Australia (Bryant et al., 2003a). However, granites from the Harney Peak

region are considered S-type, which are expected to have low ö7Li values. Therefore, the Li-

isotope composition of this pegmatite is unlikely a magmatic source signature. It may be

related to late stage fractionation and/or alteration process occurring in that region (Sirbhscu

and Nabelek, 2003). Further studies are required to solve this paradox,

2.5 Conclusions

We have presented lithium isotope data for 7 international reference rocks and some

other rocks and materials that comprise a wide range of silicate rocks. The results obtained in

this study are in good agreement with published data showing that our method for Li

separation is suitable for the basalt-to-rhyolite span of silicate rocks. The technique is also

suitable for other materials (Li minerals) that can be treated in exactly the same way as whole

rock samples without any influence on the isotope ratios. A precision of <0.5%e (2SD)

obtained over 12 months is comparable to or better than that reported for larger columns. We

conclude that much smaller acid volumes and therefore low blanks allow us to analyse -5 ng

lithium to high precision. This opens up the field for new studies on sample-size-restricted

objects and/or materials with low lithium concentration.

Acknowledgement - We thank J. Kosler (Charles University, Prague, Czech Republic) for

useful suggestions about the column chemistry and comments on earlier versions of the

manuscript and M. Vobecky (Academy of Sciences, Prague, Czech Republic) for fruitful

debates on resin problems. Mona Sirbescu (Central Michigan University, Michigan, USA)

provided sample of spodumene. J. Kosler, P. Tomascak (Univ. Maryland, USA) and R.H.

James (Open Univ., Milton Keynes, UK) kindly provided aliquots of L-SVEC. This work

was supported by ETH and Swiss Nationalfonds.

35

Seite Leer /

Blank leaf j

36

Chapter 3 Southern Cascades

Chapter 3

Lithium isotope fractionation in the southern

Cascadia subduction zone

*

published as: Magna T., Wiechert U., Grove T.L. and Halliday A.N. (2006) Lithium isotopefractionation in the southern Cascadia subduction zone. Earth and Planetary Science Letters

250 (3-4), 428-443

37


3.1 Abstract

We present lithium (Li) abundances and isotope compositions for a suite of anhydrous

olivine tholeiites (HAOTs) and hydrous basalt-andesitic (BA) lavas from the Mt. Shasta and

Medicine Lake regions, California. The values of 57Li vary from +0.9%o to +6.4%o and

correlate inversely with distance from the trench. These data are consistent with continuous

isotope fractionation of Li during dehydration of the subducted oceanic lithosphère, an

interpretation corroborated by uniformly high pre-emptive H20 contents in basaltic andésites

accompanied by high Li, Rb, Sr, Ba and Pb abundances. The subduction-derived component

that was added to these hydrous magmas is shown to be very similar beneath both Mt. Shasta

and Medicine Lake volcanoes despite characteristically distinct Li isotope compositions in

the magmas themselves. More evolved andésites and dacites from Mt. Shasta have 8 Li from

+2.8 to +6.9%o which is identical with the range obtained for HAOTs and BA lavas from Mt.

Shasta. Therefore, Li isotopes do not provide evidence for any other crustal component

admixed to Mt. Shasta andésites or dacites during magmatic differentiation and magma

mixing in the crust.

3.2 Introduction

At convergent plate boundaries cold and brittle oceanic lithosphère descends into the

more ductile asthenospheric mantle resulting in explosive volcanism, high-magnitude

earthquakes and fast morphological evolution of the overriding plate. Several processes may

ultimately lead to melt production: (i) percolation of aqueous slab fluids into the mantle

wedge (Tatsumi and MARUYAMA, 1989; GROVE et al., 2002), (ii) partial melting of the slab

(KELEMEN, 1995), (ni) reaction of slab melts generated by partial melting of the subducted

slab with the overlying mantle (Kay, 1978; Yoüodzinski et al., 1995), or (iv) more

complicated two-step processes such as hydrous metasomatism of cold lithosphère and

subsequent melting of "wet" mantle segments related to back-arc rifting (DORENDORF et al.,

2000).

Lithium isotope fractionation is mainly related to low temperature processes at the

surface and provides a geochemical tracer of recycled crust in volcanic rocks. Volcanism in

the Cascadian arc in California is dominated by basalts with subordinate andesitic to dacitic

lavas (GILL, 1981; GROVE and KiNZLER, 1986). These evolved lavas are multi-component

38


mixtures of mantle melts and andésite, dacite and rhyodacite magmas produced by fractional

crystallization. The high-alumina olivine tholeiites (HAOTs) and basalt-andesitic (BA) lavas

of this study do not show a strong overprint of fractional crystallization, crustal

contamination and/or magma mixing, making them excellent samples for studying processes

in subduction zones (GROVE et al., 2002).

Lithium exhibits large isotope fractionation in the subduction zone environment

(Moriguti and Nakamura, 1998b; Chan et al., 2002b; Zack et al., 2003). Generally, it is

believed that heavy isotope inputs are provided into the subduction zone by scrpentinized

peridotites that may have ö7Li values as heavy as +20%o (DECITRE et al., 2002; Benton et

al., 2004) and altered oceanic basalts (CHAN et al., 1992, 2002a; JAMES et al., 2003). Altered

seafloor is substantially enriched in Li over fresh ridge basalts to as much as ~80 ppm

accompanied by distinctly heavy isotope compositions ~+14%o (Chan et al., 1992, 2002a;

Seyfried et al., 1998). However, the concept that heavy Li is transferred into the subduction

zone environment has been recently questioned because of the light or "mantle-like" ô Li

values (mainly between -1 and +5%o) found in pelagic clays, oozes, muds and claystones

(Bouman ct al., 2004) combined with the observation that arc lavas have predominantly

MORB-like Li isotope ratios (Tomascak et al., 2002). Progressive shifts toward very

negative S Li had been already recognized and were related to preferential loss of Li into the

fluid phase leaving the residue isotopically light (Zack et al., 2003). Indeed, the extent of

depletion in Li is greater with larger degrees of seafloor alteration. However, the behaviour

of Li in subduction zones still remains largely unexplored and rather ambiguous Li cross-arc

trends (i.e. linked or not linked to the fluid expulsion from subducted slab) were revealed in

various subduction zones and volcanic arcs, e.g. Japan (Moriguti and Nakamura, 1998b;

Moriguti et al., 2004), northern Cascades (Leeman et al., 2004), Mariana arc (Bf.nton et

al., 2004), Kurile, Aleutian and Sunda arcs (TOMASCAK et al., 2002), Central American

Volcanic Arc - CAVA (Chan et al., 1999, 2002b) or Panama arc (Tomascak et al., 2000). It

is thought that the initiation of slab melting may produce extremely light Li isotope

compositions (Nismo et al., 2004).

In this paper Li abundances and isotope data are presented for HAOTs and primitive

BA lavas from Mt. Shasta and the Medicine Lake regions, northern California. These data

are used to place new constraints on the Li isotope geochemistry of convergent plate margins

and the inputs of Li to the source of arc volcanics.

39


3.3 Geology of the region

Mt. Shasta lies near the southern end of the Cascades chain and is flanked to the west

by pre-Tertiary rocks of the Klamath Mountains province. In the northeast and east Mt.

Shasta is bordered by Tertiary and Quaternary volcanics of the High Cascades. Medicine

Lake lies about 60 km to the east. This back-arc shield volcano has erupted large amounts of

basaltic to rhyolitic volcanics. The Lassen Volcanic Field is ~130km to the south. Tertiary

sandstones, shales, and andesitic volcanics, Mesozoic granitic rocks of the Trinity ophiolite

and Mesozoic and Palaeozoic meta-sedimentary rocks are present in the Mt. Shasta region

and have been inferred to be present beneath the volcano (Fuis et al., 1987).

Subduction zone volcanism in northern California is related to the Juan de Fuca plate,

which is being subducted at a half spreading rate of 40 mm per year (WILSON, 1988). The

location of the subducted slab can be estimated at -110 km depth below Mt. Shasta and has

been inferred to be at -200 km depth below the Medicine Lake region (Zucca et al., 1986;

HARRIS et al., 1991; Benz et al., 1992). The age of the oceanic plate beneath Mt. Shasta is

inferred to be 12-14 million years (Green and Harry, 1999). Four major eruptive phases

comprise the volcano, Sargents Ridge (<250—130 ky), Misery Hill (80-10 ky), Shastina (10-

9.4 ky) and Hotlum (6-2 ky; CHRISTIANSEN et al., 1977). Previous geochemical and

petrologic studies of Mt. Shasta include those of, for example, ANDERSON (1974).

Some of the samples from this study have been characterized by GROVE et al. (2002,

2003, 2005), Baker et al. (1994), Kinzler et ai. (2000), Wagner et al. (1995) and Elkins

Tanton et al. (2001). The reader is directed to these authors for further petrologic and

geochemical descriptions. The suite of samples comprises two Mt. Shasta and six Medicine

Lake primitive high-alumina olivine tholeiites (HAOTs), four basaltic andésites (BA) and

one primitive magnesian andésite (PMA) from Mt. Shasta and two basaltic andésites and two

samples of the Lake Basalt (WAGNER et al., 1995) from Medicine Lake. Additionally, we

have analysed six andésites and eight dacites from the Mt. Shasta stratocone.

3.4 Analytical techniques

Samples were digested in a mixture of doubly distilled concentrated HF and HNO3 (6:1

by volume), evaporated to dryness and treated three times with concentrated HNOv

Subsequently the samples were dissolved in 6M HCl and loaded on columns. A single step of

40


ion exchange chromatography (BioRad AG50W-X8, 200-400 mesh) was used to separate Li

from matrix elements, using IM HNO3 made up in 80%-methanol as elution media. A

detailed description of the method is given in Magna et al. (2004b). The purified Li

solutions were measured on the large-geometry high-resolution multiple-collector ICPMS

Nu 1700 (Nu Instruments) at m/Am -700. An in-house standard has been measured before

measuring unknown samples to monitor the quality of the mass spectrometric measurements.

International reference rock standards such as JB-2 from the Geological Survey of Japan

(GSJ) were prepared with unknown samples to check for possible lithium isotope

fractionation on the ion exchange columns and/or blank effects. Samples were measured with

a standard-sample bracketing method using L-SVEC as a reference (Flesch et al., 1973).

Lithium isotope ratios are calculated as 57Li (%o) = [(7Li/6Li)h;impic/(7Li/6Li)ilandard-l]x1000.

The in-run precision of single measurements varied typically between 0.05 and 0.10%o (2SE).

Three to five individual runs were performed for each sample from which the mean and

2a analytical uncertainty were calculated. The long-term reproducibility and precision of our

in-house Li standard is 57Li = 0.94 + 0.14%0 (2a, n=28). Analyses of JB-2 prepared together

with Mt. Shasta volcanic rocks averaged 57Li = 4.79 ± 0.29%o (2a, n=4) in four analytical

sessions. This is in excellent agreement with published data (Moriguti and Nakamura,

1998a; TOMASCAK et al., 1999a; CHAN and FREY, 2003; JEFFCOATE et ai., 2004; MAGNA et

al., 2004b).

Concentration measurements were performed on a quadrupole ICPMS VG Elemental

PQ2 prior to separation and thereafter using beryllium as an internal standard for matrix

element correction. The Li concentrations so determined were used to calculate yields and to

reject samples with <95% recovery. We also used the PQ2 to monitor each sample for the

presence of matrix elements (Na, Mg, AI, Ca, Fe) after ion-exchange chromatography.

Samples with a total matrix content that exceeded 20% of the Li abundance were not used for

isotope analysis. Trace element concentrations, with the exception of Li, were determined on

the PQ2 using rhodium as an internal standard. The accuracy and precision of trace element

analyses were checked against analysis of JB-2 (GSJ), BHVO-2 and AGV-2 (USGS).

Barium contents were higher by about 10-15% relative to reference values. Accordingly, Ba

abundances for tholeiites, basaltic and magnesian andésites were corrected down by 15% and

for andésites and dacites down by 10%. The 2a errors on concentration measurements were

better than 5% for most elements, with some scatter for Cs (<10%).

41


Strontium (Sr) and neodymium (Nd) were separated following procedures of Horwitz

ct al. (1992) and Cohfn et al. (1988), respectively. Strontium and Nd isotope compositions

were measured on a Nu Plasma MC-ICPMS (Nu Instruments). Masses 83 and 85 were used

to correct for Kr and Rb interferences, respectively, that were <1% of the respective Sr

isotope beam. The instrumental mass bias was corrected using an exponential fractionation

law. The 87Sr/86Sr ratios were corrected for mass bias using X8Sr/8f,Sr = 8.3752 and

,43Nd/144Nd was corrected using a 146Nd/144Nd of 0.7219. All 143Nd/,44Nd ratios were

normalized to the JMC-Nd standard using 143Nd/144Nd = 0.511833 (cross-calibrated to a ratio

of 0.511858 for La Jolla standard). 87Sr/86Sr ratios were normalized to the accepted ratio of

0.710245 for the NIST SRM 987. These new Sr data are within +0.01% (le unit) of previous

Sr isotope data for Medicine Lake olivine tholeiite 82-72a (Baker ct al., 1991) and basaltic

andésite 79-24E (Kinzler et al., 2000).

3.5 Results

Lithium abundances and isotope data for Mt. Shasta and Medicine Lake lavas are

presented with other radiogenic isotope data in Tables 3.1 and 3.2 (see Grove et al. (2002)

for Sr, Nd and Pb isotope data for Mt. Shasta lavas). Lithium contents of olivine tholeiites

and basaltic andésites are within the range reported for arc volcanics from southern

Washington Cascades (Leeman et al., 2004) and Central America (Chan et al., 2002b) but

are about half of those of subduction-related calc-alkaline lavas of the Panama arc

(Tomascak et al., 2000). The Li abundances range from 3.2 to 7.8 ppm in HAOTs, are

between 5.0 and 11.4 ppm in BA and Lake Basalt lavas, and yield a value of 10.2 ppm for

PMA sample 85-4lb. Most andésites and dacites have Li concentrations ranging from 6 to 11

ppm. The ô7Li values of HAOTs, BA lavas and PMA vary between +0.9%o and t6.4%0

relative to L-SVEC with only slight overlap between Mt. Shasta and Medicine Lake

volcanics (Fig. 3.1). A subset of two andésites and two dacites exhibits high Li abundances

(up to 35 ppm) compared to other evolved lavas; this is paralleled by high Rb, Cs, Ba, Pb, Th

and U concentrations and slightly elevated 87Sr/86Sr. All andésites and dacites from Mt.

Shasta have 87Li values between +2.8%o and 16.9%o which is within the range of BA lavas

and HAOTs.

42


Table 3.1 Lithium concentrations and isotope compositionsin Ml. Shasta and Medicine Lake lavas

sample type Li 57Li 2ct n'

ML Shasta

85-38a

HAOT 6.0 3.04 0.24 1

85-51" HAOT 6.3 3.02 0.62 3

82-94aa BA 7.3 6.41 1.06 3

85-44*

BA 5.5 4.04 1.23 2

95-15a BA 6.8 4.98 0.49 2

95-17a

BA 5.0 5.14 0.42 3

85-41bd PMA 10.2 4.74 0.58 2

82-94dà A 6.0 6.27 0.26

85-55a

A 11.2 2.88 0.08

85-3"

A 8.3 4.50 0.13

85-48b"

A 8.2 6.44 0.27

85-59J

A 18.4 6.35 0.15

82-85J

A 20.2 3.31 0.17

82-92aa D 35.2 2.79 0.40

82-95'*

D 28.0 5.97 0.33 2

82-9 lba D 9.7 5.21 0.22

96-1a

D 9.9 3.06 0.18

97-7a

D 8.3 6.38 0.33

83-50a

D 6.6 4.44 0.13

83-58a

D 10.3 2.99 0.14

95-13" D 9.1 6.89 0.07

Medicine Lake

82-72a" HAOT 3.2 2.42 0.29 2

95-3c

IIAOT 7.8 1.22 0.27

95-4" HAOT 6.9 2.83 0.16

95-5L

IIAOT 4.6 1.19 0.41

95-6°

IIAOT 5.0 3.20 0.09

95-9c

HAOT 5.2 5.02 0.29 2

1665md

LB 6.9 3.28 0.20

1672md LB 6.1 1.95 0.42

1399me BA 8.9 2.36 0.14

79-24Ee BA 11.4 0.93 0.30 2

Reference standards

JB-2 4.79 0.29 4

BHVO-2 4.43 0.41 2

AGV-2 8.14 0.66 2

Li concentrations arc given in ppm. Li isotope compositions expressed in per mil relative to L-SVEC standard.

HAOT - high-alumina olivine tholciitc, BA - basaltic andésite, PMA - primitive magnesian andésite, A -

andésite, D - dacite, LB - Lake Basalt

"see Grove ct al. (2005) for Sr, Nd and Pb isotopes and trace element geochemistry; bsee Bakfr et ai. (1991);c

see ELKINS TANTON ct al. (2001);''see Wagner et al. (1995);

c

see Kinzler et ai. (2000);'number of full

replicate measurements including sample dissolutions

43


Selected trace element concentrations are presented in Tables 3.3 and 3.4, respectively,

and are within 5% of the abundances previously reported for these elements (Baker et al.,

1991, 1994; Wagner et al., 1995; Kinzler ct al., 2000; Grove et al., 2002). MORB-

normalized patterns are remarkably similar with only subtle differences in Nb (Ta) depletion

and REE pattern of HAOTs (Fig. 3.2). HAOT and BA lavas from Mt. Shasta and Medicine

Lake cannot be resolved from each other in their normalized patterns.

CO

CD o

O)ô

-2

CO*

g 1

03jQ

C

Mt. Shasta

Isd„MIIIÉ

Medicine Lake

Jl.12 3 4 5 6 7 8

ô7Li (%o)

Fig. 3.1: Histogram of8Li values in tholeiites, andésites and dacitesfrom Mt Shasta and Medicine Lake. High-

alumina olivine tholeiites are in black, basaltic andésites in dark grey, primitive magnesian andésite with

pattern, andésites are in light grey and dacites are in white. Progressive enrichment in Li is found in Mt.

Shasta BA lavas, whereas BA lavas from Medicine Lake have light oLi values, probably as a result of Li

isotope fractionation during subduction.

44


3.6 Discussion

3.6.1 High-alumina olivine tholeiites

High pressure experiments (BARTELS et al., 1991) show that HAOTs from the Medicine

Lake area originate from near-anhydrous melting of spinel peridotitc consistent with low

water contents of melt inclusions in olivine (SlSSON and LAYNE, 1993). Anhydrous melting

has also been assumed for HAOTs from Mt. Shasta on the basis of very similar major

element chemistry (Grove et al., 2002) and low pre-cruptive water contents in HAOTs from

northern California Cascades (SlSSON and Layne, 1993). Analysis of primitive HAOTs from

the Cascades region has shown that the magmas underwent fractional crystallization of

olivine + plagioclase and the depth of their generation increases eastwards parallel to

modeled corner flow lines rather than dip of subduction (Elkins Tanton et al., 2001). The

Li isotope compositions of HAOTs from Mt. Shasta give an identical Ô Li = +3.0%o whereas

HAOTs from the Medicine Lake region range from +1.2 to +5.0%o. There is no indication

from existing data that Li isotopes fractionate during partial melting or magmatic

differentiation on Earth (Tomascak et al., 1999b). An exception might be the extensive

magmatic differentiation during crystallisation of the lunar magma ocean which produced a

range of cumulate sources with distinct Li isotope signatures (Magna et al., 2006b).

Nevertheless, Li isotope variations observed in HAOTs in this study most likely reflect

heterogeneity of the mantle below the Mt. Shasta and Medicine Lake regions. Such chemical

variability of the mantle sources has been drawn from trace element models in primitive

HAOTs and BA lavas from adjacent Lassen region (Clynne and Borg, 1997). There is no

evidence for assimilation of granitic crust and/or contamination of HAOTs by Trinity

harzburgites, inferred beneath Mt. Shasta (Baker et al., 1994). Therefore, relative fertility of

mantle wedge segments beneath Mt. Shasta and Medicine Lake (Borg et al., 1997) may

account for the observed several per mil scatter in HAOTs.

HAOTs are relatively enriched in fluid mobile elements such as Pb, Sr and Li compared

to MORB (Fig. 3.2). The enrichment of highly fluid mobile elements provides evidence for a

subduction-related component in HAOTs (see also Hart, 1985). This is consistent with

metasomatic enrichment of this mantle segment before melt extraction. Therefore, HAOTs

contain variable proportions of mantle and slab-derived Li. It has been argued that these trace

element patterns provide evidence for a connection with the recent subduction volcanism in

the Cascades (Donnelly-Nolan et al., 1991; Baker et al., 1994; Bacon et al., 1997). In

45


contrast, Hart (1985) noted that radiogenic Sr isotope ratios continue to the east of the arc

and seems not related to modern subduction zone volcanism. The latter is consistent with

radiogenic strontium of 87Sr/K6Sr = 0.70434 for Mt. Shasta HAOTs whereas Medicine Lake

HAOTs range from 0.7032 to 0.70345. At least some HAOTs seem to contain a component

with high time-integrated Rb/Sr that is related to metasomatism by slab-derived fluids some

lime in the past. Although the timing of the enrichment in fluid mobile elements is unclear

the anhydrous lavas provide evidence for an isotopically heterogeneous mantle wedge with

predominant isotopically light Li beneath Medicine Lake.

|~H 1 1 1 1 ) (, 4 4 t \ 1 1 1 1 1 1 1 1 1 1 1 1 1 1

Cs Rb Ba Th U Nb Ta La Ce Pb Pr Sr Nd Sm Zr Hf Eu Gd Tb Dy Li Y Ho Er Tm Yb Lu

Fig. 3.2: Trace element patterns for IIAOT and BA lavas from Mt. Shasta and Medicine Lake (not separated).

The dark grey field is that of HAOTs. The most depleted HAOT lava 82-72a is plotted separately as black

squares and shows a pattern that is similar to that of other samples 82-72al (DONNELLY-NOLAN et al, 1991)

and 82-72f (BACON et al., 1997) studied previously. The HAOTs from Mt. Shasta and Medicine Lake are

identical in their trace element patterns. The light grey field represents BA lavas; PMA lava 85-4 lb is plotted

separately (white diamonds). Two BA lavas and two Lake Basalt lavasfrom Medicine Lake cannot he resolved

from Mt. Shasta BA lavas based on their trace element patterns. All lavas exhibit similar trace elements pattern

with enriched Cs, Ba, Pb, Sr and Li and Nb, Ta depletion. For normalization N-MORB data from Sun and

Mcdonough (1989) are used.

46


3.6.2 Basaltic andésites andprimitive magnesian andésites

The pre-emptive H2O contents of these lavas provide constraints on the composition of

the slab component and the process of melt formation. The BA lavas have high Mg#s

providing evidence for equilibration with spinel peridotite (GROVE et al., 2002). Olivines

with a forsterite content of-93.6 have been reported for the primitive Mg-andesites (PMA).

Experiments show that both BA lavas (85-44; Baker et al., 1994) and PMA (85-41b;

Parman, 2001), are saturated with olivine and orthopyroxene at -1.0 GPa and I200°C. The

olivine crystals in the PMA are igneous and crystallized from a magma that was derived from

an extremely depleted mantle source (Baker et al., 1994). The experiments provide evidence

that the major-element chemistry of BA lavas and PMA from Mt. Shasta are in equilibrium,

or nearly so, with peridotites at the base of the crust (Grove ct al., 2002).

Geochemical variations in primitive BA lavas from the southern Cascades have been

accounted for by reactions of fluid phases with mantle wedge peridotites (BORG ct al., 1997,

2000). However, boron (B) abundance and isotope data place ambiguous constraints on the

nature of the fluid phase in the primitive Mt. Shasta BA lavas and do not distinguish a B-

depleted hydrous fluid from a H20-rich slab melt (ROSE et al., 2001). The major element

chemistry of BA lavas and PMA is in equilibrium, or nearly so, with peridotite. However, up

to 99% of the fluid-mobile trace elements such as Rb, Sr or Ba may come from the subducted

oceanic lithosphère. Lithium is easily mobilized from altered basalts or sediments during

shallow-depth reaction with hydrothermal fluids (Seyfried et al., 1998) or in accretionary

Table 3.2: Sr and Nd isotope compositions for selected

Medicine Lake lavas

sample type 87Sr/86Sr 143Nd/144Nd

82-72a HAOT 0.703402(11) 0.513010(8)95-3 HAOT 0.703457 (14) 0.512952(8)95-4 HAOT 0.703554(11) 0.512917(6)95-5 HAOT 0.703556(16) 0.512930(7)95-6 HAOT 0.703579(13) 0.512933(6)95-9 HAOT 0.703480(10) 0.512986(9)1665m LB 0.703762(13) 0.512860(6)1672m LB 0.703626(13) 0.512921 (5)1399m BA 0.703646(12) 0.512891 (6)79-24E BA 0.703780(11) 0.512845(7)

Numbers in parentheses arc 2o~ errors and refer to the last two

digits for Sr and the last digit for Nd.

HAOT - high-alumina olivine tholeiite, LB - Lake Basalt, BA -

basaltic andésite

47


prisms (You et al., 1995). However, during continuing subduction and change in P-T

conditions, the remaining Li is more strongly bound in the slab, presumably in mineral

phases such as omphacite, Ti-clinohumite and olivine (ZACK et al., 2003; Scambelluri et

al., 2004), with relatively high Li abundances, generally exceeding 10 ppm. Recently,

variable 57Li values have been reported for a suite of Mariana forearc serpentinites, the data

scattering between -6.1%o and +10.3%o (Benton et al., 2004). This complicates estimates of

the exact Li isotope composition of slab-derived fluids.

Strontium, Nd and Pb isotope compositions provide evidence that at least two fluid-rich

geochemical reservoirs were involved in the generation of Mt. Shasta lavas, one with

MORB-likc 87Sr/86Sr (-0.7028) and sNd and a second with high 87Sr/86Sr (-0.7037) and low

8Nd (GROVE ct al., 2002). The Ô7Li of basaltic andésites from Mt. Shasta varies between

+4.0%o and +6.4%o. Both fluid-rich reservoirs exhibit high 57Li. This is consistent with an

origin from variably altered mid-ocean ridge basalts (Tomascak, 2004) which are Li-

enriched and may show Sr isotope variation. The high- Sr/ 6Sr fluid reservoir is unlikely to

reflect detrital sediments while the upper continental crustal materials seem to be isotopically

light, averaging S7Li of ~0%o (TENG et al., 2004). A different fluid source with 57Li <0.9%o is

indicated by BA lavas and Lake Basalt lavas from the Medicine Lake region. However, all

BA lavas and PMA have low Rb/Nb (<3.7) and intermediate Sr/Y (16-73) and are enriched

in Li over similarly compatible elements like Y or Ho (Fig. 3.2).

Lithium isotopes provide a tracer of near surface fractionated materials. Therefore, the

most obvious explanation of the difference between Mt. Shasta and Medicine Lake lavas is

distinct mantle components above the Cascadia subduction zone. However, there are several

lines of evidence that this model is incorrect or at least greatly simplified.

First, the BA lavas and Lake Basalt lavas from the Medicine Lake region have Sr/ Sr

= 0.7036 and 0.7038 respectively, and cannot be distinguished from Mt. Shasta BA lavas

with the "sediment signature" (87Sr/86Sr = 0.7037 to 0.7039; Table 3.2 and GROVE et ai.,

2002). The latter indicates that the hydrous Mt. Shasta and Medicine Lake lavas were derived

from similar sources despite distinct Li isotope compositions. Second, if ö Li is behaving like

a radiogenic tracer for crustal material in the Cascadia subduction zone, then correlations

with other fluid mobile elements or element ratios, e.g. Rb, Ba or Sr isotopes, would be

expected. However, Li isotopes do not correlate with any of these (Fig. 3.3).

A Hkcly explanation for the difference between Mt. Shasta and Medicine Lake lavas is

that Li isotopic effects are dominated by fractionation during subduction of the oceanic

48


lithosphère. Dehydration of a subducting plate extracts heavy Li into the fluid and leaves

residues that are depleted in 7Li. Lithium isotope fractionation during dehydration of oceanic

crust is evident from fluids of the Costa Rica décollement zone (Chan and Kastner, 2000)

and serpentine diapirs, both with 57Li of about +20%o. ô7Li values of dehydration residues

change during prograde metamorphism with increasing degree of dehydration. Eclogites

from Trescolmen (Zack et al., 2003) and an eclogite sample from Oberkotzau, Germany

(Magna et al., 2004b) have negative ö7Li values. Both provide evidence for a 7Li-depleted

eclogitic residue in subduction zones. The pathways of Li depletion and isotope fractionation

12

V10

V

J28

Z 7

S 6 VA A

(£ J7V

V4

*0 A

2:

V

^

05 r i n i r I P n i n f t r * i 1

V

0.4 -

V

.Q 0.3A

? A AV

30.2 V

VA

V o V

A^

A

0.1o V

O

5.0

4.0

V

£3.0

_j2.0

AV

A

1.0 : ^ o VV

A A ^

*v

o

AO

0.0 ' . < » ' ' < * > J—

1.0 1.0 2.0 3.0 4.0 5.0 6 0 7.0 8.0

Ô7LÎ (%o)

Fig. 3,3; Trace element ratios vs. S7Li values in Mt. Shasta lavas. An apparent scatter of data with no

systematic behavior among diverse lava types reflects Li isotopefractionation decoupledfrom release ofafluidphase during dehydration. Closed diamonds — high-alumina olivine tholeiites; grey diamonds - basaltic

andésites; open diamond — primitive magnesian andésite; closed triangles — andésites; reversed open triangles- dacites.

49


Table 3.3: Trace element concentrations in Mt. Shasta and Medicine Lake lavas

sample type Rb Cs Sr Ba Se Y Zr Hf Nb Ni Co Pb Th U

Mt. Shasta

85-38 HAOT 4.9 0.10 169 117a

26 20.3 62 2.0 2.8 42 57 1.3 0.77 0.23

85-51 HAOT 3.5 0.10 228 123a

31 23.7 73 2.2 2.8 52 60 1.4 0.63 0.17

82-94a BA 4.9 0.27 734 258" 30 13.8 83 1.9 4.5 138 37 5.0 2.0 0.50

85-44 BA 3.1 0.07 253 108a

21 12.2 47 1.5 2.1 247 56 1.2 0.46 0.17

95-15 BA 5.6 0.14 611 325a

22 13.1 82 1.9 3.8 245 42 3.9 1.9 0.49

95-17 BA 1.2 0.03 364 53" 25 18.4 70 1.6 0.9 118 40 0.79 0.33 0.10

85-41b PMA 12.3 0.59 741 167a 19 10.1 83 2.7 3.5 128 64 3.5 1.9 0.61

82-94d A 17.6 0.58 1254 338b

11 11.1 125 2.9 3.0 176 15 6.2 3.1 0.97

85-55 A 21.0 1.18 1023 561b

20 13.8 137 3.7 4.4 242 40 7.3 3.8 1.20

85-3 A 16.0 0.52 1170 291b

13 11.6 110 2.8 4.3 236 32 4.2 2.1 0.63

85-48b A 13.4 0.27 1509 202h

12 9.7 94 2.8 3.6 187 30 3.7 2.1 0.56

85-59 A 37.3 1.97 1030 653b

14 11.8 160 4.1 7.6 255 32 11 6.1 1.80

82-85 A 40.7 2.49 914 609b

12 11.2 146 3.2 6.8 223 17 8.3 5.9 1.80

82-92a D 57.3 4.72 498 537b

8 10.0 101 2.6 5.4 123 8 10 6.5 2 50

82-95 D 69.0 2.71 492 488b

13 14.5 131 3.2 7.3 182 15 9.5 8.7 3.00

82-91b D 18.5 0.42 1022 275b

11 9.1 99 2.3 3.6 195 15 3.8 2.5 0.78

96-1 D 18.7 0.86 1098 264b

11 9.4 107 2.5 3.2 182 15 5.2 2.1 0.64

97-7 D 18.0 0.80 1051 286" 11 10.9 115 2.5 3.9 199 16 5.0 2.3 0.69

83-50 D 8.2 0.45 440 191"

22 14.3 73 1.6 3.9 341 31 2.7 1.0 0.34

83-58 D 17.4 0.84 1174 260b

11 9.2 106 3.0 3.6 186 39 5.9 2.1 0.63

95-13 D 12.8 0.62 1176 194b

10 7.3 84 2.0 1.8 174 14 3.8 1.7 0.51

Medicine Luke

82-72a HAOT 1.0 0.017 164 24" 26 17.1 41 1.5 1.4 494 75 0.36 0.21 0.08

95-3 IIAOT 11.1 0.450 473 389a

31 28.1 137 2.9 6.3 456 43 3 7 1 6 0.50

95-4 HAOT 4.4 0.190 358 240" 27 20.3 95 2.1 6.9 98 39 2.9 1.3 0.42

95-5 HAOT 2.2 0.065 390 94" 32 21.9 82 1.6 6.7 123 47 6.0 0.37 0.08

95-6 HAOT 3.1 0.043 370 127" 25 22.6 110 2.2 10.7 133 45 0.97 0.53 0.17

95-9 HAOT 1.3 0.034 287 98a

25 21.1 61 1.4 2.3 187 52 0.85 0.32 0.09

1665m LB 5.5 0.297 538 242" 20 14.3 69 1.6 2.8 60 28 3.0 1.3 0.35

1672m LB 6.0 0.350 585 241"

23 17.2 77 1.7 3.0 41 30 2.2 0.79 0.25

1399m BA 9.0 0.395 357 194" 26 21.7 97 2.1 4.3 77 37 2.7 1.1 0.37

79-24E BA 16.3 0.810 445 241"

23 16.8 88 1.9 4.4 108 35 5.0 1.6 0.51

Reference standards

JB-2 6.1 0.67 199 235a 49 23.0 47 1.3 0.67 19.7 38 5.3 0.31 0.17

BHVO-2 7.7 0.09 407 138a 30 24.8 165 4.1 19.7 126 46 1.5 1.34 0.44

AGV-2 71 1.31 712 1142b 14 19.0 225 4.5 15.5 20 16 13 6.7 1.76

All trace clement concentrations arc given in ppm.

HAOT - high-alumina olivine tholciitc, BA basaltic andésite, PMA - primitive magnesian andésite, A - andésite, D - dacite,

LB - Lake Basalt

J

corrected down for 15%

corrected down for 10%

50


Tabic 3.4: Rare earth clement concentrations in Mt. Shasta and Medicine Lake lavas

sample type La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Mt.Shasta

85-38 HAOT 4.60 11.2 1.68 8.20 2.67 1.01 2.90 0.57 3.80 0.82 2.53 0.36 2.37 0.37

85-51 IIAOT 4.74 11.9 1.90 9.30 3.22 1.26 3.40 0.62 4.24 0.90 2.81 0.41 2.66 0.38

82-94a BA 12.0 26.0 3.53 14.3 3.55 1.14 3.03 0.46 2.60 0.52 1.47 0.23 1.47 0.21

85-44 BA 3.78 9.00 1.34 5.90 1.98 0.69 2.00 0.36 2.20 0.47 1.50 0.22 1.36 0.20

95-15 BA 16.0 35.7 4.67 18.9 4.11 1.25 3.29 0.45 2.60 0.50 1.53 0.22 1.44 0.22

95-17 BA 3.50 9.40 1.56 8.10 2.62 1.05 2.90 0.54 3.50 0.73 2.18 0.33 2.21 0.30

85-41b PMA 10.3 23.1 3.06 12.3 2.70 0.92 2.30 0.34 1.88 0.39 1.14 0.16 1.05 0.16

82-94d A 18.1 39.7 5.40 20.8 4.34 1.46 3.40 0.41 2.30 0.40 1.22 0.16 1.06 0.16

85-55 A 22.8 51.1 6.81 26.7 5.73 1.75 4.30 0.53 2.76 0.52 1.58 0.21 1.38 0.20

85-3 A 15.2 33.1 4.30 17.1 3.89 1 29 3.10 0.41 2.29 0.43 1.34 0.18 1.22 0.17

85-48b A 15.6 36.1 4.77 17.9 3.47 1.20 2.95 0.37 1.93 0.37 1.11 0.14 0.97 0.14

85-59 A 28.6 57.7 7.21 27.5 5.94 1.71 4.05 0.49 2.50 0.44 1.30 0.18 1.18 0.17

82-85 A 25.1 50.9 6.30 23.8 5.08 1.52 3.79 0.47 2.34 0.45 1.26 0.17 1.14 0.16

82-92a D 15.1 29.2 3.52 13.1 3.38 106 2.51 0.37 1.98 0.38 1.13 0.16 1.06 0.17

82-95 D 19.1 38.6 4.79 18.1 4.41 1.18 3.60 0.49 2.81 0.55 1.60 0.23 1.53 0.23

82-91b D 12.8 27.2 3.50 13.9 3.21 1 05 2.56 0.35 1.92 0.37 1.06 0.15 0.94 0.14

96-1 D 12.9 27.6 3.62 14.3 3.42 1.11 2.80 0.35 1 85 0.37 1.07 0.16 0.96 0.15

97-7 D 14.3 30.9 3.98 16.0 3.63 1.18 2.90 0.40 2.12 0.42 1.20 0.17 1.10 0.16

83-50 D 7.00 15.9 2.25 10.1 2.88 1.09 2.70 0.46 2.70 0.56 1.70 0.23 1.55 0.23

83-58 D 13.3 29.0 3.75 15.0 3.40 1.11 2.70 0.35 1 90 0.36 0.99 0.14 0.93 0.15

95-13 D 9.74 21.3 2.87 11.7 2.66 0.95 2.20 0.27 1.57 0.29 0.86 0.11 0.71 0.12

Medicine Lake

82-72a HAOT 1.58 4.60 0.82 4.39 1.80 0 76 2.13 0.45 3 11 0.68 2.10 0.30 2.10 0.30

95-3 HAOT 11.1 25.9 3.74 16.4 4.94 1.90 4.90 0.80 5 10 1.04 3.10 0.46 3.00 0.46

95-4 HAOT 8.80 20.0 2.77 12.2 3.49 1.29 3.30 0.60 3 84 0.78 2.39 0.35 2.39 0.35

95-5 HAOT 5.95 15.1 2.23 10.7 3.21 1.25 3.60 0.61 4.10 0.86 2.60 0.37 2.51 0.38

95-6 HAOT 8.60 21.2 3.03 13.7 3.93 1.49 4.00 0.68 4.40 0.91 2.70 0.37 2.59 0.38

95-9 HAOT 3.58 9.39 1.48 7.60 2.51 1.08 3.60 0.53 3.77 0.84 2 47 0.37 2.40 0.36

1665m LB 7.10 15.8 2.14 9.50 2.75 1.10 2.70 0.44 2.70 0.56 1.67 0.23 1.61 0.27

1672m LB 6.85 15.8 2.30 10.5 3.18 1.40 3.20 0.52 3.20 0.64 2.02 0.27 1.90 0.27

1399m BA 7.90 18.7 2.72 12.2 3.65 1 32 3.70 0.64 4.10 0.85 2.55 0.37 2.48 0.36

79-24L BA 9.30 20.4 2.80 11.8 3.37 1.20 3.20 0.51 3.16 0.67 1.90 0.28 1.86 0.27

Reference standards

JB-2 2.35 6.88 1.24 6.4 2.40 0.97 3.12 0.61 3.95 0.88 2.60 0.41 2.55 0.39

BHVO-2 15.5 38.6 5.56 24.9 6.23 2.29 6.4 1.08 5.54 1.06 2.64 0.42 2.12 0.38

AGV-2 39.4 71.7 8.6 31.3 5.8 2.04 5.6 0.69 3.6 0.71 1.95 0.26 1.69 0.27

All rare earth element concentrations are given in ppm.

HAOT - high-alumina olivine tholciitc, BA - basaltic andésite, PMA primitive magnesian andésite, A - andésite, D - dacite, LB -

Lake Basalt

51


are not expected to be constant at all pressure and temperature conditions in subduction

zones. However, dehydration at low temperatures will fractionate 7Li into the fluids and

produce a 6Li-enriched residue. Therefore, fluids that derived from the slab below ~75 km

are expected to be isotopically lighter than unaltered MORE as also recently experimentally

verified (Wunder et al., 2006).

3.6.3 Andésites and dacites

Andésites (A) and dacites (D) from Mt. Shasta have been modeled as products of

multiple recharge of magmatic reservoirs by parental H20-rich PMA magmas (e.g. 85-41)

accompanied by fractional crystallization of olivine + pyroxenes + amphibole + plagioclase

throughout (or within) the continental crust. Possibly some BA lavas were also introduced

into the system (Grove et al., 2005) and served as parent magmas. Experimental work

provides evidence for high pre-eruptive H2O contents (up to 10 wt.%) in some evolved

magmas (Grove et al., 2003). Similar water contents have been observed in northern

Cascadia (Blijndy and Cashman, 2001). Although crustal rocks are present in the area

(Zucca et al., 1986), Sr and Nd isotope data do not indicate contamination by continental

crust. This is also endorsed by U-Th-Ra disequilibria for young volcanic rocks (Volpe,

1992). Andésites and dacites are therefore derived from a H20-rich PMA-magma by

fractional crystallization and mixing with related, but more evolved magmas (Grove et al.,

2005). This mode of origin differs from that proposed for the generation of andesitic, dacitic

and rhyodacitic lavas in the High Cascades where melting of amphibolitic and granulitic

rocks at various crustal levels without significant mafic contributions has been suggested

(Conrey et al., 2001). Average continental crust and Trinity peridotite as contaminants have

been excluded based on major and trace elements (Grove et al., 2005). Continental crust has

been advocated formerly (Newman et al., 1986). The likely contaminant has been identified

as highly evolved differentiated residual liquids from long-lived magmatic systems that

underwent fractional crystallization, recharge of magma and mixing (Grove et al., 2005).

Andésites and dacites show a similar range of trace element characteristics with, for example,

depletion in Nb and Ta, accompanied by enrichment in Ba, Pb, Sr and Li (Fig. 3.4). The S7Li

range measured for andésites and dacites is identical to the Li isotope composition of

tholeiites and basaltic andésites. This precludes a voluminous contamination of andésites and

dacites by old or upper crustal material and restricts its volume to less than 5% (andésites: Li

= 10 ppm, Ô7Li = 7%o, Sr = 1200 ppm, 87Sr/86Sr = 0.7028, continental crust: S7Li = 0%o, Li =

52


35 ppm (Teno et al., 2004), 87Sr/86Sr = 0.740, Sr = 276 ppm, CC2 composition in Table 5

from Hart, 1985). Although the upper continental crust appears to have ö7Li ~0%o (Teno et

al., 2004), the Sr and Nd isotope compositions will have evolved in a manner that is distinct

from those of the mantle and basaltic systems. Evolved calc-alkaline lavas of Panama arc

show a considerable range in ö7Li from +3.9%o to +11.2%o (Old Group in Tomascak et al.,

2000). This has been ascribed to release of isotopically heavy Li from previously subduction-

fertilized mantle in an environment where ridge subduction was being initiated. In contrast,

adakitic lavas from the Panama arc exhibit a much more limited range in 87Li from +1.4%o to

+4.2%o and may represent slab melts. Such an interpretation is broadly consistent with the

model presented in this study.

1000 c

100

CÛce

o

TO>CO

10

0.1 -I—I—I—I—I—I- -I—I—I—\- _4 1 1 ( ( 1 1 » I -I i (-

Cs Rb Ba Th U Nb Ta La Ce Pb Pr Sr Nd Sm Zr Hf Eu Gd Tb Dy Li Y Ho Er Tm Yb Lu

Fig. 3.4: Trace element patterns for andésites and dacitesfrom Mt. Shasta. Andésites (light grey) and dacites

(while) display nearly identical patterns with only subtle differences. Andésites are more homogeneous than

dacites with less variable trace element ratios, e.g. (La/Sm)N or Sr/Y, however are they generally

indistinguishablefrom dacites in N-MORB-normalizedpatterns.

A subset of four Li-rich samples (85-59, 82-85, 82-92a, 82-95) shows distinct 87Sr/86Sr

ratios compared to the rest of andesites-dacites group. These samples carry a Sr isotope

signature of a sediment-derived fluid-rich component (Grove et al., 2005). Lithium

53


abundances between 18 and 35 ppm also provide evidence for varying degrees of

contamination by rhyolitic rocks. However, their Ô Li values scatter substantially and this

may reflect heterogeneity in the rhyolitic contaminant. It is possible that variability in the

57Li is a characteristic of this rhyolitic residuum (Grove et al., 2005).

We suggest that the mineralogy of the subducted slab does not have a large effect on

the Li isotope composition of andésites and dacites, as these are products of multiple mixing

of parental H20-rich magmas with fractionally crystallized lavas at crustal depths (Grove et

al., 2003, 2005).

3.7 Li isotope fractionation in the southern Cascadia zone

The high temperatures of melting in the spinel peridotite source regions beneath Mt.

Shasta and Medicine Lake (1300-1450°C; ELKPNS TANTON et al., 2001) rule out isotope

fractionation of Li isotope compositions from the parental peridotite. Fluids added to the

source region of HAOTs from Mt. Shasta are probably responsible for the variations in S7Li

and represent a source material that is different from that provided by the subduction-derived

fluid-rich components that were added when the hydrous basalts and basaltic andésites from

Mt. Shasta and Medicine Lake formed. The BA, PMA and Lake Basalt lavas display a

decrease in S Li away from the arc trench (Fig. 3.5). This may reflect heterogeneity of the

mantle wedge and continuous Li isotope fractionation as fluids are released from the slab at

progressively greater depths (Zack et al., 2003).

It is a reasonable generally accepted scenario that fluids with heavy Li are injected from

the subducted slab into the viscously coupled overlying mantle at shallow depth (Fig. 3.6).

This hydrous mantle segment (probably serpentinite) is dragged down by the subducting

plate (Elliott et ai., 2004) and metamorphosed. Just below Mt. Shasta the mantle segment

begins to dehydrate as a consequence of the breakdown of serpentinite and / or chlorite at

temperatures >600°C (ULMER and TROMMSDORFF, 1995). The high temperatures of these

dehydration reactions would be expected to lead to only a small fractionation of Li isotopes

between the source serpentinite and the fluid. It is therefore unlikely that this mantle segment

after losing its Li has very low 87Li. The low Ô7Li values for Mt. Shasta lavas probably were

added during the subsequent vapor-saturated melting that occurred as the fluids ascended and

encountered higher temperatures in the mantle wedge, i.e. the low-57Li lavas from Mt. Shasta

included a mantle wedge component with relatively "normal" Li. That is, the Li is diluted

54


a§

°.

itoCM

9.0

80

70

60 -

5,0 -

40

30

2,0

1 0

00

7.0

6,0

50

40

3.0

2.0

1 0 [

slab fluids

<p.p

MANTLE

slab fluids

or melts M

f ."i

B

( t f,

;L -^ . ( ,

(l

_t tm

, jii f

200 225 250 275 300 325

distance from trench (km)

350

Fig. 3.5: Distance from the arc trench vs. ifLi values (panel A) and pre-eruptive water contents (panel B),

respectively. Distinct hi isotope signatures are found dependent on distance from trench. A general trend of

decreasing üfLi with increasing distance from the trench (and conversely, depth of Wadati-Benioff zone) is

similar to that observed for the lzu arc (MORIGUTI and NAKAMURA, 1998b) and has not been found in the

northern Cascadia arc (LEEMAN et al, 2004) or Kurile and Sunda arcs (TOMASCAK et al., 2002). This has been

interpreted in terms of a lack of Li isotopefractionation because Li is relatively strongly retained in subducted

slab (Tomascak et al., 2002; hEEMAN et al, 2004). Similar conclusions were drawn by Moriguti el al. (2004)

for northeastern Japan. Diamonds - Mt. Shasta lavas (closed - HAOTs; grey - BAs; open - PMA); squares -

Medicine Lake lavas (closed - HAOTs; grey BAs; open- Lake Basalts). The mantle composition shown as a

bar is defined as tfLi = 4.1 ±0.5%t> as inferred from analyses ofspinel Iherzolites and their olivines (MAGNA et

al, 2006b). In panel B, pre-eruptive water contents can be seen to decrease in "flux-melted" magmas as a

function of distance from the trench. Significant water contents are recognized in basaltic andésites and Lake

Basalt lavas from Medicine Lake.

55


Mt Shasta Medicine

Fig. 3 6: Schematic figure ofevolution of hi isotopes in the southern Cascadia subduction zone. Altered oceanic

crust (AOC) provides source of heavy hi isotope input into the subduction zone due to 7Li being preferentially

adsorbed in secondary clay minerals. This Li is easily released during low-T events just after passing the arc

trench. Overlying serpentinite diapirs become isotopically heavy. Progressive hi depletion in successive steps

ofslab dehydration and/or slab melting is reflected in lighter hi isotope composition ofBA lavasfurther away

from the trench. HAOTs from both Mt. Shasta (with a source locatedjust beneath the base of the continental

crust, depicted as a vertically lined oval) and Medicine hake (horizontally patterned oval indicating a deeper

source of these HAOTs) represent near-anhydrous melting of spinet peridotites. Slightly different cfLi values

mirror heterogeneity ofthe mantle segments beneath the two volcanoes.

with a mantle component with Ô7Li ~ 4%o. Therefore, S7Li variation observed in Mt. Shasta

lavas reflects both: different proportions of Li derived from the slab and Li isotope

heterogeneity in the mantle wedge. Further evidence of a heterogeneous mantle wedge is

provided by anhydrous volcanics from the Mt. Shasta-Medicine Lake region that exhibit a

large range of ö7Li values. The complexity of this heterogeneous mixing process may explain

why 57Li does not correlate with fluid mobile trace elements and strontium isotopes.

The H20-rich lavas from Medicine Lake (which also have a large contribution from the

slab in their incompatible element abundances) arc uniformly low in S Li. These lavas are

likely generated by vapor-saturated melting in the mantle wedge as the slab-derived fluids

ascend. That the metamorphosed subducted slab contains reasonable amounts of Li has been

recently shown (SCAMBELLURI et al., 2004). In particular mineral phases such as staurolite

56


(DUTROW et al., 1986), serpentine, phengite, chlorite and Ti-clinohumite can hold high Li

abundances. The breakdown of those minerals that are H20-rich such as lawsonite or

phengite may occur in subducted metamorphosed oceanic crust just beneath Medicine Lake.

In addition, breakdown of serpentine in the subducted oceanic slab can be an effective source

of Li at Medicine Lake. Such fluids will be isotopically light because heavy Li is lost at

shallow depths causing metasomatic overprinting of the overlying mantle wedge with

predominantly heavy Li. Thus, progressive dehydration reactions with the accompanying

higher temperatures deeper in the downgoing slab causes the distinct 57Li observed in Mt.

Shasta and Medicine Lake basaltic andésites.

Studies of eclogites from Trescolmen, Switzerland (Zack et al., 2003) and Oberkotzau,

Germany (Magna et al., 2004b) have revealed negative ô7Li values. Fluids derived from

such an eclogitic source are Li rich and expected to be light (Zack et al., 2003) because the

fractionation between rock and fluid is small at high temperatures. Therefore, the data are

consistent with the Wadati-Benioff zone being at greater depth and higher temperature below

Medicine Lake than Mt. Shasta. The isotope fractionation in subduction zones would explain

why Li isotopes do not correlate with fluid mobile trace elements and strontium isotopes.

Such a decoupling has also been described for oceanic arcs (Tomascak et al., 2002; Chan et

al., 2002b) and seems to be a characteristic of most arc magmas.

3.8 Comparison with other arcs

Lithium isotopes have been studied in several arcs. However, only the Izu arc

(Moriguti and Nakamura, 1998b) and the southern Cascades (this study) show a decrease

in ö7Li values with increasing distances from the trench (or depth of the slab). These two arcs

deviate from the other localities by abundant proportions of basalts (Fig. 3.7). For the Izu arc

a mixing model has been proposed in which mantle wedge Li mixes with slab-derived Li

(Moriguti and Nakamura, 1998b). The fluids have a signature mainly of altered oceanic

crust. Such a mixing model is consistent with a study on trace elements and radiogenic

isotopes by HöchstAEDTER et al. (2001). However, the model (Moriguti and Nakamura,

1998b) relies on a mantle end member that has Ô7Li ~ 0 to i-2%0 which is inconsistent with

the vast majority of oceanic basalts and unmetasomatised peridotite xenoliths. Therefore,

continuous Li isotope fractionation in the slab during subduction and an interpretation of the

light end member being a slab signature seems to provide a more satisfactory explanation of

57


#

0.0

Mt. Shasta

Medicine UK«

9100 150 200 250 300

Distance from trench (km)

350

6.0

£4,0

0.0

122.5

' *"' ifUKB

"'1 mmÊ 1mmÜ- Group 1

~~*""\

m* **

\\

Group II

''' TRENCH]

122.0 121.5 121.0

W Latitude

120 5

20

0.0

AIMuusl.

C

\\ \

v,f)

100 150 200 250

WBZ (km)

300 350

200 250

WBZ (km)

350

80

6.0

*r- 4.0 i*_i \

4g predominantlyJ basalte andésites

/*"' ".—

ioo 200 250 300 350

WBZ (km)

s.o

6.0

rf

«5

'

*l

F

I*

t*

* i

^4

I-_ --*)

150 200 250 300

WBZ (km)

350

Fig. 3.7: Comparison ofLi isotope profiles ofvarious arcs, hithium isotopes fractionate strongly with increasing

distances (or increasing depth of Wadati-Benioffzone) from the arc trench in northern California Cascades (A)

and in Izu arc (C) with predominantly basaltic volcanism (GiLL, 1981). On the other hand, arcs with prevalent

basalt andesitic and andesitic lavas show across-arc patterns of a7hi values that are similar among these latter

settings, i.e. restricted invariant range of ö7hi values. A - hi isotope data from Mt. Shasta (this study), B -

southern Washington Cascades (see htiEMAN et al. (2004) for distinctions between Group I and Group II), C -

Izu arc (MORIGUTI and NAKAMURA, 1998b), D - northeastern Japan arc (MORIGUTI et ai, 2004), E - Kurile arc

(TOMASCAK et al, 2002), F - Sunda arc (TOMASCAK et ai, 2002). Symbols are plotted irrespective of the

petrologic evaluation and major element chemistry of corresponding samples. Insets: basalts are in black,

basaltic and acid andésites are in dark grey and dacites are in white. Inset 7C: basaltic andésites are in dark

grey and acid andésites are in light grey. Note the variations ofrelative proportions of major magma types in

the insets (Gill, 1981). WBZ-depth of Wadati-Benioffzone.

58


the data for the Izu arc and the southern Cascades.

For andesitic lavas from northeastern Japan a slightly smaller range of S7Li values from

+2.0 to +4.4%o is reported but no systematic variation across the arc has been detected

(Moriguti et al., 2004). The subducted slab in northeastern Japan sinks with a dip angle

of~30° into the mantle whereas the dip angle for the Izu arc is -50°. The different geometry

and a higher temperature at the top of subducting slab in the northeastern Japan (MORIGUTI et

al., 2004) may result in a dehydration of the slab at very shallow, crustal levels which are not

sampled by the volcanism. It may be that in northern Japan only Li from a 7Li-depleted slab

is incorporated in subduction zone lavas. Therefore, very heavy Li isotope ratios of up to 7%o

are missing. Indeed, Li-Pb isotope modeling shows that Ô7Li of slab-derived fluid and mantle

wedge end members are potentially very similar resulting in lack of Li isotope fractionation

during continuous subduction. A convex-upward trend of Li/Y was observed across the

northeastern Japan arc and was explained by the possible influence of lawsonite breakdown

(Moriguti et al, 2004). This could be a consequence of shallower subduction and thus,

different P-T conditions for northeastern Japan compared to Izu arc. Whether this could

influence the 57Li path across the arc is presently unclear. A concave-upward profile of 5nB

across the northeastern Japan arc has been explained by involving tourmaline as a minor

mineral phase that carries a substantial portion of the boron in the subducted slab (Moriguti

et al., 2004).

Across the southern Washington Cascadia 57Li values of tholeiites to basaltic andésites

range from +2.5 to +4.2%o and are randomly distributed across the arc (LEEMAN et ai., 2004)

which is distinct from the Li isotope patterns observed in lavas from this study. The fluid-

mobile element abundances decrease from the forearc lavas towards the backarc. A

corresponding effect on Li isotopes is not observed but B isotopes change systematically

across the arc. It has been suggested that this is because Li is retained in the subducted slab

more efficiently than B. Such an inference is in good agreement with previous observations

made for other arc settings (Tomascak et al., 2002; Moriguti et al., 2004). Hence, the Li

isotope signatures in these lavas may represent the composition of the mantle wedge because

there has been only a small input of isotopically distinct fluids from the slab (Leeman et al.,

2004).

Lithium isotope compositions of basaltic andésites and andésites in the Kurile arc range

from +2.4 to +4.9%o and are randomly distributed across the arc (Tomascak et al., 2002). In

contrast, B isotope ratios vary systematically with the distance to the trench (Ishikawa and

59


TERA, 1997). It has been suggested that this is either due to very low Li contents in fluids or

Li has been removed from fluids by interaction with minerals in the mantle wedge

(Tomascak et al., 2002). Apart from one sample, tholeiites and high-K alkali lavas of the

Sunda arc display a narrow range of 57Li values from +2.1 to +5.1%o (Tomascak et al.,

2002). This range is as observed in lavas from southern Washington Cascades, Kurile and

northeastern Japan arc. Mixing of isotopically heterogeneous mantle Li and a homogeneous

slab-derived Li in fluids or melts is proposed for the observed B/Be and Sr-Nd-Pb isotope

systematics in Sunda arc (EDWARDS et al., 1993). Possible explanations for the distinct Li

isotope patterns in arcs involve cither Li-poor slab fluids or removal of Li from the fluids by

interaction in the mantle wedge (Tomascak et al., 2002). In the first case, Li would be

retained in the slab. Indeed, some evidence is found for such behaviour of Li during

subduction (sec sections 3.6.2 and 3.7, respectively). This Li would not contribute

significantly to the Li isotope budget of arc volcanics (TOMASCAK et al., 2002). In the latter

case, Li is sequestered in rims of minerals in the mantle wedge. Decoupling of Li and B

isotope patterns indicates a chromatographic separation of Li isotopes as fluids percolate

through the mantle wedge. This may erase the slab signature of a moderately incompatible

element like Li (Tomascak et al., 2002) but not that of highly incompatible element like B.

In conclusion, the dip angle and the temperatures in subduction zones will affect the

extent of dehydration in the subducting slab. Most 7Li may be lost into the crust where

subduction is shallow. In only two locations do Li isotopes show an inverse correlation with

depth of Wadati-Benioff zone (or distance from trench). In these arcs, basalts are the

dominant lava type. As discussed above, arcs dominated by basalts provide evidence that Li

isotopes continuously fractionate in the slab during subduction. In contrast, Pb isotopes and

other radiogenic tracers provide evidence that ascending fluids or melts interact and mix with

distinct components in the mantle wedge.

3.9 Conclusions

1) High-alumina olivine tholeiites are near-anhydrous melts of spinel peridotites. Lithium

isotopes are not fractionated during melting and most likely reflect Li isotope

composition of the uppermost mantle beneath Mt. Shasta.

60


2) Lithium in the hydrous BA lavas in the Mt. Shasta region is heavy whereas BAs and

basalts from Medicine Lake have isotopically light Li providing evidence for isotope

fractionation during subduction of the lithospheric slab.

3) Basaltic andésites exhibit the same Ô7Li range as andésites and dacites from Mt. Shasta

reflecting the composition of the mantle source. Thus, our data provide further evidence

that solely magmatic differentiation processes are not fractionating Li isotopes.

Acknowledgement - We are indebted to Irene Ivanov-Bucher for her help with the sample

preparation, Darrell Harrison for his careful reading of early versions of the manuscript, Felix

Oberli for technical assistance on Nu 1700 and Marcus Gutjahr for his help with Sr and Nd

isotope analysis. Lui-Heung Chan, James Gill, Bill Leeman and Paul Tomascak are

acknowledged for their valuable information. We are grateful to Paul Tomascak for detailed

and insightful review, Tim Elliott for offline comments to the manuscript and David Price for

editorial handling. ETH and Swiss Nationalfonds (SNF) supported this study.

61


Supplementary figure

i22°w Oregon 121"W

Supplementary figure SI: Schematic map ofsample locations ofMt. Shasta and Medicine hake lavas analyzed

for lithium concentrations and isotope compositions. Solid Ones with numbers represent roads, dashed lines

represent contour lines with elevation in feet.

62

Chapter 4 North Atlantic

Chapter 4

Combined Li and He isotopes in Iceland and

Jan Mayen lavas and the nature of the North

Atlantic mantle

*

to be submitted for publication in Geochimica et Cosmochimica Acta as Magna T., Halliday

A.N., Wiechert U., Stuart F.M. and Harrison D.: Combined Li and He isotopes in Iceland and

Jan Mayen lavas and the nature of the North Atlantic mantle

63


4.1 Abstract

We present lithium (Li) isotope data as well as supporting isotope and trace element

compositions for two suites of basaltic volcanic rocks from Iceland and Jan Mayen. Basaltic

glasses from the Hengill fissure system are depleted in 180, U, Th and Li, enriched in Ba, Nb

and Eu and have variable Sr/Nd. They can be divided into enriched (LaN/Sm^Ll) and

depleted (LaN/SmN<0.52) lavas. The latter have superchondritic Sr/Nd and are anomalously

depleted in Pb. The two groups define a collinear trend when 57Li (+3.8 to t6.9%o) is plotted

against 3He/4He (12 to 20 Ra). There is no well-defined correlation between the Li or He

isotope compositions and other isotope or trace element parameters. The isotope variability in

the Iceland suite contrasts strongly with the homogeneous Li (87Li = +3.9 to +4.7%o) and He

(3He/4He = 5-7 Ra) isotope compositions of basalts from Jan Mayen. These lavas show

large degrees of enrichment in incompatible trace elements (LaN/SmN>3.2) with patterns

similar to those of the Hengill Enriched Group, the only striking difference being a large

positive Pb anomaly.

The linear relationship between Li and He isotope systematics in Hengill would indicate

that the end members had similar Li/He. The highest 3He/4He of 20Ra is associated with

"normal" mantle 57Li of +3.8%o. There is no reason why this should be the case unless it

represents an end member composition. Therefore, despite the light oxygen isotope

compositions and Eu and Nb anomalies, the high-3He/4He end member carries no Li isotope

trace of being the product of recycling of altered slab or wedge material. Neither does it

appear to reflect a high-^He/'He volatile-rich lower mantle reservoir else the correlation with

Li isotopes would be non-linear. The Jan Mayen basalts provide evidence of a third source

characterised by EM2-like trace element enrichments and 3He/4He that is lower than in

MORB accompanied by the same normal 87Li as the high^He^He Hengill basalts. Therefore,

MORB and OIB can display a dramatic range in He isotope compositions from 5 to 20 Ra

and large variations in trace element compositions associated with recycled components

whilst exhibiting no sign of surface altered slab or wedge material in the Li isotope

composition. Furthermore, there is no sign of a high-3He/4He reservoir selectively dominating

the variability in He isotope composition. One possible explanation for these data is that the

He isotope compositions reflect small variations in (U+Th)/He in ancient low-180 lower

oceanic crust and upper mantle material. The high 87Li (+6.9%o) component with 3He/ He =

12 RA includes samples that tend to have higher Nb/U, Sr/Nd and Eu anomalies providing

64


evidence of a cumulate enriched source that may also be part of an ancient altered ocean floor

slab.

4.2 Introduction

The role of recycled material in the formation of ocean island basalts (OIB) has been the

focus of considerable debate (Hofmann and White, 1982; Halliday et al., 1995;

HOFMANN, 1997; NlU and O'HARA, 2003). Lithium (Li) isotope fractionation in nature is

mainly the result of low-temperature processes near the Earth's surface (Chan et al., 1992,

2002a). Fractionation of Li isotopes accompanying aqueous alteration of oceanic crust, and

its subsequent dehydration during subduction has the potential to be used to trace subducted

material in the convecting mantle (Elliott et al., 2004; Magna et al., 2006a). Dehydration

leads to an enrichment of 7Li in the overlying mantle wedge with substantial loss of Li from

the subducted slab into the fluid phase (ZACK et al., 2003). However, a significant proportion

of the Li inventory is still retained in the residual slab even after large-degree dehydration

(Scambelluri et al., 2004). Lithium tends to be depleted and isotopically heavy in gabbros

relative to unaltered seafloor basalts (Chan et al. (2002a) and T. Magna, unpublished data).

Less altered portions of recycled crust show ö7Li that is elevated over unaltered basalts (up to

+7.4%0; Nishio et al., 2005) whereas highly altered oceanic crust (up to 114%0; Chan et al.,

1992) is thought to leave a residue that has extremely negative 67Li, as low as -1 l%o, after

dehydration (Zack ct al., 2003). Therefore, the exact estimate of the Li isotope compositions

of the sources feeding ocean island basalts is problematic due to variable ö Li within the

oceanic crust pile and the ill-defined nature of what portion contributes to the recycled

signature of OIBs.

Unaltered mid-ocean ridge basalts (MORB) show a relatively large range of S7Li (-11.5

to +5.6%0; Tomascak et al., in press; Chan et al., 1992; MORIGUTI and Nakamura, 1998b;

ELLIOTT et al., 2004, 2006) with only subtle differences between N-, T- and E-MORB lavas

from different ridge systems. This range is similar to that observed in ocean island basalts,

e.g. Hawaii (+2.5 to +5.7%o; Tomascak et al., 1999b; Chan and Frey, 2003; Kobayashi et

ai., 2004) and indistinguishable from previously published data for Icelandic basalts (+3.1 to

+4.0%o; Ryan and Kyle, 2004). HIMU-OIB lavas from Polynesia yield slightly elevated ö7Li

>5.0%o which has been explained as a result of being derived from recycled altered oceanic

crust that underlies the recycled uppermost crust (Nishio et al., 2005). The range of Li

65


isotope compositions in HIMU and other OIB types (EMI and EM2) remains obscure

(Tomascak, 2004).

Lithium is isotopically heavy in seawater (ö7Li ~ +31%o) and in meteoric water (see

comprehensive review by Tomascak, 2004). Lithium is strongly fractionated in the ocean

floor and during subduction and has the potential to be used for tracing the source

components in ocean island basalts. However, extremely light Li isotope ratios are not

observed in any OIBs (Elliott et al., 2004). Here we report new Li isotope data for basalts

from the Western Rift Zone of Iceland and from Jan Mayen in order to characterise the source

of recycled components in the mantle beneath the North Atlantic Ocean and improve our

understanding of how Li isotopes fractionate during subduction and recycling of oceanic

crust.

4.3 Samples and geologic background

Icelandic volcanism has long been attributed to the presence of a mantle plume, a

hypothesis supported by seismic imaging of low-velocity material extending to ~700 km

depth (Wolfe et al., 1997). The presence of recycled oceanic crust in the Iceland plume has

been proposed on the basis of incompatible trace elements (Chauvel and HÉmond, 2000),

Pb isotopes (Thirlwall et al., 2004) and ,80-depleted crustal components (Skovgaard et

al., 2001). Support for this is provided by trace elements and He-0 isotopes in basalts from

central Iceland that display a large range of 3He/4He (1.6-34.3 Ra) and low 5180 values

(<4.8%0; Macpherson et al., 2005). Several origins for the recycled component have been

proposed:

(i) heterogeneous basaltic oceanic crust (Tiiirlwall et al., 2004),

(ii) a mixture of recycled basaltic oceanic crust and low-ö'sO Icelandic mantle

(Thirlwall et al., 2006),

(iii) both gabbroic and basaltic part of the oceanic crust (CHAUVEL and HÉMOND,

2000; Skovgaard et al., 2001),

(iv) lower oceanic crustal gabbros (Breddam, 2002).

Model (i) has been proposed based on detailed mapping of coupled Pb-Sr-Nd isotopes

in Iceland and on adjacent ridges. The Pb isotope systematics are believed to reflect

preferential hydrothermal addition of U to the Palaeozoic oceanic crust and contemporaneous

depletion of lead during dehydration of the subducted slab (Thirlwall et al., 2004). Model

66


(ii) is predicated on several lines of evidence (Thirlwall et al., 2006). The uniquely low

5lxO of Icelandic basalts could be considered to reflect a contribution of recycled Ordovician

crust with oxygen isotope compositions that are vastly different from modern oceanic crust.

Alternatively, host mantle may be the carrier phase of low 8180 (and, simultaneously, high

3He/4He) signature whereas other geochemical features of enriched Icelandic mantle would be

dominated by incompatible element enriched recycled crust (Thirlwall et al., 2006). The

idea that low ôlsO is generated by oxygen isotope fractionation in the D" layer was discussed

and dismissed by Thirlwall et al. (2006). In model (iii) variable but low volumes of ~3 Gyr

gabbros and basalts are required to account for the mixing relationship observed in e.g.

Theistareykir lavas as modelled for combined O-Os isotopes (Skovgaard et al., 2001). The

source of depleted lavas is thought to involve the gabbroic part of the oceanic crust because

of excess Sr, Ba and Eu, positive (Sr/Nd)N and deficiencies in Hf (and Zr) relative to the

REE. These characteristics closely resemble plagioclase-clinopyroxene-rich gabbros

(Chauvel and HÉMOND, 2000). It has been proposed that some enriched Icelandic basalts

have a HIMU-affinity (Stracke et al., 2003) and could be sourced in Palaeozoic oceanic

crust (Thirlwall et al., 2004). This is inconsistent with other evidence for more ancient

recycled oceanic crust the isotope fingerprint for which develops through long-term storage in

the mantle (Chauvel and Hémond, 2000; Skovgaard et al., 2001). Model (iv) has been

proposed as the explanation for the compositions of depleted Icelandic lavas representative of

high-degrec melts (Breddam, 2002). It is supported by the presence of excess strontium

L(Sr/Nd)N>lJ and a positive correlation between (Sr/Nd)N and 143Nd/l44Nd that discriminates

between enriched and depleted plume components.

The Icelandic basalts studied here are olivine phenocryst- and Cr-diopside megacryst-

bearing olivine tholeiites and picritic glasses from the 15 km long Maelifell-Miöfel fissure

system (MMF), Hengill, approximately 45 km east of Reykjavik (Harrison et al., 1999).

The lavas erupted rapidly without significant crystallisation within the crust (TR0NNES, 1990;

Hansteen, 1991). Oxygen isotope compositions of glasses from MMF show remarkably low

ô180 (GURENKO and CHAUSSIDON, 2002; BuRNARD and HARRISON, 2005) that may reflect

either alteration by 1ow-180 aqueous fluids or recycling of an old recycled oceanic

lithosphère. That MMF glasses escaped significant contamination by Icelandic crust has been

shown by broadly uniform and light boron isotope compositions in glass inclusions entrapped

in olivine with an average 8MB of-1 l%o (Gurenko and Chaussidon, 1997) matching that

67


of "primitive mantle". The magmas have been explained in terms of <20% fractional melting

at depths of between 75 and 30 km (Gurenko and Chaussidon, 1997).

The basalts have depleted trace element patterns (Hemond et al., 1993; Chauvel and

HÉMOND, 2000) and lithophile radiogenic isotopes provide evidence of several mantle

components that contribute to the observed variability of Icelandic basalts (e.g. Hemond et

al., 1993; Kempton et al., 2000; Fitton et al., 2003). The 3He/4He ratio increases from 12.7

to 20.2 RA in SW-NE direction along the fissure system. Samples with the highest 3He/4He

also show the most extensive degassing. However, the degassing does not exceed 30%

(Harrison et al., 1999, 2003). An inverse correlation between 3He/4He and CaO/Al203 has

been interpreted as reflecting variable clinopyroxene assimilation (BURNARD and HARRISON,

2005). Correlations between the isotope compositions of argon and oxygen are consistent

with low 5,80 (~4.6%o) in the source regions of these magmas (BURNARD and HARRISON,

2005).

Jan Mayen (71°N, 8°30'W) is situated south of the NW-SE trending Western Jan Mayen

Fracture Zone (WJMFZ) which separates the Kolbcinscy Ridge and the Mohns Ridge. The

island is composed of several cycles of potassic trachybasalts erupted by the Beerenberg

stratovolcano (2227 m above sea level). Volcanism is dominated by primitive alkali basalts

that evolve from the volumetrically dominant ankaramites through alkali olivine basalts to

trachytes (Maaloe et al., 1986). Partial melting of a spinel Iherzolitic source at > 1400°C and

~ 20 kbar followed by late-stage gravitational settling of alkali basalt magmas at low pressure

has been proposed as a way of generating the variability of Jan Mayen basalts and trachytes

(Maaloe et al., 1986). The Sr-Nd-Pb-Hf isotopes and incompatible trace element systematics

of Jan Mayen and the adjacent platform differ markedly from the nearby Kolbeinsey and

Mohns Ridges (SCHILLING et al., 1999; Tronnes et al., 1999; Mertz et al., 2004; BLICHF.RT-

Toft et al., 2005) in that they represent low-degree partial melts of enriched mantle reservoir

(Tronnes et al., 1999). Lead isotope signatures of Jan Mayen basalts provide support for a

HIMU-like source (Tronnes et al., 1999) whereas the absence of a strongly radiogenic Pb

isotope signature (mPb/204Pb ~ 18.6) reflects a short isolation time (approximately 500 Myr)

for the subducted material, consistent with conclusions from Icelandic basalts (THIRLWALL,

1997). On the basis of Pb isotope composition of basalts dredged from the Jan Mayen

platform, Blichert-Toft et al. (2005) describe its source as a mixture of depleted mantle

(DM), component "C" (as defined by HANAN and GRAHAM (1996) based on Pb-Hf-Nd

isotope systematics) and an enriched mantle end member (EM2).

68


We have analysed clinopyroxene and olivine phenocryst-bcaring basalts (strictly

ankaramites) from Jan Mayen that show minimal post-eruption alteration and have MgO >

8% to avoid the effects of wall-rock assimilation. The 87Sr/86Sr, 143Nd/144Nd and 3He/4He

variation is limited (0.7034-0.7035, 0.51289-0.51291 and 5-7 RA, respectively; STUART et al.,

submitted manuscript), but these samples represent some of the most enriched basalts in the

North Atlantic Ocean mantle (Stuart et al., submitted manuscript). Lead isotopes (Stuart

et al., submitted manuscript) are consistent with the young HIMU-like source proposed

previously (Tronnes et al., 1999). The narrow range of 3He/4He (5-7 RA) is typical of basalts

with a strong HIMU affinity, however similar values are recorded for the OIBs with enriched

mantle (EM) affinities (GRAHAM, 2003).

4.4 Analytical techniques

Lithium was separated using an HNOî-methanol mixture with 2.1-ml cation-exchange

resin BioRad AG50W-X8 (mesh 200-400). The second step of cleaning was applied to Jan

Mayen samples using 0.6-ml cation-exchange resin BioRad AG50W-X12 (mesh 200-400) for

additional purification as described elsewhere (Magna et al., 2006b). Lithium isotope

measurements were performed on Nu1700 (Nu Instruments, Wrexham, UK) multiple-

collector inductively coupled plasma mass spectrometer (MC-ICPMS) using simultaneous

collection of 6Li (low-mass Faraday cup L7) and 7Li (high-mass Faraday cup H8). Samples

were measured relative to the L-SVEC international reference standard (Flesch ct al., 1973)

and reported as 87Li (%o) = [(7Li/6Li)sampic/(7Li/6Li),tandaid-lJx1000. A detailed description of

the method is given in Magna et al. (2004b). Two reference materials were measured

together with samples to assess reliability of the analytical method. JB-2 (Geol. Survey of

Japan) and BHVO-2 (US Geol. Survey) yielded 87Li = 4.84 ± 0.15%o (2a) and 4.58 + 0.32%o

(2a), respectively. This is in excellent agreement with previously published values

(Moriguti and Nakamura, 1998a; Tomascak et al., 1999a; Jeffcoate et al., 2004;

Magna et al., 2004b).

Lithium concentrations were measured on a quadrupole ICP-MS PlasmaQuad 2 (VG

Elemental) using beryllium as internal standard. For other trace element concentrations

rhodium was used as an internal standard. The precision for most trace clement measurements

was better than 5% (2a), for Cs, Ta, Pb, Th and U better than 10% (2a). Accuracy of the data

69


was monitored with BHVO-2 (US Geological Survey) and JB-2 (GSJ), which were analysed

together with samples.

Strontium (Sr) and neodymium (Nd) were separated from the Icelandic basalts using the

procedures of HORWITZ et al. (1992) and Cohen et al. (1988), respectively. Strontium and

Nd isotope compositions were measured on a Nu Plasma MC-ICPMS. Masses 83 and 85

were used to correct for Kr and Rb interferences, respectively. The instrumental mass bias

was corrected using an exponential fractionation law. Measured 87Sr/86Sr ratios were

corrected for mass bias using 88Sr/86Sr = 8.3752 while 143Nd/44Nd were corrected using

146Nd/,44Nd = 0.7219. The ,43Nd/,44Nd ratios were further normalised to the JMC-Nd

standard using a 143Nd/144Nd = 0.511833 (cross-calibrated to La Jolla standard value

0.511858) and 87Sr/86Sr were normalised to the ratio of 0.710245 for the NIST SRM 987.

4.5 Results

Lithium concentrations and isotope compositions from the Iceland and Jan Mayen

basalts are presented in Table 4.1 along with other pertinent radiogenic isotope data. Trace

element data for Iceland and Jan Mayen lavas are presented in Tables 4.2 and 4.3,

respectively. The trace elements define two distinct patterns (Fig. 4.1). The "Enriched Group"

samples show elevated Lan/SmN > 1.1, while the "Depleted Group" is characterised by

Laivi/SmN < 0.6. The Depleted Group has pronounced Zr and Hf depletions relative to

similarly incompatible REEs (Fig. 4.1). The samples display a narrow range of 87Sr/S6Sr =

0.7030-0.7032 and ,43Nd/,44Nd = 0.51302-0.51312 (Table 4.1).

The Li isotope compositions scatter substantially in Icelandic picritic glasses. The 87Li

values of the Enriched Group range between +3.9 and t-5.8%o whereas the Depleted Group

shows slightly larger variation in ô7Li (+3.8 to +6.9%o). There is a striking linear and negative

correlation between 87Li and 3He/4Hc (Fig. 4.2) in which low 57Li values, typical of MORB-

like mantle, are associated with high 3He/4He that are diagnostic of the Iceland plume. We

explore the implications of this later.

Lithium concentrations of Jan Mayen ankaramitic basalts vary between 2.2 and 5.1 ppm

but, in contrast to the Icelandic basalts, they show extremely homogeneous ö7Li (+3.9 to

+4.7%o; Table 4.1 ). This is consistent with the homogeneous trace element (Table 4.1) and Sr-

Nd isotope compositions (Stuart et al., submitted manuscript) of these samples, and Jan

Mayen basalts in general (Tronnes et al., 1999; Blichert-Toft ct al., 2005). The rare earth

70


element patterns (LaN/SmN> 3.25) are steeper than in the Iceland lavas (Fig. 4.1) and imply a

more enriched source. Unlike the Icelandic basalts, the Jan Mayen ankaramites show Pb

enrichment relative to neighbouring elements (Fig. 4.1). The Jan Mayen basalts arc enriched

in Zr and Nb relative to MORB and possess strongly positive ANb (Fitton et al., 1997)

values (0.42-0.49; Table 4.3). Subtle depletions in Sr, Zr and Hf relative to neighbouring

elements of similar incompatibility are observed.

Table 4.1 : Li concentrations and Li, Sr

Jan Mayen basalts

and Nd isotope compositions of Iceland and

Li (ppm) 57Li (%o)a 2a S7Sr/86Srb 143Nd/l44Ndc

Iceland

DICE 5 2.8 6.90 0.34 0.703040 (14) 0.513115 (6)

DICE 7 2.6 6.29 0.19 0.703015 (12) 0.513115 (9)

DICE 8 3.8 5.84 0.37 0.703104 (14) 0.513066 (6)

DICE 9 4.5 3.88 0.39 0.703141 (17) 0.513027 (5)

DICE 10 2.6 4.57 0.37 0.703089 (14) 0.513081 (7)

DICE 11 2.4 3.72 0.19 0.703095 (14) 0.513083 (6)

DICE 13 4.1 4.01 0.18 0.703209 (12) 0.513023 (6)

DICE 15 4,1 5.21 0.17 0.703120 (11) 0.513038 (5)

DICE 43 5.0 4.38 0.26 0.703194 (17) 0.513019 (7)

MAE 2.8 5.50 0.09 0.703007 (13) 0.513101 (7)

Jan Mayen

JM14 5.1 3.88 0.25

JM44 4.8 3.92 0.41

JM78 3.9 4.00 0.16

JM207 4.3 3.86 0.20

JM208 4.0 4.38 0.33

JM 16837 3.8 4.70 0.39

JM 16841 2.9 4.49 0.23

JM16852 2.2 4.04 0.44

Sr and Nd isotope compositions of Jan Mayen ankaramitic lavas from this study are

given elsewhere (STUART et al., submitted manuscript).J

Li isotope compositions are expressed in per mil relative to Li reference standard

L-SVEC (NIST SRM 8545).bNumbers in parentheses are 2rj errors and refer to the last two digits.

'

Numbers in parentheses are 2a errors and refer to the last digit.

71


Table 4.2: Trace element concentrations of Iceland picritic lavas

DICE 5 DICE 7 DICE 8 DICE 9 DICE 10 DICE 11 DICE 13 DICE 15 DICE 43 MAE

Rb 0.37 0.34 2.28 4.00 0.28 0.20 4.30 2.61 9.10 0.31

Cs 0.014 0.008 0.020 0.025 0.005 0.007 0.052 0.025 0.098 0.006

Sr 117 114 168 233 124 113 278 239 324 122

Ba 9.3 8.7 48 85 7.2 6.7 87 66 139 8.5

Sc 29 28 42 34 26 28 43 42 43 27

Y 14 13 23 25 12 12 27 27 33 14

Zr 25 23 64 105 19 18 114 106 160 25

Hf 0.73 0.68 1.54 2.4 0.63 0.57 2.59 2.41 3.6 0.69

Nb 1.15 1.24 8.4 17 1.00 0.95 20 16 30 1.25

Ta 0.14 0.11 0.48 0.95 0.08 0.08 1.07 0.81 1.64 0.09

Co 55 56 58 68 53 55 59 55 59 53

Ni 218 250 112 287 180 202 152 116 102 199

Pb 0.09 0.07 0.34 0.63 0.09 0.04 0.61 0.54 1.07 0.13

Th 0.06 0.06 0.37 0.83 0.04 0.03 0.74 0.50 1.35 0.05

U 0.02 0.02 0.12 0.23 0.01 0.01 0.21 0.15 0.36 0.02

La 1.16 1.15 5.63 10.7 0.99 0.94 12.0 9.73 18.4 1.14

Ce 3.38 3.26 13.9 24.9 2.92 2.67 28.2 23.4 42.4 3.44

Pr 0.68 0.62 2.12 3.59 0.60 0.53 3.99 3.43 5.93 0.64

Nd 3.77 3.45 9.90 16.2 3.29 3.03 17.9 15.9 26.0 3.66

Sm 1.65 1.44 3.21 4.64 1.37 1.32 5.10 4.69 6.94 1.62

Eu 0.68 0.63 1.30 1.70 0.63 0.59 1.90 1.79 2.45 0.70

Gd 2.00 1.83 3.49 4.72 1.74 1.75 5.32 4.82 6.86 1.97

Tb 0.38 0.36 0.63 0.80 0.34 0.33 0.88 0.83 1.10 0.38

Dy 2.64 2.46 4.28 4.82 2.32 2.29 5.46 5.35 6.57 2.64

Ho 0.56 0.50 0.89 0.95 0.52 0.47 1.09 1.05 1.31 0.56

Er 1.66 1.53 2.61 2.82 1.40 1.48 3.07 3.07 3.82 1.66

Tm 0.24 0.21 0.38 0.38 0.21 0.20 0.43 0.44 0.52 0.24

Yb 1.59 1.43 2.40 2.47 1.36 1.26 2.68 2.85 3.37 1.44

Lu 0.22 0.21 0.35 0.36 0.19 0.19 0.39 0.41 0.47 0.22

(Sr/Nd)N 1.99 2.12 1.09 0.92 2.42 2.39 1.00 0.96 0.80 2.14

(La/Sm)N 0.45 0.52 1.13 1.49 0.47 0.46 1.52 1.34 1.71 0.46

EuN/Eu*'

1.14 1.19 1.18 1.10 1.25 1.19 1.11 1.14 1.08 1.20

Trace element concentrations are given in ppm. Trace element ratios arc normalised to primitive mantle (SUNand Mcdonough, 1989).

a

EuN/Eu is calculated from the interpolated Eu* defined by the primitive mantle-normalised abundances of Sm

and Gd.

72


Table 4.3: Trace element concentrations of Jan Mayen ankaramitic lavas

JM 14 JM44 JM78 JM207 JM208 JM 16837 JM 16841 JM 16852

Kb 51 50 32 39 40 38 19 21

Cs 0.48 0.44 0.27 0.30 0.38 0.35 0.14 0.18

Sr 673 658 479 523 531 501 378 304

Ba 828 823 537 667 669 620 438 346

Y 26 26 19 21 21 20 16 13

Zr 218 222 140 181 180 169 124 91

Hf 4.7 4.7 3.2 3.9 3.8 3.6 2.9 2.3

Nb 79 79 54 66 66 61 42 32

Ta 4.7 4.8 3.2 4.0 3.9 3.6 2.5 2.0

Co 50 54 52 55 55 56 59 59

Ni 712 623 274 772 742 733 621 344

Pb 12 5.0 2.8 6.5 6.4 3.3 2.7 1.8

Th 5.9 5.8 3.6 5.0 5.1 4.6 2.5 1.4

U 1.4 1.3 0.87 1.2 1.1 1.0 0.70 0.52

La 52 51 36 44 44 41 28 22

Ce 101 100 70 84 84 78 56 43

Pr 12 12 8.3 9.9 9.9 9.2 6.8 5.3

Nd 45 45 31 37 37 35 27 21

Sm 8.1 8.1 5.9 6.7 6.5 6.4 5.3 4.3

Eu 2.6 2.7 1.9 2.2 2.1 2.0 1.6 1.3

Gd 8.0 7.6 5.6 6.6 6.6 6.1 5.0 4.1

Tb 1.0 1.1 0.74 0.85 0.84 0.83 0.70 0.57

Dy 5.2 5.4 3.8 4.2 4.3 4.0 3.5 2.9

Ho 0.96 0.98 0.71 0.81 0.82 0.76 0.65 0.53

Er 2.7 2.6 1.8 2.2 2.1 2.0 1.6 1.3

Tm 0.36 0.35 0.24 0.29 0.29 0.28 0.23 0.19

Yb 2.2 2.3 1.5 1.8 1.9 1.7 1.4 1.2

Lu 0.32 0.32 0.23 0.27 0.26 0.26 0.21 0.18

(Sr/Nd)N 0.96 0.94 0.98 0.91 0.91 0.92 0.90 0.92

(La/Sm)N 4.13 4.10 3.94 4.27 4.35 4.10 3.44 3.25

Eum/Eu*° 0.98 1.04 0.99 1.00 0.98 0.99 0.96 0.92

ANb" 0.42 0.41 0.49 0.41 0.41 0.41 0.42 0.49

Trace element concentrations arc given in ppm. Trace element ratios are normalised to primitivemantle (SUN and MCDONOUGH, 1989).

a RuN/Hu* is calculated from the interpolated Eu* defined by the primitive mantle-normalised

abundances of Sm and Gd.bANb is calculated after FlTTON et al. ( 1997) as ANb = 1.74 + log(Nb/Y) - 1.92xlog(Zr/Y)

73


1000

*->100

c

CO

f"*

0>^^J 10

E*i_

Cl-

CO>CO 1

0.1 ' 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 H

Cs Rb Ba Th U Nb Ta La Ce Pb Pr Sr Nd Sm Zr Hf Eu Gd Tb Dy Y Li Ho Er Tm Yb Lu

Fig. 4.1 • Primitive mantle-normalised trace elements from Iceland and Jan Mayen. Trace element patterns for

basaltic glasses from MMF 'system can be divided into the Enriched Group (open) and Depleted Group (black).

Both lava groups are depleted in hi and Pb and enriched m Ba, Nb and Ta compared to similarly compatible

elements. Depleted Group is also enriched in Sr and Zr- and Hf-depleted relative to neighbouring elements. A

similar trace element pattern has been observed independently of our study in the Hengill picrite HEN5 and

ascribed to recycled oceanic crust as one constituent for Icelandic plume (CllAUVEl and Hémond, 2000). A

small positive Eu anomaly has been detectedfor both lava groups: hi is depleted in both groups. Trace element

pattern ofJan Mayen ankaramites is shown in grey colour. It closely resembles that ofJan Mayen lavas from

TR0NNES et al. (1999). Note Li depletion and Pb enrichment relative to neighbouring elements. Enriched

character ofthe lavas is documented by steepness of the primitive mantle-normalised pattern with high LaN/SmN

values >3.2. Normalisation values are takenfrom SUN and McDonough (1989).

4.6 Discussion

4.6.1 Lithium isotope composition ofthe North Atlantic mantle constrainedfrom Icelandic

basalts and Jan Mayen ankaramites

Surprisingly few data exist for Li isotope compositions of Icelandic basalts. Two studies

reveal ô7Li values that are indistinguishable from the results in this paper (Pistiner and

74


Henderson, 2003; Ryan and Kyle, 2004). One basalt with major clement compositions

similar to the Depleted Group (AU389) gives 57Li of +4.0%o (Ryan and Kyle, 2004),

consistent with the high-3He/4He DICE 10 and DICE 11 lavas. The interior part of the

effectively unaltered Laki basalt shows slightly lighter Li isotope compositions (S Li =

+3.2%o) with some scatter throughout the sample transect (+l.l%o; Pistiner and Henderson,

2003), a value in agreement with data from this study. A preliminary study of olivine

tholeiites from Theistarcykir and Krafla, two geographically distinct volcanic centres in

Iceland, shows a large spread of 57Li at both localities (+0.9 to +4.3%o; -0.8 to +6.3%o,

respectively; Hansen et al., 2005). Light Li isotope compositions of primitive high-MgO

(>8%) post-glacial Theistareykir lavas do not reflect contamination by Icelandic crust but are

believed to represent intrinsic features of the magmatic sources as also evidenced by

radiogenic isotopes (e.g. ELLIOTT ct al., 1991; STRACKE et al., 2003). On this basis the

Icelandic plume may have ö7Li in a similar range to that found for Theistarcykir lavas

(Hansen et al., 2005).

Data from this study show Ö7Li values for high- He/ He lavas that are similar to those of

the above studies (see Table 4.1). Lavas with 3He/4He higher than 17 RA have an average 67Li

of ~+4.3 + 1.0%o. This is in excellent agreement with the inferred mantle S7Li value at ~+4%o

based on data from pristine spinel lherzolites (Magna et al., 2006b). Recent Ô Li data of

fresh MORBs seem to be inconsistent with any provinciality in distribution of Li isotopes

among Mid-Atlantic Ridge, Southeast Indian Ridge and East Pacific Rise (TOMASCAK et al.,

in press). Furthermore, chemical variability (K20/Ti02 = 0.05-1.55) is not mimicked in any

way by large Li isotope variations in MORB. Three fertile peridotites from the study of

Jeffcoate et al. (2007) show an average Ô7Li of +3.6%o and are consistent with previous

investigations of Li isotopes in pristine peridotites (Seitz et al., 2004; Magna et al., 2006b).

These peridotitic data further validate the recently postulated Li isotope composition of the

upper mantle at +4.2 ± 0.8%o (Chan, 2003) as inferred from fresh MORBs and a single

Hawaiian lava.

That intra-plate basaltic volcanism from Hawaii and Iceland have Ô Li close to +4%0

(TOMASCAK et al., 1999b; Chan and Frey, 2003) implies a single value for the Li isotope

composition of the terrestrial mantle which is constant and independent of trace element

depletion/enrichment. This is corroborated by Jan Mayen data from this study that show an

75


10.0

8.0

o

6.0

r*.

to4.0

D

25

2.0

0.0

100

75

25

4.8

J 4.6

oo 4.4

to

4.2

4.0

80

40

00

4 0

-8 0

o

r"v -

--

,

«x>

I -^

"

'•

. 10

, \l 0 V

10 20 30 40

•%-,.0>

presumedLotht source

-j i i j- i i i i i I i i i i 1 i

o

o

o

o o

10 15 20

3He/4He (R/RA)

c

'

0

-

o

'

o

0

1 i 1 1

0

1 1 1

25

76


overall homogeneity in <57Li (+4.2 + 0.6%o). Therefore, Li seems quite homogeneous,

regardless of the fact that "'He^He < MORB in these lavas (Stuart et al., submitted

manuscript) and 3He/4He > MORB in the majority of MMF lavas (Burnard and Harrison,

2005). Both of these suites involve distinct components required to produce the observed

trace elements and radiogenic isotopes.

The explanation for this could be that the Li isotope compositions are to some extent

buffered by a dominant mantle component even after entrained material with different trace

elements and isotope ratios was added. Deviations to either light or heavy ô Li values would

in turn provide evidence for incorporation of volumetrically significant pollutants with largely

different Li isotope signatures (eclogites, seawater-altered crust, continental crust). However,

lack of Li isotope data for other OlBs does not permit precise estimates of the S7Li of such

environments. Collectively, the data from this study suggest only limited variations in Li

isotope compositions with a global mantle average ô7Li ~ +4%o.

4.6.2 Li-He isotope correlation in Icelandic basalts

The Icelandic basalts from this study display an inverse and linear correlation between

S7Li and 3Hc/4He (Fig. 4.2) that is, however, not well reflected in other trace element and/or

Fig. 4.2: Plot ofHe isotope variations with Li isotopes, Nb/U and O isotopes: The linear correlation between Li

and He isotopes (a) is assumed to reflect mixing of two reservoirs (see text for further discussion). Open

diamonds - Enriched Group; closed squares - Depleted Group; grey circles - Jan Mayen lavas. Grey rectangle

represents the MORB array (hi range from TOMASCAK, 2004; He range from GRAHAM, 2003). The two grey

hexagons represent the presumed hoihi source lavas (hi isotope datafor Mauna Kea and presumed hoihi source

arefrom ClIAN and FREY (2003), He isotope data arefrom Kurz et al,, 2004). Dashed line is a least square best

fit for the data points. Eclogites lend to have very low, even negative cfhi combined with radiogenic 3HefHe as

marked in inset ofa by red and green line that, by keeping linear relationship, give an estimate of the minimum

ö hi requiredfrom the highest 3He/4He in Iceland (red line; 37 RA and —2%o) and worldwide (green line; 50 RA

~-7%o). Blue horizontal bar represents cfLi of the Earth's mantle estimatedfrom available pristine peridotitic

data (+3.6 to +4.2%o). In b, He isotopes are plotted against Nb/U, which is generally homogeneous in MORB

and OIB reservoirs (Hofmann, 1997). Hence, high Nb/U is most likely inherited from depletion in U and

residual enrichment in Nb during alteration ofocean floor and subduction. 'The Jan Mayen lavas lie away from

the Icelandic lavas because their Nb/U ratios fall close to average value of-50 inferred for oceanic basalt

suites (HOFMANN, 1997). In c, broadly negative correlation ofHe and O isotopes may be observed, in particular

obvious for Enrichedgroup lavas (see textfor discussion).

77


isotope ratios. The high 3Hc/4He of Iceland plume-derived basalts arc thought to reflect the

contribution from a deep mantle that is less degassed of primordial volatiles than upper

asthenospheric mantle (Stuart et al., 2003). Therefore, the simplest explanation for the

observed He-Li isotope trend is that it represents a mixing between a primordial gas-rich deep

mantle source and altered Icelandic crust. Altered crust represents a reservoir with enrichment

in 7Li and high Li abundances in tens of ppm. The mantle likely has 87Li ~ +4%o with

minimal scatter. Parts of the deep mantle may carry high 3He/4He ~50 Ra (Stuart et al.,

2003) whereas crustal materials have a large contribution of radiogenic 4He tending to low

3He/4He. However, four lines of evidence suggest that this model is too simplistic.

Firstly, the öi80 of all samples (4.1 to 4.6%0; Burnard and Harrison, 2005) is lower

than most oceanic basalts. The depletion in 180 transcends the trace element

enrichment/depletion and is common to the whole sample suite. It has been suggested that Ar-

O isotope systematics in these lavas could reflect surficial processes (e.g. assimilation of

meteoric fluids or material altered by meteoric fluids; Burnard and Harrison, 2005).

However, as pointed out above this is not considered to be a likely explanation for the

ubiquitous low-180 Icelandic volcanism. It is very possible that large-scale hydrothermal

alteration with meteoric fluids provided source rocks with low 5180 (Gautason and

Muehlenbachs, 1998) that were formed in the past, as in Palaeozoic ophiohtes

(Muehlenbachs ct al., 2004) for example, and then recycled. A low-8,80 mantle end-

member is required that is depleted in incompatible trace elements and carries 3He/4He < 12

Ra together with elevated ö7Li values, Eu enrichment and superchondritic Sr/Nd.

Incompatible-tracc-elcment-enriched lavas could well represent a mixture of such a recycled

low-o180 component with a less degassed mantle component having high 3He/4He (Burnard

and Harrison, 2005) and normal mantle Sr/Nd. In this case, a negative correlation between

8180 and 3He/4He would be observed (Thirlwall et al., 2006). Indeed, such a tendency can

be seen in the MMF lavas (Fig. 4.2).

Second, trace elements provide evidence of recycling and feldspar accumulation, even

though none of the basalts analysed contain feldspar as a phenocryst phase. The Eu anomaly

is always positive (Table 4.2) and correlates with Sr/Nd (Fig. 4.3), consistent with plagioclase

accumulation. Recycled gabbroic cumulates that have been converted to eclogite or

pyroxenite could carry the positive Eu anomalies and high Sr/Nd imparted from an earlier

cumulate history, as has been proposed for Hawaii (HOFMANN and JOCHUM, 1996). The

evidence for feldspar accumulation in the prehistory of these magma sources is striking. The

78


samples that characterise the low-3He/4He high-57Li end-member (DICE 5, 7 and MAE) are

rich in feldspathic cumulates. In principle, these could be either sub-Icelandic crust or

recycled gabbroic lower oceanic crust. If these magmas reflected vastly different portions ofa

recycled component it would be hard to explain the restricted range of 87Sr/86Sr = 0.7030-

0.7032 and 143Nd/,44Nd = 0.51302-0.51312 (Table 4.1) and similar overall trace element

patterns (Fig. 4.1 ).

1.30

1.20

3 1 10LU

jjj 1.00

0.90

0.85

3.00

t 2.00

CD

CO

1.00

0.00

Enriched group 0- f m

O m

ooo

Depleted group

o

fyo

Jan Mayen Island

o

o

m

m

Soo

oo

m

m

0.0 0.5 1.0 1.5 2.0 2.5 3.0

(Sr/Nd)N

Fig. 4.3: EuN/Eu und (Ba/Th)N vs. (Sr/Nd)N. The plot reveals that Enriched Group lavas have tawer Sr/Nd

ratios than Depleted Croup lavas. Within each group a steep positive trend can be seen. The largely distinct

Sr/Nd element ratios require different sources. EuN/Eu* is calculated from the interpolated Eu* defined by the

chondrite-normalised abundances of'Sm and Gd. (Ba/Rb)N vs. (Sr/Nd)N plot (not shown) has very similar pattern

as (Ba/Th)N vs. (Sr/Nd)N plot. Jan Mayen lavas mostly lie on the intercept of mantle values for (Sr/Nd)^ and

EuN/Eu but two samples (16841, 16852) show remarkable Th depletion relative to other Jan Mayen lavas which

results in their substantially higher (Ba/Th)N ratios. The symbols are the same as in Fig. 4.2.

79


Third, high Nb/U and negative Pb and Li anomalies are a characteristic of most

samples. High Nb/U in particular implicates U-loss during a prior subduction history.

Niobium is effectively retained in the subducted residue (Rijdnick et al., 2000) whereas U is

removed into the mantle underlying arc and back-arc lavas (Kelley et al., 2005). Alteration

of the upper oceanic crust elevates U abundances (Elliott et al., 1999) whereas the Nb

budget remains largely undisturbed. The mobility of U was only augmented after the Archean

as a result of dramatic increase in oxygen fugacity (Elliott et al., 1999). All but one sample

has Nb/U higher than the upper range of the global average (47 + 10) compiled by Hofmann

( 1997) and eight out of ten Icelandic lavas from this study have Nb/U > 70 (see Fig. 4.2),

This superchondritic Nb/U must have evolved in an ancient subduction event and cannot

reflect present-day processes.

Fourth, if the Li-He isotope correlation represents a mixing between altered Icelandic

crust and deep mantle, the linearity of the data array requires that the [He]/[Li] of the end

members are broadly similar. From the He-0 isotope relationship, Burnard and Harrison

(2005) inferred a mixing between depleted MORB mantle and hypothetical Iceland source

region where [He]niviM ^ 10x[HeJicc- Keeping this relationship would, accordingly, entail a

ten-fold difference in Li abundances between the two respective end-members which is not a

likely option for the mantle components because the bulk distribution coefficient for Li in

mantle melting is probably not far from unity with Li abundances that differ by a factor of ~2-

3 at most between depleted MORB mantle and Iceland source region.

Hydrothermally altered rocks carry in general isotopically light Li (Chan et al, 2002a).

However, at this time insufficient is known to evaluate whether this is a normal feature of

180-depleted altered oceanic crust. Petrologic studies and major element chemistry that show

that all these samples ascended through the Icelandic crust quickly because they contain

numerous gabbroic xcnoliths and both DICE 10 and DICE 11 contain partially resorbed

olivines (Tronnes, 1990). Sparse analyses of gabbroic rocks show that they may have 57Li >

fresh MORB while maintaining low Li contents (CHAN et al. (2002a) and T. Magna, unpubl.

data) which could be a result of addition of late stage low-temperature alteration products

(Chan et al., 2002a). The Icelandic silicic rocks carry Sr and Nd isotope compositions like

those of the basalts (e.g. O'Nions and Grönvold, 1973; Sigmarsson et al., 1991, 1992;

GUNNARSSON et al., 1998; MARTIN and SIGMARSSON, 2007). However, these evolved rocks

are strongly depleted in Ba, Sr and Eu (JÔnasson, 1994, 2007; Gunnarsson et al., 1998;

Martin and Sigmarsson, 2007) and carry 5180 values heavier than those measured in DICE

80


lavas (Sigmarsson ct al., 1991, 1992; Burnard and Harrison, 2005; Martin and

Sigmarsson, 2007). These features combined with the major element compositions, render

the silicic rocks an improbable contaminant.

By extrapolating the correlation to the highest He/ He measured in Iceland («37 Ra;

Hilton et al., 1999; Ellam and Stuart, 2004) or globally («50 Ra; Stuart et al., 2003),

the high-3He/4He end-member would possess 57Li < ~-2%o or ~-7%o, respectively (see inset in

Fig. 4.2). So far, the only rocks with negative Li isotope compositions are eclogites which

have suffered dehydration during subduction (Zack et al., 2003). However, the loss of mantle

He during subduction of oceanic crust, and the subsequent ingrowth of radiogenic He from U

(and Th) decay will, in all likelihood, leave eclogites with extremely low 3He/4He making

them unlikely candidates as the high-3He/4He reservoir in the Earth. More importantly, the

fact that the lowest 57Li found in the Iceland samples is the same as "normal" mantle renders

as extremely unlikely the notion that this composition is a fortuitous mixture of the exactly

correct mixture of light and heavy Li. Accepting this to be the case the "light" end member

actually has ô7Li ~ 4%o and 3He/4He = 20 RA.

Experimental partitioning data for noble gases (Heber et al., 2007) show that, like Li,

they are less incompatible than K, U and Th in olivine and clinopyroxene, and tentatively also

in orthopyroxene. Hence, recycling of previously melt-depleted ancient lithosphère with

reasonably constant He/Li but depletions in K, U and Th may lead to variable ^He^He while

maintaining low total He abundances and lead to a linear functional relationship between Li

and He isotope compositions. This would be consistent with the relationships between the Nd

and Sr isotope compositions of the Iceland samples and their parent / daughter ratios (Sm/Nd

and Rb/Sr, respectively; not shown) which are consistent with fractionation during the

Phanerozoic.

Recent work has highlighted the relative importance of diffusion-driven Li isotope

fractionation (Richter et al., 2003; Lundstrom et al., 2005) that may add to Li isotope

variability in the limited scales of transects. Although it may prove important for local-scale

sampling and for intra- and inter-mineral diffusive isotope fractionation as illustrated recently

(e.g. Jeffcoate ct al., 2007; Rudnick and Ionov, 2007), lavas from MMF system are

dispersed along some ~15 km and are not directly related to each other. Also, any changes in

ô Li due to diffusion vanish after some one hundred meters away from contact and Li

abundances vary only within few meters of the contact (TENG et al., 2006a). Therefore, a

direct assessment of diffusive Li isotope fractionation is irrelevant here. Diffusive

81


fractionation of He isotopes was advocated for a suite of samples from Siberia which was

underlain by assuming faster diffusion of He and by positive correlation between [ He] and

3He/4He (Harrison et al., 2004). Also Burnard et al. (2004) have shown that some limited

fractionation of He isotopes due to diffusion is plausible though not explaining total variation

in 3He/4He. There is, however, no similar correlation observed in MMF lavas from this study

and we infer diffusion-driven isotope fractionation is an unlikely mechanism for the observed

Li-He isotope correlation.

4.6.3 Implicationsfor the origin ofJan Mayen basaltsfrom Li isotopes

Further evidence that recycled components are important in the North Atlantic mantle

despite normal Li isotopes comes from the Jan Mayen ankaramites. The incompatible trace

element concentrations in Jan Mayen basalts (Table 4.3) are at least twice as high as those in

basalts from adjacent Mohns and Kolbeinsey Rigdes (Mertz et al., 1991; Haase ct al., 1996;

Schilling et al., 1999). Combined with highly radiogenic Sr and the least radiogenic Nd

(Stuart et al., submitted manuscript) and Hf (Blichert-Toft et al., 2005) measured in the

North Atlantic they provide evidence for a particularly enriched mantle beneath Jan Mayen

that is exotic and largely dissimilar to other localities in the North Atlantic Ocean. The Jan

Mayen basalts have trace element compositions that may indicate a contribution of HIMU-

like mantle in the source region (Thirlwall, 1997; TR0NNES et al., 1999). However, the Pb

isotopes are not highly radiogenic (Stuart et al., submitted manuscript) and the 87Li values

(+3.9 to +4.7%o) appear to be markedly different from previous analysis of basalts with a

HIMU affinity. Heavy ô7Li values (ô7Li > +5%o) are recorded for the HIMU component in

basalts from the Polynesian islands (Nishio et al., 2005). Ryan and Kyle (2004) have

modelled the HIMU component in alkali basalts from St. Helena to have Ô Li = +8.5%o.

Although these basalts show clear signatures of weathering which might explain the high 87Li

(Ryan and Kyle, 2004), the lavas record no loss of alkali elements which in turn means they

likely reflect the original magmatic signature. The simplest interpretation for the source of the

heavy Li isotope component in HIMU basalts is that they are derived from melting of a

partially-dehydrated deep oceanic crust (Nishio et al., 2005), a suggestion in agreement with

input of heavy Li into the subduction zones (Bouman et al., 2004). This is consistent with the

origin of HIMU in dense oceanic crust that has been hydrothermally depleted in lead, on the

basis of Pb isotopes and trace elements (HOFMANN, 1997; WILLBOLD and STRACKE, 2006).

82


Although the absence of heavy Li isotope component in the Jan Mayen basalts contrasts

with HIMU basalts, similar 87Li values (+3 to +5%o) have been measured in intra-plate

Antarctic basaltic lavas with HIMU affinity (Ryan and Kyle, 2004). Their Li isotope

compositions were explained as a mixture of a HIMU component with moderately elevated

87Li (~ +8.5%o) and a depleted mantle source of MORB. Such heavy Li signatures (S7Li ~

+7.5%o) have been observed in Polynesian HIMU lavas (Nishio et al., 2005) that have been

explained by entrainment of a less altered basaltic part of oceanic crust in the source of these

lavas. Hawaiian lavas with HIMU-like Pb isotope signatures possess 57Li values that are

indistinguishable from MORB (Kobayashi et al., 2004). However, glass inclusions trapped

in some of these lavas do carry extremely negative Li isotope compositions (87Li ~ -10%o)

explained by recycled dehydrated oceanic crust in the source of Hawaiian mantle plume

(Kobayashi et al., 2004). This is similar to a conclusion made for two anomalously light

samples from Koolau volcano (S7Li ~ +2.6%0; Chan and Frey, 2003).

The Jan Mayen ankaramites display large Pb enrichments and Th-U depletion relative to

neighbouring elements which is inconsistent with HIMU. Such basalts in general show

significant Pb depletion (Hofmann, 1997; Willbold and Stracke, 2006). In fact, the trace

element pattern of Jan Mayen ankaramites (Fig. 4.1) and trace element ratios (e.g. Rb/Sr,

Th/U, Ce/Pb, U/Pb, Ba/La, Rb/La) are more similar to those of the EM2 component

(Workman et al., 2004; Willbold and Stracke, 2006). The Jan Mayen basalts may be a

mix of normal depleted mantle with enriched component where the 87Li is dominated by Li

contribution from the depleted mantle, following Ryan and Kyle (2004). This is supported

by the presence of a similar chemistry in off-axis volcanism elsewhere in the North Atlantic

Ocean (HAASE and DEVEY, 1994; TR0NNES et al., 1999).

Tronnes et al. (1999) have argued that Jan Mayen basalts are melts of laterally

dispersed Iceland mantle plume. Although the S7Li values for the Enriched Group of MMF

basalts (+3.9 to 15.9%o) overlap the Jan Mayen ankaramites (+3.9 to +4.7%o), several lines of

evidence are inconsistent with Jan Mayen and Iceland having the same source. First, the

projection of Icelandic mantle plume material to the north of Iceland is spatially restricted to

several hundreds of kilometres northwards (Thirlwall et al., 2004). Second, radiogenic

isotope trends in Iceland do not approach those measured in Jan Mayen (Mertz et al., 1991).

This is observed in more radiogenic Sr and less radiogenic Nd isotope compositions for Jan

Mayen ankaramites relative to MMF picritic glasses. Third, He isotopes show no sign of the

high values typical of the Iceland plume and are well below MORB (R/RA<7; Stuart et al.,

83


submitted manuscript). Fourth, Jan Mayen basalts have been produced by larger degrees of

enrichment in Nb-Zr than Icelandic basalts with resulting high ANb values (Table 4.3). This

parameter cannot be simply changed by entraining or contamination of a Nb-Zr-rich

component but largely reflects a source characteristic (Fitton et al., 1997, 2003).

It is obvious from trace element data (Table 4.3) that HIMU is not an ideal end-member

for Jan Mayen. Indeed, global dispersion of HIMU is strictly limited to the southern

hemisphere and is by no means conspicuous in the northern hemisphere. The trace element

pattern more likely resembles that of other enriched mantle end-members (see Willbold and

Stracke, 2006) and a similar prognosis has been made based on detailed Pb isotope studies

of basalts dredged from the Jan Mayen platform (BLICHERT-TOFT et al., 2005). EMI seems to

carry strongly negative S7Li values (NISHIO et al., 2004), as if originating from eclogitic melts

whereas the EM2 end-member might be expected to be slightly heavier than MORB

(Tomascak et al., in press; Nishio et al., 2004). The available Pb isotope data fall well into

the field of North Atlantic MORB given by Blichert-Toft et al. (2005) who assume that Jan

Mayen may reflect mixing of depleted mantle and EM2 component. A tentative estimate for

EM2 end-member is given at ô7Li ~+6%o (Nishio et al., 2004), much like the still heavier

HIMU end-member (Nishio et al., 2005). Available Pb isotope data show rather limited

variations for EM2 compositions (2l)f'Pb/2(,4Pb between -17.8 and -19.2 as compiled by

Hofmann (1997) and thus, do not exclude a spike of such material in the source of Jan

Mayen lavas (Fig. 4.4). A metasomatic origin has recently been proposed for EM2

(Workman et al., 2004) in contrast to an involvement of discrete portions of the upper

continental crust mixed with recycled oceanic lithosphère (WILLBOLD and Stracke, 2006).

However, the upper continental crust has rather narrow average of 87Li (0 + 2%o) and high Li

abundances of -35 ppm (Teno et al., 2004). This limits the maximum material from the

upper crust to <5%.

84

Chanter 4 North Atlantic

8.0

206pb/204pb

Fig. 4.4: Relation of coupled Li-Pb isotopes for Jan Mayen lavas. HIMU, EM2 and DM are involved as end-

member compositions. We propose EM2 component may add to chemically enriched character ofJan Mayen

lavas. Indeed, Pb isotope ratios in EM2 component (see data compilation in Hofmann, 1997) are consistent

with such a view. Grey bars represent different mantle end-members as given in TOMASCAK (2004). HIMU as an

enriched component is uneasy to satisfy both Li-Pb relationship as well as trace element patterns. Lead isotope

data arefrom (Stuart et al., submitted manuscript). See text forfurther discussion.

4.7 Conclusions

We have presented Li isotope data for two suites of basaltic lavas from the North

Atlantic. Lithium isotope compositions of lavas from Maelifell-MiÖfel Fissure (MMF) system

(Hengill area, Iceland) show a striking inverse and linear correlation with He isotopes. The

high- HerHe lavas show rather narrow range of S7Li averaging - +4.3%o whereas low-

3He/4He lavas possess isotopically heavier Li. The simple interpretation of the observed

relationship as being a mixture between mantle and surface altered crust is rendered

improbable by trace element data. Instead, recycled gabbroic oceanic crust is required in the

source ofMMF lavas from coupled Li isotopes and trace element systematics. Although exact

qualification of these end members is not straightforward, the observed ubiquitous Eu

enrichment hints to an old recycled component, efficiently mixed into the source of Depleted

85


DICE lavas. Data from this study do not support diffusive fractionation of both Li and He

isotopes, nor do they provide evidence of contamination by silicic volcanic rocks. The Jan

Mayen ankaramitic lavas are homogeneous in terms of Li isotopes and trace elements that

show remarkable enrichments in LREE. A possible HIMU origin of the enriched component

of Jan Mayen Island lavas is not easily reconciled with the data and we suggest these lavas

may include an EM2-likc component.

Lithium isotope compositions of lavas from this study support a view of relatively

homogeneous mantle 8 Li value of 14%o irrespective of their chemical enrichment / depletion

and evidence of recycling. The fact that He isotopes vary greatly in North Atlantic lavas all

displaying the same mantle-like Li isotope composition, O depletion and variable trace

clement enrichments reflecting recycled cumulates provides evidence that recycled lower

oceanic crust with variable (Ui-Th)/He was an important part of the source region. The fact

that Li isotopes are correlated linearly with He isotopes in Iceland provides evidence of an

additional recycled altered oceanic crustal component with different (U+Th)/He ratio but the

same He/Li. This is consistent with the fact that He and Li are both slightly but not highly

incompatible. There is no evidence that the He isotope compositions have been affected by a

high-3He component from a relatively undegassed reservoir.

Acknowledgement We thank Marie-Theres Bär for making the columns, Heiri Baur, Donat

Niederer, Bruno Rutsche, Urs Menet and Andreas Süsli for the electronics and mechanical

workshop services, Irene Ivanov-Bucher for help with sample preparation, Marcus Gutjahr

and Martin Frank for assistance by Sr and Nd isotope analyses and, in particular, Felix Oberli

for maintenance and troubleshooting of Nu 1700. We are grateful to Mark Rehkämper, Paul

Tomascak and Helen Williams for comments on previous version of the manuscript and

Francis Albarede, Janne Blichert-Toft, Karsten Haase and Kaj Hoernle for helpful

suggestions and/or information. Kaj Hoernle kindly provided samples from Canary Islands.

ETH and Swiss National Fonds (SNF) supported this work financially.

86

Chapter 5 Terrestrial planets

Chapter 5

New constraints on the lithium isotope

compositions of the Moon and terrestrial

planets

* published as: Magna T., Wiechert U. and Halliday A.N. (2006) New constraints on the

lithium isotope compositions of the Moon and terrestrial planets. Earth and Planetary Science

Letters 243 (3-4), 336-353

87

Chapter 5 Terrestrialplanets

5.1 Abstract

High-precision lithium (Li) isotope data are reported for samples from the Earth, Moon,

Mars and Vesta and provide evidence of broadly similar compositions that are slightly heavy

relative to those of chondrites. Mare basalts exhibit a large range of Li isotope compositions

(87Li = +3.4 to +6.4%o) that correlate with indices of magmatic differentiation. Three samples

(quartz-normative basalts and picritic orange glass) that are thought to have formed by

melting of relatively primitive source regions yield a mean 87Li = +3.8 ± 0.4%o taken as the

best estimate for the average composition of the Moon. Other samples are isotopically heavier

correlating with increases in Rb and Hf and probably reflecting transport of isotopically

heavy Li that formed in specific high-Ti cumulate - melt layers during crystallisation of the

magma ocean. The most extreme lunar 87Li is found in a ferroan anorthosite ()8.9%o).

Terrestrial mantle olivines fall into a tight range between +3.6%o and t-3.8%o. If these olivines

reflect the composition of the bulk Earth, the Li isotope compositions of Earth and Moon are

identical. The Li isotope compositions of samples from the Moon, Earth, Mars and Vesta

provide no evidence for differences between large inner solar system mantle reservoirs. This

in turn provides evidence that core formation, volatile loss and the presence of a crust and

hydrosphere have not significantly influenced the bulk Li isotope composition of the mantles

of these objects. The fact that chondrites are isotopically light compared with differentiated

planetary bodies of the inner solar system is consistent with a small but significant Li isotope

fractionation within the accretionary disc or chondrite parent bodies, the origin of which is

presently unclear.

5.2 Introduction

Theoretical models and isotope approaches for studying the origin of the solar system

are becoming increasingly sophisticated and powerful (CANUP and AGNOR, 2000; HALLIDAY,

2003; Chambers, 2004), but a major issue of current debate is the origin of the

heterogeneous distribution of volatiles, which it is now thought may partly relate to accretion

processes. It has been proposed on the basis of Sr isotopes that there were significant losses of

moderately volatile elements during the late (>10 Myrs) accretionary history of the material

that formed the Earth and Moon (HALLIDAY and PORCELLI, 2001). More recently it has also

been argued that the slightly heavier Fe isotope composition of basalts from the Moon,

88


relative to those from Vesta and Mars, may reflect volatilisation of Fe during the Giant

Impact (POITRASSON et al., 2004). Iron is only slightly volatile and the degree to which this

effect reflects vaporisation of molten iron metal as opposed to silicate is unclear. If the effect

is produced by volatilisation of silicate it should be present in other elements of similar half

mass condensation temperature. Lithium is such a slightly volatile element. O'Neill (1991)

predicted complete or nearly complete condensation of Li while most Na is evaporated during

formation of the Moon. Being very light, Li is expected to show sizeable isotope

fractionations during partial volatile loss.

Lithium (Li) isotope abundances have already been reported from a few meteorites

using high-precision multiple-collector inductively coupled plasma mass spectrometry (MC-

ICPMS; MCDONOUGH et al., 2003; SEPHTON et al., 2004). However, most of these are

chondrites. The few existing analyses suggest that different chondrite groups may have

distinct Li isotope compositions. This variation could reflect aqueous alteration on parent

bodies resulting in isotopically light Li in the residual highly metamorphosed chondrites and

isotopically heavy Li related to chondrites with high alteration index (McDonoijgh et al.,

2003). CM chondrites show large internal Li isotope variation (Sephton et al., 2004)

providing evidence of a complex history of evaporation and condensation of Ca- and Al-rich

refractory inclusions and chondmles in the solar nebula. These two constituents of chondrites

seem to have negative ö7Li values resulting partly from incorporation of spallogenic Li into

chondritic matter in the early solar system (ChausSidon et al., 2001 ; 2006).

The exact Li isotope composition of the bulk silicate Earth or primitive mantle has

proven problematic because the variation in Li isotope compositions in mantle rocks is large

(-17 to >+10%o; CHAN et al., 1992; 2002b; MORIGUTI and NAKAMURA, 1998b; TOMASCAK et

al., 1999b; Chan and Frey, 2003; Zack et al., 2003; Kobayashi et al., 2004; Nishio et al.,

2004; SEITZ et al., 2004). Most estimates of the Li isotope composition of the terrestrial

mantle are made on the basis ofMORB and very limited data for Hawaii (S7Li of 4.2 + 0.8%o;

Chan, 2003). Available analyses of peridotitic rocks seem to be consistent with a 57Li close

to +4%o for the Earth (BROOKER et al., 2004; SEITZ et al., 2004) although samples from the

latter study may have been influenced by subduction zone processes with heavy Li. Isotope

compositions of peridotites, calculated from mineral data, showed rather variable values (ô7Li

from 1.1 to H.l%0; Seitz et al., 2004). Two fertile lherzolites from Vitim (Russia) and

Dreiser Weiher (Germany) have calculated bulk Ô7Li of \ 3.8 and i4.1%o, respectively (SEITZ

89


et al., 2004), consistent with the above mantle value. Also Elliott et al. (2004) stressed that

the 87Li in N-MORB lavas displays a tight range of 3-4%o with slightly higher values in OIB.

We have analysed a range of samples to assess whether there are systematic differences

in Li isotope composition between planetary objects, in particular the Moon and to constrain a

87Li value for the bulk Earth. In addition to data for lunar basalts, eucrites, martian meteorites

and chondrites, we present Li isotope ratios for a suite of olivines, clinopyroxenes and

orthopyroxenes from spinel lherzolites in order to better estimate the 8 Li of the Earth's

mantle.

5.3 Methods and samples

All samples have been dissolved in a mixture of HF and HNO3 (6:1 v/v) and after

complete decomposition evaporated and refluxed with HNO3 several times to remove

fluorides. We used small-volume (2.1 ml) ion exchange columns packed with BioRad

AG50W-X8 (mesh 200-400) cation-exchange resin and employed IM HNO3 - 80% methanol

(v/v) mixture as an elution media. Isotope measurements were performed on the large-

geometry high-resolution multiple-collector ICPMS Nu 1700 (Nu Instruments). Details of

analytical and mass spectrometric procedures are given in MAGNA et al. (2004b).

Additional small columns (0.6 ml) packed with BioRad AG50W-X12 (mesh 200-400)

were used for final cleaning and effective removal of large quantities of chromium that

appeared in the Li fraction after the first step. This problem was notable for clinopyroxenes

from spinel lherzolites which contain up to 1-2% Cr. The use of 0.5M HNO3 eluted all Cr in 2

ml such that the subsequent Li fraction was completely "clean" (sec Figure 5.1). Although

this second step was precautionary and we did not quantify the influence of Cr on the

accuracy of Li isotope measurements it appears to be important for the acquisition of high

quality data.

Lithium abundances were measured using a PQ-2 quadrupole ICPMS (VG Elemental)

with beryllium as an internal standard for matrix effect corrections. Other trace element data

for lunar rocks, eucrites, Allende and martian meteorites were obtained using the PQ-2

quadrupole ICPMS (VG Elemental) with rhodium as an internal standard. Basaltic reference

standards were measured prior to samples and 2a errors (two standard deviations) were better

than 5% for most elements. Uncertainties in U and Pb determinations were better than 10%

90


(2a), Cs determinations were better than 20% (2a). The Eu anomaly Eun/Eu was calculated

from the interpolated Eu defined by the chondrite-normalised abundances of Sm and Gd.

Terrestrial samples comprise mineral separates of olivine, clinopyroxene and

orthopyroxene extracted from spinel lherzolites from San Carlos and Kilbourne Hole (USA),

Vitim (Russia), Atsagin-Dush and Tariat (Mongolia). Additionally, three lherzolites and one

harzburgite from Atsagin-Dush were analysed for whole-rock 87Li composition. San Carlos is

a primitive spinel lherzolite from Arizona, USA, studied for its major and trace elements by

Jagoutz ct al. (1979). Vitim (Siberia, Russia) has been depicted as a fertile off-craton

peridotite and in detail studied by IONOV et al. (2005). Samples from Atsagin-Dush,

Mongolia, were previously described in detail by WIECHERT et al. (1997). 8520-09 is a spinel

lherzolite for which Li concentration and isotope data for olivine and orthopyroxene were

acquired. 8520-12, -19 and -20 are spinel lherzolites whereas 8520-30 is a harzburgite. Tariat

spinel lherzolite is located in Tariat depression, Mongolia, and has been described in detail by

Press et al. (1986). Kilbourne Hole (New Mexico, USA) is a spinel lherzolite, another

aliquot of which has been studied by Jagoutz et al. (1979). Mineral separates from spinel

lherzolites were ultrasonically cleaned in weak HCl for 20 minutes to remove surface

contamination, subsequently rinsed with de-ionised water and ultrasonically washed for 30

minutes in de-ionised water. The lunar samples include the more important magmatic rock

types collected during various Apollo missions: high-Ti mare basalts, picritic glasses, olivine-

,ilmenite- and quartz-normative basalts as well as one ferroan anorthosite. In addition, 2

martian meteorites, 6 eucrites and 2 chondrites were analysed.

Lithium isotope compositions were calculated using the following formula: S7Li (%o) =

[(7Li/6Li)sampic/(7Li/6Li)standard-l]xlOOO. We used L-SVEC as a reference standard (FLESCH et

al., 1973). Accuracy and precision of the method has been monitored by multiple analyses

(full replication including sample dissolution) of BHVO-2 and JB-2, two international

reference rocks of the US and Japanese Geological Surveys, and OK-1 an in-house eclogite

standard from Miinchberg (Germany) with well-defined 57Li = -8.27 + 0.20%o (Magna et al.,

2004b). During this study 87Li values of 4.52 ± 0.18%o (2a; n=10), 4.65 + 0.39%o (2a; n-9)

and -8.24 + 0.33%o (2a; n=3) have been obtained for BHVO-2, JB-2 and OK-1, respectively.

These values are in good agreement with previously published data and results reported by

other groups (Moriguti and Nakamura, 1998a; Tomascak et al., 1999a; Jeffcoate et al.,

2004).

91


5000

4000

3000

2000

1000

Ü

w 20000

a.o

OS 10 15 20 26 30 35 <10 45 50 55 6.0

ml

Fig. 5.1: Two-stage Li elution and purification from clinopyroxene 8520-09, Panel A displays first step of hi

elution where Cr dominates hi fraction. Note two-order difference in counts per second (cps) for Cr and hi.

BioRad AG 50W-K8 cation-exchange resin was used and elution media IM HNOj in 80% (v/v) methanol

provided excellent separation ofhi from other matrix and trace elements apartfrom Cr. Second step (panel B)

used 0.5M HNO] to separate hi from Cr successfully with employing BioRad AG 50W-X12 resin. Chromium is

marked as squares with doited tine, lithium as diamonds with solid line.

5.4 Results

Lithium abundances and isotope compositions for non-terrestrial samples are given in

Table 5.1. Trace element data for non-terrestrial samples are provided in Table 5.2. Two

chondrites Allende (CV3) and Bruderheim (L6) contain 2.1 and 2.7 ppm Li, respectively.

Nakhla, a martian clinopyroxene cumulate, and Zagami, a martian shergottitic basalt, contain

4.8 and 4.0 ppm Li, respectively. Eucrites show substantially higher Li abundances (8.3-13.1

ppm) as do lunar rocks and glasses (6.0-13.3 ppm). Two ferroan anorthosites have 0.8 ppm Li

(62255) and 0.14 ppm Li (65315), respectively. Total repeat analyses of five samples

including the full analytical procedure replicated very well (see Table 5.1).

92


Table 5.1 : Lithium concentrations and isotope compositions of non-terrestrial samples

sample split Li (ppm) 57Li (%o) 2a

Moon

Anorthosites

Fe-anorthositcs 62255 134 0.8 8.89 0.18

65315 30 0.14 n.a.

Basalts

High-Ti marc basalts 70035 164 8.6 5.09 0.09

replicate 5.11 0.10

74255 178 9.2 6.39 0.41

75075 158 10.3 5.48 0.15

replicate 5.21 0.31

Qz-nonnativc basalts 15058 7.7 3.74 0.12

15475 31 7.2 3.35 0.26

Ilm-normative basalt 12045 13 9.5 4.43 0.09

replicate 4.28 0.07

Ol-normative basalt 15555 115 6.0 4.32 0.16

Glasses

Green glass 15426 163 6.7 4.74 0.19

Orange glass 74220 804 13.3 4.19 0.30

Grey glass 74240 17 10.5 5.60 0.15

Mars

Nakhla 4.8 4.99 0.47

Zagami 4.0 3.88 0.18

Vesta

Bereba 10.6 4.23 0.22

Bouvantc 13.1 3.84 0.48

Cachari 10.5 5.75 0.89

Juvinas 9.0 3.65 0.29

replicate 3.61 0.34

Pasamonte 10.1 3.65 0.20

ALHA 78132 8.3 3.04 0.24

Chondrites

Allende 2.1 3.13 0.25

replicate 3.08 0.19

Bruderheim 2.7 1.49 0.29

n.a. - not analysed

93


Table 5.2: Measured trace elements concentrations of non-terrestrial samples12045 15058 15426 15475 15555 62255 70035 74220 74240 74255 75075

Li 9.5 7.7 6.7 12 6.0 0.81 8.6 13.3 10.5 9.2 10.3

Co 55 44 69 114 67 1.9 17 66 31 25 25

Ni 55 32 195 1341 67 9.2 6.6 72 94 9.0 8.2

Cu 19 10 10 11 25 - 44 36 32 40 43

Ga 4.8 3.6 3.5 4.1 3.9 - 4.1 8.7 8.1 5.2 5.3

Rb 1.2 0.88 1.3 0.92 0.93 0.33 1.2 1.5 2.3 2.2 1.3

Sr 149 98 65 91 44 170 89 215 152 181 183

Y 53 24 24 23 27 0.35 82 49 76 97 81

Zr 124 73 88 63 89 2.6 231 173 216 246 211

Nb 7.1 5.2 6.6 5.0 6.9 0.10 23 16 19 25 23

Cs 0.04 0.04 0.04 0.04 0.04 0.003 0.03 0.05 0.10 0.07 0.03

Ba 61 50 72 45 47 8.6 54 85 103 85 72

Hf 3.9 1.9 2.1 1.5 2.3 0.03 7.6 5.2 6.3 8.0 7.0

Pb 0.28 0.20 0.70 0.22 0.22 0.57 0.16 0.93 1.0 0.28 0.21

U 0.17 0.11 0.25 0.10 0.13 0.002 0.09 0.15 0.37 0.16 0.09

La 5.9 4.5 6.0 4.0 5.0 0.17 4.4 6.3 9.0 7.2 5.2

Ce 17 12 15 11 13 0.44 15 19 26 24 17

Pr 3.0 1.9 2.2 1.6 2.1 0.05 3.1 3.4 4.3 4.6 3.4

Nd 16 9.7 10 8.2 11 0.23 18 19 22 28 20

Sm 6.1 3.1 3.0 2.8 3.6 0.10 8.0 6.8 8.5 11 8.1

Eu 1.5 0.93 0.60 0.78 0.58 0.67 1.2 2.0 1.6 2.1 2.0

Gd 7.7 3.8 3.5 3.5 4.4 0.06 11 8.2 10 14 11

Tb 1.6 0.71 0.67 0.63 0.84 0.01 2.3 1.6 2.1 2.8 2.3

Dy 9.5 4.2 4.2 4.0 4.8 0.05 14 9.2 13 17 14

Ho 2.2 0.91 0.90 0.88 1.0 0.01 3.2 1.9 2.9 3.7 3.1

Er 5.6 2.4 2.4 2.4 2.7 0.04 8.7 4.7 7.8 9.8 8.4

I'm 0.84 0.36 0.39 0.36 0.38 0.005 1.4 0.70 1.2 1.4 1.2

Yb 5.5 2.3 2.6 2.4 2.4 0.03 8.8 4.4 7.9 9.5 8.2

Lu 0.77 0.34 0.39 0.33 0.37 0.005 1.3 0.64 1.2 1.3 1.2

Allende Bruderheim Bereha Bornante Cachari Juvinas Nakhla Pasamonte Zagami ALHA 78132

Li 2.1 2.7 10.6 13.1 10.5 9.0 4.8 10.1 4.0 8.3

Sc 29.1 29.3 33.7 28.4 54.7 29.3 62.6 26.7

Co 682 9.0 3.7 9.9 7.6 55 7.5 43 7.8

Ni 14600 5.1 4.7 5.0 4.9 91 22 82 20

Cu 117 3.8 3.7 6.3 4.7 6.6 2.3 15 1.8

Ca 7.5 2.4 3.0 2.1 2.9 4.2 2.9 15 2.5

Rb 1.3 0.40 0.76 1.2 0.33 4.1 0.40 5.7 0.42

Sr 15 78 91 101 78 55 79 35 73

Y 2.7 17 29 21 16 3.7 17 13 16

Zr 16 41 87 49 48 12 48 65 45

Nb 2.3 3.9 7.3 5.1 4.2 1.8 3.9 4.9 3.9

Cs 0.09 0.01 0.03 0.03 0.02 0.53 0.01 0.37 0.01

Ba 5.3 32 57 86 30 26 32 25 31

Hf 0.33 0.95 2.2 1.2 1.2 0.28 1.1 1.7 1.0

Pb 2.1 0.18 0.47 0.26 2.30 0.58 0.31 1.2 0.38

II 0.02 0.11 0.17 0.28 0.09 0.05 0.09 0.07 0.12

La 0.58 3.0 5.7 3.8 2.9 2.0 3.0 1.4 3.0

Ce 1.4 7.6 14.0 9.4 7.1 5.1 7.4 3.3 8.2

Pr 0.23 1.2 2.2 1.5 1.1 0.80 1.2 0.52 1.3

Nd 1.2 5.9 10.7 7.6 5.4 3.7 5.7 2.8 6.2

Sm 0.43 1.9 3.5 2.6 1.7 0.82 0.86 1.1 1.8

Eu 0.12 0.70 0.93 0.75 0.70 0.25 0.66 0.44 0.68

Gd 0.40 2.3 4.3 3.0 2.4 0.77 2.4 1.6 2.3

Tb 0.08 0.48 0.79 0.56 0.43 0.12 0.46 0.35 0.45

Dy 0.47 3.1 5.0 3.6 2.7 0.71 3.0 2.2 2.7

Ho 0.12 0.67 1.1 0.79 0.64 0.15 0.66 0.51 0.64

Er 0.28 1.8 3.0 2.3 1.8 0.43 1.8 1.4 1.8

Tm 0.06 0.29 0.42 0.34 0.28 0.07 0.27 0.22 0.28

Yb 0.35 1.9 3.0 2.3 1.8 0.38 1.8 1.4 1.7

Lu 0.05 0.29 0.41 0.34 0.26 0.07 0.29 0.20 0.26

All trace element data are listed in parts per million (ppm). Precision of the concentration data for Ec-anorthosite 62255 is lowered to

20-40% (2a) due to very low concentrations of most trace elements.

94


The Li isotope compositions of the samples range from 41.5 to +8.9%o 57Li (Fig. 5.2).

The chondrites are isotopically light relative to all the other samples, Bruderheim (L6)

yielding ô7Li = +1.5%o, whereas two duplicate measurements of Allende (CV3) average ô7Li

= +3.1 %o. Although this is an inadequate characterisation of chondrites, the data are fully

consistent with the recent and more comprehensive study of chondrites reported by

Mcdonough et al. (2003), who also report compositions that are variable but light relative to

the values for the various silicate samples reported in Table 5.1. The best estimate for average

chondritic compositions is 87Li = +2.0 J- 1.5%o (W. McDonough, personal communication to

ANH, 2005). The martian meteorites Nakhla and Zagami yield S7Li = +5.0%o and +3.9%o,

respectively. Eucrites span a narrow range from +3.0%o to +4.2%o excluding Cachari with a

high 57Li. The same aliquot of Cachari also has anomalously heavy oxygen (Wiechert et al.,

2004). The range of Ô7Li values in lunar basalts and glasses (i.e. excluding the anorthosites)

exceeds 3%o ( i 3.4 to +6.4%o) and appears to vary with lithology. Quartz-normative basalts

have the lightest Li isotope compositions whereas high-Ti mare basalts are isotopically heavy.

Picritic glasses, olivine- and ilmenite-normative basalts have intermediate ö Li values. An

extreme Ô7Li value of+8.9%o has been measured in a ferroan anorthosite (62255).

Lithium abundances and isotope compositions of terrestrial samples are given in Table

5.3. Terrestrial basalts, e.g. MORB and OIB, are variable in Li concentrations but most have

3-6 ppm Li consistent with the comprehensive study of Ryan and Langmuir (1987).

Olivines from spinel lherzolites are notably homogeneous with 1.5-2.2 ppm Li. Pyroxenes are

depleted in Li compared to olivines, and orthopyroxenes are more depleted than

clinopyroxenes, consistent with findings of Seitz and WOODLAND (2000) for peridotitic

lithologies. All olivines and bulk peridotites have Ô7Li values indistinguishable from the

currently accepted mantle 87Li of 4.2 + 0.8%o (Chan, 2003). Olivine separates are

particularly homogeneous with Ô7Li ranging from 3.6 to 3.8%o. Whole-rock analyses of

peridotites from Mongolia show constant Li abundances of 1.4-1.7 ppm Li and a remarkably

narrow range of S7Li = 3.8-4.3%o. The 87Li of the harzburgite is slightly lighter (by ~0.4%o)

than the values for spinel lherzolites.

95


-10123456789 10

Earth*> O owwes^sp Ihenolites)

' -O 8HVO-2

i • t JB-2V

rFe-arwrltiraile . S2MS * -Ù '

O,

; 70035

Of

> - O - i 74255 tiigh-Tt basalte

'O"

75075

, o .

Moon <

'O- 15058 ,

\ qz-normauw basalts

t-O i 15475 '

&

12045 . ilm-nofmalwe basaJl

o

KH 15555 ol-nomiatwe tsasaft

i O-f 16426(911811)

O • 74220 (Drang») gla&ses

LO 74240 (gtey)

Marsi -TJ- i Nakhla

t-D-. Zagami

r>—Q— Bereûa

Q pouvante

h q , Cadran

Vesta < i O- i

,> Juvinas

> Pasafnonte

k. Q ( MHA 78112

Attende '

'<

+ i Brudeffieim

Chondrites

-3 -1 0 1 1023456789

57Li (%o)

Fig. 5.2: Lithium isotope composition of terrestrial and extraterrestrial materials. MORB range is fromTOMASCAK (2004) and ELLIOTT et al. (2004), Iceland range isfrom PlSTlNHR and HENDERSON (2003), RYAN and

KYLE (2004) and high-3He/4He samples from MAGNA et al (2004a), Hawaiifrom CHAN and FREY (2003), Shi

range of chondrites is from MCDONOUGH et al. (2003) and JAMES and PALMER (2000). Olivines from spinetlherzolites are from this study. BHVO-2 and JB-2 are displayed as long-term reproducibility standards (see

Chapter 2 and MAGNA et al,, 2004b), Apparent variations of cfhi among lunar rocks cannot be related to

possible surface exposure to isotopically heavy solar wind particles, even not in case of low-Li ferroananorthosile 62255 due to its extremely short exposure history <2Ma. Grey glass 74240 comes from the same

location as orange glass 74220 and is very similar in chemical composition to this spatially related orange glass

(apart from slightly elevated Na contents; RHODES et al, 1974; KOROTEV and KliEMSFR, 1992). harge

proportions of basaltic fragments and grey "ropy" glass spherules cause macroscopically grey colour, orange

glass component represents only a very small portion ofgrey glass. Clinopyroxene is the most abundant mineral

phase. It has been suggested that the grey glass is a product of a local large highland impact (HEIKEN and

MCKAY, 1974).

96


In contrast to these uniform values, pyroxenes show a considerable range in ö7Li from

2.6 to 6.0%o and two reference peridotites measured during the course of this study show

completely distinct ö7Li values. JP-1 (GSJ) has S7Li = 3.97 + 0.41%0, consistent with the

olivine and Mongolian bulk peridotite data, whereas PCC-1 (USGS) yields 57Li = 8.85 +

0.41%o (see Table 5.3) which is indistinguishable from a value of 8.9%o for this reference rock

given by Seitz et al. (2004). However, PCC-1 is a partially serpentinised harzburgite, which

can readily account for its heavier 87Li value compared to the other peridotites. Bulk Li

isotope compositions of peridotitic rocks must be considered with caution as they may be

influenced by metasomatism and alteration at various possible stages.

Table 5.3: Lithium concentrations and isotope compositions of terrestrial samples

Li (ppm) 5'Li (%o) 2a

Peridotites - minerals

San Carlos ol

San Carlos opx

Vitim 314-58 ol

Vitim 314-58 opx

Vitim 314-58 epx

Atsagin-Dush 8520-09 ol

Atsagin-Dush 8520-09 epx

Tariat MPH 79/1 ol

Tariat MHP 79/1 epx

Kilbourne Hole 96-2 ol

Peridotites - whole rocks

Atsagin-Dush 8520-12 wr




Peridotites - standards

PCC-1 (USGS)

JP-1 (GSJ)

1.6 3.64 0.15

0.52 3.32 0.35

1.9 3.76 0.05

0.39 3.78 0.15

1.4 5.21 0.23

1.5 3.55 0.15

0.82 5.96 0.26

2.0 3.75 0.04

1.4 2.57 0.32

1.7 3.56 0.39

1.6 4.11 0.50

1.7 4.23 0.25

1.6 4.27 0.07

1.4 3.77 0.40

1.1 8.85 0.41

1.8 3.97 0.41

ol - olivine; opx - orthopyroxene; epx - clinopyroxene; wr - whole rock

5.5 Discussion

5.5.7 Lunar Li and the isotope effects ofexposure

Among the various analysed planetary bodies most Li isotope variation is found in lunar

materials (Table 5.1). This could of course reflect the fact that we have analysed many more

97

Chapter 5 Terres trial planets

lunar samples. The variability on the Moon is still surprising however. Oxygen isotopes

provide evidence that the Earth and Moon formed from the same raw material (Wiechert et

al., 2001). Therefore, the Li isotope compositions for the total Earth and Moon should be

identical, assuming Li isotopes have not been fractionated by the putative Moon-forming

Giant Impact. Lithium isotope fractionation on Earth is dominated by the large effects

associated with the hydrosphere. Therefore, Li isotope fractionation within the dry Moon is

expected to be much smaller with basaltic compositions similar to the best estimate for the

Earth's mantle (fi7Li ~ +4%o). In fact lunar basalts might offer greater insights into the Li

isotope composition of terrestrial planetary mantles because of the anticipated absence of

large fractionations.

The overall range of Li isotope composition among lunar samples is, however,

surprisingly broad. The most extreme composition found among the various lunar samples is

for the lunar highlands anorthosite 62255 (57Li = +8.9%o). The issue arises as to whether this

spread might be caused by secondary irradiation / implantation. Solar wind (SW) and galactic

cosmic rays (GCR) may alter those rocks at the lunar surface not protected by a magnetic

field. The bulk Sun has 7Li/6Li ~106 (Delbourgo-Salvador et al., 1985) and solar wind has

a 7Li/6Li -31 (Cuaussidon and Robert, 1999). 7Li/6Li -2 have been reported in galactic-

cosmic-ray-induced spallation reactions (Eugster and Bernas, 1971), and the interstellar

medium 7Li/6Li has been estimated at ~I2 (Knauth et al., 2003), and ~2 (Knauth ct al.,

2000).

Cosmogenic effects would produce roughly equal amounts of f'Li and 7Li by fission of

heavy nuclides because spallation production Li/'Li ratio can be as low as ~1.0 (e.g.

(Lemoine et al., 1998). Rather higher values of ~2 were measured for low-energy galactic

cosmic rays (Ramaty et al., 1996). Spallogenic Li has been detected in lunar soils yielding

values of 7Li/6Li - 2 (Eugster and Bfrnas, 1971). Therefore, Li ofpurely spallogenic origin

is extremely light ö7Li ~ -850%o (Chaussidon et al., 2001). This component has been

detected in small quantities in refractory inclusions from several meteorites (Chaussidon and

Robert, 1996; Chaussidon et al., 2001) and in four lunar soils studied with depth profiling

(CHAUSSIDON and ROBERT, 1999). It is impossible that spallogenic effects arc responsible for

Li enrichment as detected in lunar samples.

An effect from exposure to heavy solar wind (CHAUSSIDON and ROBERT, 1999) would

of course generate higher Li/'Li. However, the depth of SW penetration into regolith

particles is on the order of just tenths of microns (Chaussidon and Robert, 1999).

98


Therefore, this effect is generally expected to be very small. More accurately and

independently from grain size and solar flux one can calculate the proportion of solar-wind-

derived Li using 36Ar abundances. Elevated 6Ar abundances are usually found in fine-grained

fractions of lunar soils with up to 5-fold " 6Ar enrichment in the finest fractions relative to

coarser-grained fractions (ASSONOV et al., 2002). An abundance of 0.32* 1016 atoms / gram

was reported for the 25-42um fraction of the lunar soil 67601 (Wieler and Baur, 1995). By

applying the enrichment factor of 5 to the <25um fraction we obtain [ Ar] of 1.6x10 atoms

/ gram. The lunar soil 67601 <25 micron fraction does in fact yield an Ar concentration of

60000x10s ce / g or 1.6xl016 atoms / gram, in excellent agreement with this prediction (R.

Wieler, personal communication to TM and UW, 2005). The solar abundances of 7Li and v'Ar

normalized to 106 silicon atoms (Anders and Grevesse, 1989) are 52.8 and 8.5xl04

respectively. From these a solar Li/ 6Ar of 6.2x10" can be calculated. Using the measured

Ar abundances and solar Li/ 'Ar one obtains 9.9x 10 of Li atoms / gram derived from the

solar wind in the <25u,m fraction. This number has been multiplied by a factor of 3 in order to

account for lower ionization energy of Li compared with argon (FIP factor). This gives

3xl0'3 of 7Li atoms per gram that is the expected portion implanted by the solar wind

(fraction <25u.m of sample 67601). An average of 5.2 ppm Li of the 3 coarse fractions of

67601 is used to calculate the number of lunar 2*1016 Li atoms / gram in the <25um grain

size fraction. The solar wind portion of Li corresponds only to about 0.15% of the Li in

fraction <25um micron. On the basis of a solar wind composition of 7Li/'Li = 31 and

assuming no fractionation by the implantation process a Ô Li = +7.6%o is calculated for the

<25pm fraction based solely on measured 36Ar abundances and the lunar lithium

concentration. This calculation provides strong evidence that solar wind is detectable in grain

size fractions <25um of lunar soil using MC-ICPMS.

To test this theoretical approach we have conducted independent studies of the extent of

the effects of solar wind implantation on lunar soils. The Li abundances and isotope

compositions have been measured for size fractions from immature lunar soil 67601 (Table

5.4), which has an exposure age of ~55 Myrs, and in bulk soil 64421 with a higher exposure

age of -210 Myrs (KIRSTEN et al., 1973). We find clear 7Li enrichment in soil 64421 (Ô7Li =

8.3%o) and in the finer-grained fractions of 67601 (Figure 5.3 and Table 5.4) but this effect

diminishes rapidly with increasing grain size. The three coarser fractions (>42um) of soil

67601 have uniform 8 Li ~ +5%o. The measured S7Li = 8.4 %o for the fraction <25u,m is in

99


good agreement with predicted value in the above calculations, given the range of

uncertainties. This provides strong evidence for solar-wind-derived Li in lunar soils. The

explanation for such trend toward heavy S7Li values could be a large proportion of abraded

o

to

<25 25-42 42-90 90-250

Size fraction (jam)

>250

Fig. 5.3: Size fractions vs. c?Li in Apollo 16 soil 67601. Apparent trend toward heavier ifhi values with finer

fractions of the North Ray crater soil reflects most probably implantation of isotopically heavy solar wind

(CHAUSSIDON and Robert, 1999). Dashed line is the average ofthree coarse-grained fractions with Shi = 5.20

1 0.46%o (2a; grey bar) that may reflect the starting isotope composition.

surface layers of soil particles that are enriched in 7Li. This effect is insignificant in coarse¬

grained fractions with volumetrically negligible quantities of such eroded material. Therefore,

bulk rock samples of igneous material, even with long exposure age should show no

significant effects and the two components (solar wind, spallogenic reactions) can be

neglected when interpreting the data set of lunar rocks from this study.

Table 5.4: Lithium concentrations and isotope compositions of lunar soils

sample split fraction (um) Li (ppm) 67Li (%o)

67601 7 <25 5.4 8.44

25-42 5.6 7.03

42-90 5.2 5.35

90-250 5.2 4.93

>250 4.7 5.31

64421 5 7.4 8.26

100


This is confirmed by the observation that the Li isotope compositions of the bulk rocks

studied here display no correlation with exposure age (Figure 5.4). In particular the very

heavy 57Li measured in ferroan anorthosite 62255 is accompanied by an extremely short ex

posure on the lunar surface with Kr exposure age of <2 Myrs (Drozd et al., 1977). In some

respects this is not surprising since neither implantation of solar wind nor spallogenic effects

can account for the apparent relationships between Li isotope composition and rock type

(Table 5.1). Similarly, it would be hard to explain the correlations with trace element indices

of magmatic differentiation, discussed in the next section.

IU.U

9,0 362255

8.0

70

e—>. 6 0*

742S'>

ê

to

5.074240

70035'

75075

4.0»

-A7422Ü

15555

15426

15058

3.0 15475

2.0

1.0

n n

50 100 150 200 250

Exposure age (Ma)

300 350

Fig. 5.4: Exposure ages vs. Shi for lunar samples. Closed diamonds • high-Ti mare basalts; triangles -

volcanic glasses (open orange glass 74220, dark grey-

grey glass 74240, tight grey - green glass 15426),

polygon - olivine-normative basalt 15555, closed circles - quartz-normative basalts, open square -ferroan

anorthosite 62255. No systematic trend is observed among the data. Exposure ages are taken from DROZD et al.

(1977), EUGSTER et al. (1977), HÖRZ et al. (1975), HUNEKE et al. (1973), STETTLER et al. (1973) and ÏOKOYAMA

etal (1974).

5.5.2 Lithium isotope composition ofthe Moon

A thorough model and review of geochemical and magmatic evolution of the lunar

magmasphere has been given elsewhere (TAYLOR and JAKES, 1974; SHEARER and PAPIKE,

1999). Lithium isotope variation among lunar basalts most likely reflects heterogeneity in the

source. The mare basalts analysed have primitive major element compositions and probably

101


segregated only relatively small amounts of olivine and/or pyroxene. This is unlikely to have

caused a detectable change in 57Li (Tomascak et al., 1999b). The mare basalts are the

products of melting of cumulate sources that crystallised from the Lunar Magma Ocean

(LMO) at distinct stages. All investigated mare basalts show a negative Eu anomaly (Figure

5.5) providing evidence that the cumulate sources were formed after significant crystallisation

of plagioclase in the LMO. It is believed that floating plagioclase formed the early lunar crust

(TAYLOR and JAKES, 1974). The Eun/Eu* value provides a monitor of how much plagioclase

crystallised from the magma ocean. Early cumulates have REE patterns with virtually no Eu

anomaly (Shearer and Papike, 1989). Later cumulates should develop a REE pattern with

negative Eu anomalies and be systematically depicted in Ni and Co and enriched in

incompatible elements. The inverse trend of EuN/Eu* vs. ô7Li for lunar mare basalts (Figure

5.6) provides evidence of a relation between magmatic differentiation of the LMO and Li

isotope fractionation. However, the trend is scattered and cannot be explained by feldspar

fractionation alone. Because the Li content of an anorthosite mainly composed of calcic

plagioclase is very low (see Table 5.1), the overall contribution of the lunar crust to the Li

budget of the Moon is particularly limited. Therefore, plagioclase segregation is an unlikely

mechanism for generating a significant Li isotope effect.

The sources of the low-Ti mare basalts were composed of early-formed ultramafic

cumulates (predominantly olivine and low-Ca orthopyroxene). Flotation of plagioclase

imposes a negative Eu anomaly on the crystallising mafic cumulates and produces a

complementary positive Eu anomaly in the feldspathic lunar crust. Most of the models for the

generation of primary high-Ti basalts require the melting of source rock types consisting of

olivine, clinopyroxene, ilmenite ± plagioclase + orthopyroxene (Snyder et al., 1990;

SHEARER and PAPIKE, 1993). Accurate data of relevant mineral-melt fractionation factors

(amjn-mcit) are currently unavailable for Li isotopes but the data on mantle minerals reported in

this study provide evidence of ~2%o equilibrium fractionation between olivine and

clinopyroxene at magmatic temperatures. The dominance of clinopyroxene in the source of

high-Ti mare basalts and the ultramafic olivine-dominated cumulates of the low-Ti mare

basalts could already account for most of the 67Li range detected in marc basalts. This is fully

consistent with trends between S7Li and REE ratios or Zr/Hf (Figure 5.6) because

clinopyroxene and to a lesser extent titanite are the major Zr/Hf-fractionating phases (David

et al., 2000; Weyer et al., 2003).

102

Chapter 5 terrestrial planets

100

Fig. 5.5: CI -normalised REE patterns of lunar rocks Symbols are the same as in Figure 5.4. Grey squares-

ilmenite-normative basalt 12045. Also REEpatternfor Fe-anorthosite 65315 is plotted. Note apparent scatter in

trace element data for both anorthosites caused by larger analytical uncertainties. Bottom pattern of Fe-

anorthosite 65315 is approximative only due to analytical uncertainties for most elements exceeding 50% (2a).

Strong correlations between Li isotopes and trace elements arc found for indices that

span a greater relative range of variability in basalts. For example Li/Yb and Ga/Hf vary

widely in the basalts and clearly correlate inversely with ô7Li (Figure 5.6). Gallium and

lithium are considered to be at most only slightly incompatible and no phase fractionates

them strongly. This is endorsed by the basalts, which yield Ga and Li concentrations that vary

by a factor of two at most (Tables 5.1 and 5.2). In contrast the REEs and high field strength

elements (HFSEs) vary in concentration by more than a factor of four. Therefore, the

variation in Ga/Hf and Li/Yb is actually dominated by a large degree of variability in HFSEs

and REEs. The samples with lowest Ga/Hf and Li/Yb are the Apollo 17 glasses and high-Ti

mare basalts. This is fully in line with a source consisting mainly of late clinopyroxene-rich

cumulates. These samples are also more enriched in alkalis leading to a positive correlation

between Rb concentration and 57Li (Figure 5.6) and therefore, clearly more evolved than low-

Ti basalts. On the basis of a study of a terrestrial volcanic rock sequence it has been stated

earlier that magmatic differentiation has no effect on Ô7Li (Tomascak et al., 1999b).

103


However, Tomascak ct al. (1999b) explored only some ten percent segregation of olivine ±

pyroxene + plagioclase whereas the sources of high-Ti basalts are thought being formed after

>95% of the LMO had been crystallised. That such extensive magmatic differentiation can

fractionate stable isotopes has been shown with oxygen isotopes. For example, a significant

shift has been produced by olivine segregation in highly evolved melts (MgOmdt<2%; Eiler,

2001). Assuming a fractionation factor a0iivme-meit = 0.999 between olivine and silicate melt

(as inferred from olivine measurements in this study) a simple Rayleigh fractionation model

produces significant Li isotope fractionation. If 95% of the LMO is crystallised, the ö7Li

value of the residual melt increases by more than 2.5%o compared with the starting

composition. This may in addition contribute to the 57Li range observed for the lunar basalts.

The lunar crust is about ~60km thick and consists mainly of anorthosite that is

complementary to the source of mare basalts. This is also reflected in the REE pattern (Figure

3 in Tayeor and Jakes, 1974). Lithium concentration of 0.8 ppm for the studied anorthosite

implies a Li concentration of about 4 ppm for the magma ocean assuming an element

partition coefficient for plagioclase and basaltic melt Dpiag-mcii of ~0.2 (Bindeman et al.,

1998). The S7Li value of 8.9%o for the lunar anorthosite is somewhat surprising. This gives a

fractionation factor between plagioclase and basaltic melt apiag.mcit of 1.0049 (assuming a 87Li

value of 4.0%o for the lunar magma ocean) and represents an extreme value for equilibrium

fractionation at temperatures >1150°C. Whether the high S7Li of the lunar crust reflects

equilibrium fractionation by plagioclase crystallisation is unclear.

It is possible that the Li isotope composition of 62255 is not in equilibrium with the

LMO. The Li concentration of 0.8 ppm measured for this sample is much higher than the 0.14

ppm obtained for a second Fe-anorthosite 65315 which is believed to be one of the most

pristine lunar highland crust samples available as deduced also from low C1-normalised REE

patterns (Figure 5.3 and Ebiiiara et al., 1992). Such low Li abundances as in 65315 would

mean approximately 0.8 ppm Li for the LMO if Dpiag_meit ~0-2 but such a low Li content of the

LMO is highly unlikely. The possibility that the high ô7Li value of 62255 reflects Li derived

from late fluids or melts that percolated through the lunar crust cannot be excluded.

Unfortunately, the amount of Li extracted from 65315 was too small for an isotope ratio

measurement. However, the extremely low Li abundances arc hard to reconcile with the

conclusions of DREIBUS et al. (2003) who suggest preferential Li incorporation into a

plagioclase crystallising very early from the LMO.

104


3.0 r-

2.5

— 2.0 t

S 1-5

LX. -in

0.5

4.0

3.5 r

3.0

2.5

2.0

1.5

1.0

0.5

1.0

0.8 -

'ë °-6

l3 0.4

0.2

0.0 L-J-J-

74220A

2045

• • O1M/S

1MS81SS5i

74740 +742SS

5542G

4 75075

7003*

A

A

D

A

• A

O

A

SOS« o15&&5

>^b7424G

74720

12043~ - +«»35" "7SS

K075

_l 1 *_..i....i.....A..„J.„

A3 A

D A

A3

O

« A

XtN

50

40

35

30

3.0

2.5

2.0

1.5 laO

1.0

0.5

1000

800

600 n

QC

-i 400

- 200

2.0 3.0 4.0 5.0 6.0 7.0 8.0 3.0 4.0 5.0 6.0 7.0 8.0

S7Li (%o) Ö7Ü (%o)

Fig. 5.6: Indices of magmatic fractionation vs. cfLi in lunar basalts and glasses. Symbols are the same as in

Figure 5.4. Ferroan anorthosite 62255 is not plotted. Variations in Rb abundances, hi/Yb, EuN/Eu , Ga/Hfand

Zr/Hf ratios are related to precipitation ofclinopyroxene, ilmenite, armalcolite, spinel and titanite andpossibly

anorthile from the hMO. Excess Eu in 62255 (EuN/Eu >25) relative to other rocks is caused by Eu substitution

for Ca in anorthosite that is nearly pure anorthile (>97%). K abundance data are from PaI'IKE et al. (1974),

RHODES and Hubbard (1973), Rhodes el al. (1974), Rose et at. (1974; 1975) and TAYLOR et al. (1973).

The ô Li value of the bulk Moon can be estimated from basaltic lavas thought to be

derived early from the LMO crystallised cumulates. Among the first reservoirs that separated

105


from the LMO are the sources of two quartz-normative basalts (15058, 15475) and an orange

glass (74220) with Eun/Eu as high as -0.8. These three samples have ô Li values of 3.7 ±

0.1 %o, 3.4 + 0.3%o and 4.2 + 0.3%o, respectively. One of the most primitive and little

fractionated material available from the Moon is often considered to be the green picritic

glass 15426 released from depths >700km (Longhi, 1992). However, relatively high Ti

content of this glass provides evidence for late stage mixing with magmas derived from

clinopyroxene- and ilmenite-rich cumulates. This is consistent with a slightly elevated 5 Li

(4.7 ± 0.2%o) compared with the former samples. Finally, olivine basalt 15555 is viewed by

some as being representative of the Nd and W isotope composition of pristine, unfractionated

lunar mantle (Nyquist et al., 1995; Kleine et al., 2003). But the large negative Eu anomaly

indicates equilibration with a more evolved lunar magma ocean and its Li isotope

composition (57Li = 4.3 ± 0.2%o) may be slightly affected by magmatic segregation of

olivine. Therefore the mean of the two quartz-normative basalts and one orange glass sample,

taken as providing the most reliable estimate of the bulk composition of the Moon, is 5 Li =

+3.8±0.4%o(la).

5.5.3 Lithium isotope composition ofthe terrestrial mantle

In order to constrain the Li isotope composition of the Earth's mantle, we studied

mineral separates of spinel lherzolites and whole-rock peridotites. Spinel lherzolites probably

best represent the Li isotope composition of the upper mantle. Modal proportions of olivines

in spinel lherzolites from this study range from 47 to 62 modal% (San Carlos: 47.2%

(KÖHLER and BREY, 1990), Vitim: 62.3% (IONOV, 2004), Atsagin-Dush: 60.6% (Wiechert

et al., 1997), Tariat: 57.1% (Press et al., 1986), Kilbourne Hole: 57%). Therefore, olivine

contains between 80 and 90% of bulk sample Li and is a good approximation for the Li

isotope composition of peridotitic mantle rocks. The average S Li of olivines from spinel

lherzolites is 3.65 ± 0.20%o. Whole-rock analyses of three spinel lherzolites (8520-12, 8520-

19, 8520-20) and one harzburgite (8520-30) are also consistent with a Ô Li value of the

terrestrial mantle close to +4%o (see Figure 5.7). These values are similar to the calculated

bulk 57Li for fertile Vitim garnet lherzolite of 13.8%o (SEITZ et al., 2004). This is also similar

to data for peridotites from Zabargad Island (Egypt), which yield ö7Li «4-5%o for the

terrestrial mantle (CHAN et al., 2002b; BROOKER et al., 2004). All these results are consistent

with the predicted upper mantle value of close to +4%o (CHAN, 2003).

106


In contrast, a whole-rock analysis of a fertile San Carlos spinel lherzolite given by SEITZ

et al. (2004) gave light 57Li = -l.l%o. The whole-rock data for this San Carlos sample appear

inconsistent with the Li concentration and isotope composition determined from mass balance

of the three major minerals analysed. Therefore, it would appear that some poorly understood

effect is biasing the result from the whole rock. Our Li isotope data for mineral separates

provide some evidence of heavier Ô Li in clinopyroxene although more data are needed to

place better constraints on inter-mineral fractionation. The extensive study of Seitz et al.

(2004) yielded different results for reasons that are presently unclear. Olivines from their

study are heavier in terms of Li isotope compositions than co-existing clinopyroxene and

orthopyroxene but span a substantial range of S7Li from 1.4 to 4.5%o. Co-existing

orthopyroxenes were isotopically lighter and clinopyroxenes were in most cases isotopically

the lightest major mineral. Although this could reflect variously depleted and metasomatised

mantle as suggested by Seitz et al. (2004), a more likely explanation for the observed light Li

isotope signatures in pyroxenes would encompass late stage diffusion. That such a process

can lead to significant shifts in S7Li has been shown for silicate melts (Richter et al., 2003)

as well as for solid-liquid systems, as documented in peridotitic lithologics where Ô Li

variations as large as 10%o were observed (Lundstrom et al., 2005). The latter study also

proved experimentally that Li isotope fractionation on the order of tens of per mil is readily

achievable in a very short-term timescale by diffusion and showed that scatter in ö7Li data for

clinopyroxenes is more pronounced than for olivines. This is consistent with fast Li diffusion

in clinopyroxene as recently observed (Coogan et al., 2005).

Elliott et al. (2004) sketched a possibility that heavy mantle Li may be caused by

viscous coupling of hydratcd isotopically heavy mantle with complementary dehydrated

subducted slab. Some heterogeneity in the mantle does indeed exist as sampled by MORBs

and OIBs at various locations. Fresh MORB analyses are sparse but show significant

heterogeneity in Ô7Li from +1.5 to 5.6%o (Chan et ai., 1992; Moriguti and Nakamura,

1998b; TOMASCAK and Langmuir, 1999) even greater than for OIB with 87Li values from

+2.5 to 5.8%o (Tomascak et al., 1999b; Chan and Frey, 2003; Magna et al., 2004a). OIBs

may be derived as unrepresentative small-volume melts from plumes rising from the core-

mantle boundary, ultimately fed by subducted slabs (Baker and Jensen, 2004). Thus, they

107


8.0

7.0

6.0

o

o1-

'3

to

5.0

4.0

3.0

2.0

1.0

0.0

A

.

D

'. O

a *T

V

0.0 0.5 1.0 1.5 2.0 2.5 3.0

Li (ppm)

Fig. 5.7: Li abundances vs. oLi ofperidotitic rocks and mineral separates. Closed symbols — olivines; grey

symbols - clinopyroxenes; open symbols - orthopyroxenes. Diamonds - San Carlos; squares - Vitim; triangles

- Atsagin-Dush; reversed triangles Tariat; circle - Kilbourne Hole. Solid star represents one harzburgite

(8520-30), open stars display spinel lherzolites (8520-12, 8520-19, 8520-20); all are from Atsagin-Dush,

Mongolia.

may not represent typical upper mantle compositions (ELLIOTT et al., 2004). MORBs repre¬

sent shallow mantle melting of depleted lherzolites with interspersed mafic veins (LUND-

STROM et al., 2000). As large volume melts they should be more representative than OIB.

However, the question arises of how much variably fractionated Li resulting from recycling is

affecting the compositions if metasomatism from recycled components is selectively affecting

the compositions ofpyroxene, a phase that is preferentially melted during basalt production.

To summarise, the existing data for the Earth's upper mantle are consistent with a 87Li

~+-4%o. Non-metasomatised peridotites appear to be strikingly uniform in their Li isotope

composition. Four such whole-rock peridotites as well as one peridotitic standard from this

study are extremely homogeneous in their S7Li (~4.1%o) varying at ±0.4%o (2a), consistent

with the upper mantle fi7Li at 4.2%o defined by Chan (2003). Similar results are found for

olivines of spinel lherzolites that average at +3.6%o and are identical with the mantle value.

108


We therefore suggest that the Earth's upper mantle has 57Li ~+4%o as inferred from our

olivine and whole-rock peridotite data. The Li content of the upper mantle as inferred from

non-metasomatised spinel lherzolites of this study is ~1.6 ppm and is in excellent agreement

with the estimate of Ryan and Langmuir (1987) and Sun and McDonough (1989).

Whether the upper mantle is also representative of the lower mantle is less clear. The

deep mantle may differ in 57Li as inferred from the slab dehydration and Li isotope fractio¬

nation model of Zack et al. (2003). Therefore, the possibility exists that the mantle is strati¬

fied with respect to Li isotopes. This would be caused by progressive 7Li depletion in the

deep mantle due to its loss from the dehydrated subducted slab. A metasomatic overprint in

clinopyroxenes has been shown in some mantle xenoliths from the Far East Russia and Japan

(Nishio et al., 2004) which carry extremely negative Ô7Li values inherited from metasoma¬

tism by eclogitic melts. The existence of light domains in the mantle is consistent with data

for clinopyroxenes from ultramafic xenoliths. Nishio et al. (2004) explained this feature by

metasomatic overprint with Li from subducted highly altered basalts that yielded extremely

negative values. Similarly, a relationship between light Li and assimilation of subducted crust

has been assumed on the basis of glass inclusions analysis in olivines and orthopyroxenes

from Iblean plateau tholeiites (GURENKO and SCHMINCKE, 2002). Recently, extremely

negative ô7Li values have been found in glass inclusions of olivines from Hawaiian lavas

(Kobayashi et al., 2004). Kobayashi et al. (2004) have attributed these light values to a

recycled component, supported also by B and Pb isotopes. Therefore, recycling does appear

to transfer a fractionated Li component into the mantle. What its average composition is

remains unclear. How different the bulk upper mantle and bulk lower mantle are as a result of

recycling is under-constrained.

5.5.4 Lithium isotope composition ofMars and Vesta

Two martian meteorites give distinct Ô Li of 3.9%o (Zagami) and 5.0%o (Nakhla).

Zagami is a basaltic martian meteorite and most likely reflects the composition of the source

material within the martian mantle, whereas Nakhla is a clinopyroxenite cumulate that closely

resembles shallow ultramafic intrusives (Harvey and McSween, 1992). Study of light

lithophile elements (Li, B) performed on pyroxenes from martian basalts (Zagami, Shergotty)

provided evidence for incompatibility of Li in major mineral phases of martian basalts which

has been explained by fluid degassing (Herd et al., 2004). The parental melt of Shergotty is

109


supposed to have high Li abundances >20 ppm based on recent mineral-melt partitioning data

estimated for martian pyroxenes (HERD et al., 2005). A similar study of light lithophile

elements distribution (Li, Be, B) in pyroxenes from nakhlites and shergottites led LENTZ et al.

(2001) to conclude that the parental magmas of shergottites had significant water contents

whereas the parental magmas of Nakhla were rather dry. Core-to-rim decreases in Li

abundances in pyroxenes of Zagami have been explained by possible involvement of late-

stage fluids (LENTZ et al., 2001) that could also potentially fractionate Li isotopes. Indeed,

low-temperature alteration hydrous phases, e.g. "iddingsite" (a mixture of smectites and

ferrihydrites), gypsum, calcite and halite have been found in nakhlites (Bunch and Reid,

1975). The high ö7Li of Nakhla may therefore be related to the former presence of water on

Mars. Significant changes in 8 Li between pyroxene cores and rims could be caused by the

presence of water during petrogenesis of martian basalts (BECK et al., 2004). However, a

recent detailed ion-microprobe study of distribution of Li and its isotopes in olivine and

pyroxene phenocrysts in lunar meteorite NWA 479 revealed large variations of Ô Li within

single crystals (Barrat et al., 2005). It has been concluded that diffusive isotope

fractionation is responsible for observed Li isotope patterns. By analogy, martian meteorite

NWA 480 studied previously (Beck et al., 2004) could also reflect a similar process rather

than degassing. However, further work is necessary to verify this.

Eucrites are thought to be excavated from asteroid 4 Vesta by an impact (Melosh,

1984). Eucrites have high Li abundances compared to martian and terrestrial basalts. Basaltic

eucrites possibly represent large-degree melts from chondritic source (BLICHERT-TOFT et al.,

2002). Given its incompatibility, the high Li content is therefore surprising (Table 5.1). An

unlikely explanation is that the terrestrial and martian mantles are relatively depleted in Li

because of crust formation. This would require that about half of the terrestrial lithium resides

in the crust. Although Li is enriched in crust there is little evidence that Li in the crust can

account for half the abundance of the terrestrial mantle. There is also no evidence that Li is

lost to the core. The most likely explanation might be that the asteroid 4 Vesta is internally

heterogeneous. The Li abundances of eucrites are in fact very similar to high-Ti mare basalts

of the Moon. The 6 Li value of studied eucrites (see Table 5.1) averages at +3.7%o, identical

within error to Mars (Zagami), the Moon and the Earth's upper mantle (sec above). This

provides strong evidence that whatever the reason for the high Li abundance of eucrites, it

does not relate to a process that produces large isotope fractionation.

110


5.5.5 Lithium isotope composition ofchondrites

Two chondrites have also been measured (see Figure 5.2 and Table 5.1). The

carbonaceous CV3 chondrite Allende gives 57Li = +3.1 %o whereas a lower 67Li = +1.5%o has

been obtained for the L6 chondrite Bruderheim. These results arc in accord with a more

comprehensive study of chondrites, which reports a range from -4 to +3%o (McDonough et

al., 2003). Recently, ZACK et al. (2003) demonstrated that metamorphic dehydration at

slightly higher temperatures will shift the Li isotope composition of rocks toward negative

values, as low as -ll%o. The variation among chondrites may reflect aqueous alteration on

parent bodies with isotopically light Li in the residual metamorphosed chondrites (e.g.

Bruderheim) and isotopically heavy Li related to chondrites with high aqueous alteration

index (Orgueil, Allende; JAMES and PALMER, 2000; McDonough et al., 2003). The

isotopically light Li of chondrites may reflect the bulk solar system Ô Li, assuming that if

anything aqueous processing has shifted some samples to heavier compositions. The bulk

silicate Earth is expected to have an average Ô Li value of +4%o, estimated from MORB (e.g.

Chan et al., 1992). A solar system with 57Li ~0%o (McDonough ct al., 2003) or Ô7Li ~2%o

(W. McDonough, personal communication to ANH, 2005) would imply either that the planet

formation process involves some small but significant Li isotope fractionation or there is a

currently unidentified isotopically light reservoir within the Earth.

5.6 Conclusions - inner solar system Li isotope compositions

From these preliminary studies some very clear basic conclusions can be drawn.

(1) Basaltic and ultramafic rocks from the Earth, Moon, Mars and Vesta yield broadly

similar Li isotope compositions consistent with a mantle source composition of Ô Li ~ +4%o.

There is no reason why each upper mantle should appear to be the same in its Li isotope

composition unless each entire terrestrial planetary object is also identical with 87Li ~ +4%o.

Therefore, a first order conclusion is that the Earth's lower mantle is unlikely to be different

in composition from its upper mantle given the similarity between this and other inner solar

system upper mantles.

(2) The terrestrial, martian and some low-Ti lunar mare basalts have very similar Li

abundances whereas eucrites have high Li abundances similar to high-Ti basalts. This

111


indicates a heterogeneous interior of Vesta. The reason for this is unclear at present but may

be related to di fferentiation during a magma ocean stage.

(3) The Earth, Moon, Mars and Vesta have different structure and composition of their

respective crusts and cores. They also have dramatically different geology that includes a

long history of subduction of isotopically fractionated Li into the Earth's mantle. However,

none of these differences appear to have resulted in major changes in the Li isotope

composition of the bulk planetary mantles.

(4) The Li isotope variation among lunar rocks reflects magmatic differentiation during

the lunar magma ocean stage. Lithium isotope fractionation is likely caused by crystallisation

of olivine, pyroxene and feldspar from the magma ocean and produces a relatively large

isotope effect.

(5) There is no evidence for lithium isotope fractionation related to the depletion of

volatiles in the inner solar system. This is similar to the result obtained for K (Humayun and

Clayton, 1995).

(6) There is also no indication that the Moon-forming Giant Impact resulted in Li

isotope fractionation. Therefore, the effects found for Fe (POITRASSON et al., 2004) may be

caused by fractionation during vaporisation of iron metal.

(7) Chondrites, as indicated by the results of this and an earlier study, are isotopically

light compared with differentiated planetary bodies of the inner solar system. This is consis¬

tent the with results of a more comprehensive study on chondrites (McDonough et al.,

2003). Therefore, there appears to be an important fractionation that takes place in the

accretionary disc or in chondrite parent bodies during their thermal metamorphism and

aqueous alteration.

Acknowledgement - We kindly thank Dmitri lonov for providing splits of mineral separates

from Vitim and Tariat peridotitic xenoliths, Mark Rehkämper for split of Kilbourne Hole

sample and Albrecht von Quadt for splits of two peridotitic standard rocks. We appreciated

discussions with Thorsten Kleine, Ingo Leya, Bill McDonough, Rainer Wieler and

HelenWilliams. Felix Oberli is acknowledged for maintenance of Nul700. The manuscript

benefited from careful and stimulating reviews by Tim Elliott and one anonymous reviewer.

Wc arc grateful to David Price for editorial handling. SNF and ETH funded this study.

112

Chapter 6 Summary and outlook

Chapter 6

Summary and outlook


6.1 Summary and outlook

In this project a new method for low-blank separation and isotope measurements of

lithium (Li) was set up which has subsequently been applied to terrestrial rocks, meteorites

and lunar samples. These included continental arc volcanic rocks (primitive olivine tholeiites,

magnesian basaltic andésites, evolved andésites and dacites), ocean island basalts from

Iceland and Jan Mayen, Mongolian spinel peridotite xenoliths and their minerals, and a set of

samples from inner solar system bodies (Moon, Mars, eucrites, chondrites). For most of these

Li isotope data have been complemented by trace element abundance measurements. The

main achievements are summarised in the following:

1 ) A new method for separation of Li from a wide variety of silicate rocks was developed

that employed a single-step cation-exchange chromatography and mixture of IM HNO3

blended with 80% methanol as elution media for concomitant separation of Li from other

major and minor elements. Use of low quantities of elution media (-16 ml) resulted in

insignificant and negligible contributions of blank Li. The robustness of the method has

been proven by identical elution schemes for variety of rocks with distinct silica contents

(basalts to rhyolites). The isotope homogeneity of Li international standard (L-SVEC) has

been verified based on aliquots from three international labs compared to the aliquot used

at ETH; the level of L-SVEC uniformity is better than 0.05%o (2a). No significant matrix

effect on Li isotope measurements due to the presence of large amounts of major elements

(e.g. Na, K, Mg, Ca, AI and Fe) has been observed. The external reproducibility of S7Li as

assessed by multiple replicate measurements of the reference rocks JB-2 and BHVO-2 is

+ 0.3%o (2a). Several other reference rocks provide long-term reproducibility better than

0.5%o (2a) and differences between distinct generations of reference rocks (e.g. BHVO-1

vs. BHVO-2 and AGV-1 vs. AGV-2) are greater than the analytical error. The latter thus

provides evidence of small-scale heterogeneity in the various generations of rock

standards. An improvement of the analytical method has been presented for Li separation

from individual mineral species of peridotitic rocks that employed small volumes of

cation exchange resin combined with 0.5M HNO3 as elution media. It provided additional

purification of Li from the large amounts of Cr present in clinopyroxenes. Future

innovation in Li analytical chemistry could be provided by extending the diversity of

organic media and/or searching for Li-sclectivc ion-exchange resins. Characterisation of

114


Li isotope compositions of a wider variety of reference rocks would certainly be helpful

for inter-laboratory comparison that is now covered by a restricted number of available

well-characterised standards. Recent developments in secondary ionisation mass

spectrometry (SIMS; CHAUSSIDON and ROBERT, 1998, 1999; DECITRE et al., 2002;

Kase.MANN et al., 2005) allow closer insights into variations of Li isotopes within

individual minerals and also chondrules (Beck et al., 2004, 2006; Barrat et al., 2005;

Chaussidon et al., 2006; Jeffcoate et al., 2007). Improvements of this alternative

analytical technique in terms of accuracy and precision of measuring Li isotopes could be

important for further research on a micron scale.

2) Subduction zones represent complex geological environments. Stable isotopes are

becoming an important tool to unravel and quantify the material transport in subduction

zones, i.e. which materials are recycled to the surface regions and which arc subducted

into the deep mantle. Seawater alteration preferentially adds 7Li to ocean-floor basalts.

Therefore, isotopically heavy Li is subducted into the mantle. Although Li is mobilised

during dehydration, there is no clear correlation with other parameters of fluid mobility.

Comparison of several arcs clearly shows distinct across-arc profiles providing further

evidence that Li in subduction zone volcanics is not a simple mixture of mantle and

recycled crustal Li (Moriguti and Nakamura, 1998b; Chan et al., 2002b; Leeman et

al., 2004). In the southern Cascadia subduction zone Li isotopes are decoupled from fluid

mobile elements (as indicated by e.g. Ba/Nb, Sr/Y) suggesting continuous isotope

fractionation during subduction and in the mantle wedge. The high-alumina olivine

tholeiites that originated from near-anhydrous melting of the spinel peridotite are

isotopically lighter in Medicine Lake volcanics compared to Mt. Shasta lavas. This

provides evidence for a heterogeneous mantle although both HAOT suits are enriched in

fluid mobile elements (e.g. Sr, Pb and Li) relative to MORBs. The timing of this

metasomatic enrichment is unclear but the anhydrous character of these lavas requires

non-uniform mantle with predominantly light Li beneath Medicine Lake. Basaltic

andésites and magnesian andésites show a large range in 57Li with lighter values for

Medicine Lake volcanics. Two fluid compositions with primarily heavy Li isotope

signature are reported for Mt. Shasta BA lavas that have distinct Sr isotope ratios

(MORB-like vs. "sediment-labelled"). Different fluids with light 57Li<0.9%o have been

deduced from analysis of Medicine Lake BA lavas. Indistinguishable Sr isotope ratios for

115


Medicine Lake BA lavas and "sediment-labelled" Mt. Shasta BA lavas provide evidence

for a similarity in the sources of hydrous Medicine Lake and Mt. Shasta lavas. The

depletion of lavas in 7Li follows the concept of progressive loss of heavy Li into the

overlying mantle wedge (Zack ct al., 2003). Light signatures away from the arc trench

are interpreted as a slab signature that is the direct consequence of continuous Li isotope

fractionation during dehydration and subduction. Nevertheless, a better understanding of

processes beneath Mt. Shasta region is rendered difficult by a lack of precise information

on Li partitioning and, in particular, isotope fractionation among major minerals and fluid

phases for the whole range of p-T conditions and should be improved by acquiring

relevant data (e.g. Lynton et ai., 2005; Marschall et al., 2006). Diffusion-driven

fractionation of Li isotopes may be an important mechanism that could provide

explanation for isotopically heavy fluids or melts and, in parallel, isotopically lighter

minerals generally (LUNDSTROM ct al., 2005; JEFFCOATE et al., 2007) or, more

specifically, in arc settings (TOMASCAK et al., 2002). The extent of this fractionation

between clinopyroxene and coexisting fluid is reduced with increasing temperature but

can still be resolved at T>900°C (Wunder et al., 2006).

3) Some noble gas data of Earth's interior require "hidden" reservoirs deep within Earth

(STUART et al., 2003). Such regions should recharge the upwelling mantle plumes with a

characteristic He-enriched signature, which is often considered as remnants from the

formation times of the Earth. Alternatively, ancient recycled oceanic lithosphère has been

called upon as a source of high 3Hc/4He ratios (ANDERSON, 1998). Coupled Li-He isotope

data from picritic basalts from Hengill area, Iceland, suggest an alternative explanation

for high 3He abundances that runs contrary to theories of "hidden" noble gas reservoirs at

the base of the Earth's lower mantle and/or in D" layer. It is suggested that the observed

inverse and linear Li-He relationship is most likely generated in the mantle. Isotopically

heavy Li carried by the low-3He/4He lavas could reflect a portion of ancient recycled

oceanic lithosphère in the source of the Icelandic plume. This is corroborated by positive

Eu anomaly and super-chondritic Sr/Nd which requires involvement of plagioclase in the

earlier history although none of the lavas contain feldspar as a phenocryst phase. High-

HerHe lavas have low ô Li but are diverse in terms of REE systematics. The (La/Sm)N

of Enriched lavas are similar to that of the upper mantle whereas Depleted lavas show

homogeneous (La/Sm)N consistent with the estimate of the depleted mantle. The Jan

116


Mayen lavas show particular enrichment in incompatible trace elements combined with

limited variations in Ô Li. This enrichment is suggestive of incorporation of a mantle

domain with EM2 affinity. Hot-spot origin of Jan Mayen cannot be ruled out at present

based on Li isotope data. Overall, the data for Iceland and Jan Mayen suggest a globally

uniform 87Li ~ +4%o independent of trace element enrichment or depletion. A recent

study of distribution of Li isotopes in a suite of East Pacific Rise lavas (Elliott et al.,

2006) provides an alternative explanation for the origin of the recycled component in

MORB which is supposedly a result of entrainment of fluid-modified subarc mantle into

the source region of MOR basalts. Hence, these diverse opinions reflect the need for

further studies.

4) A study of Li isotopes in non-terrestrial materials indicates a homogeneous composition

of differentiated inner solar system bodies whereas chondrites may be isotopically slightly

lighter. Thus, the planet formation processes may have a small effect on Li isotopes. The

large range in 57Li among lunar rocks reflects extreme degrees of crystallization (>95%)

of the lunar magma ocean. The early- and late-stage cumulates with variable amounts of

major minerals with isotopically different Li will produce basalts with distinct 6 Li. The

presence or absence of clinopyroxene may be a major factor controlling the Li isotope

compositions of the mare basalts. The bulk Moon ö7Li value at +3.8%o has been inferred

from lunar rocks showing the least Eu depletion. The lunar crust represented by ferroan

anorthosite seems to yield an extremely heavy Li isotope signature that may reflect

equilibrium fractionation during plagioclase segregation from the early lunar magma

ocean. The lunar regolith may have recorded the implantation of the isotopically heavy

solar wind as evidenced by progressive enrichment in 7Li within finer fractions of a lunar

soil. Pristine peridotitic rocks provide evidence for a 8 Li ~ -t4.0%o for the Earth's upper

mantle. This is further supported by a limited range of ô7Li measured for a worldwide

collection of olivines from spinel lherzolites. There is no resolvable difference in

estimated Ô Li between terrestrial spinel peridotites and basaltic rocks and/or meteorites

from Moon, Mars and Vesta despite large differences in the volatile content of these

bodies. This provides evidence for uniform S Li of mantles of the inner solar system

planets. Available data for chondrites tend to reveal small but resolvable Li isotope

fractionation within the accretionary disc or chondrite parent bodies. Extending the Li

isotope database to chondritic meteorites of various petrologic classes may further

117


contribute to our knowledge of the processes in the young inner solar system

(MCDONOUGH et al., 2003, 2006; SEITZ et al., 2006b). The Moon currently represents an

open field for exploration by Li isotopes because there was so far no attempt for detailed

study of Li isotopes in the lunar materials. Relation of the lunar crust and the lunar mantle

may be deconvoluted using Li isotopes for a larger suite of lunar crustal rocks. Evolution

and cessation of the Lunar Magma Ocean could be explored further with Li isotope data

for other lunar rocks and in particular mineral phases. Characterisation of Li isotope

compositions for variety of other rock types (Mg-suite, low-Ti and high-Ti mare basalts,

KREEP basalts) could further evaluate the processes that lead to genesis of these

important lunar magmatic phases.

118

References

References

Anders E. and Grevesse N. (1989) Abundances of the elements: Meteoritic and solar. Geochim.

Cosmochim. Acta 53, 197-214

Anderson A.T. (1974) Evidence for a picritic, volatile-rich magma beneath Mt. Shasta, California. J.

Petrol. 15,243-267

Anderson D.L. (1998) The helium paradoxes. Proc. Natl. Acad. Sei. USA 95, 4822-4827

Assonov S.S., Franchi I.A., Pillinger CT., Semenova A.S., Shukolyukov Y.A., Verchovsky A.B. and

lassevitch A.N. (2002) Nitrogen and argon release profiles in Luna 16 and Luna 24 regolith

samples: The effects of regolith reworking. Meteoritics Planet. Sei. 37, 27-48

Aston F.W. (1927) Bakcrian lecture. A new mass-spectrograph and the whole number rule. Proc. Roy.Soc. London A 115,487-514

Aston F.W. (1932) The isotopic constitution and atomic weights of caesium, strontium, lithium,

rubidium, barium, scandium and thallium. Proc. Roy. Soc. London A 134, 571-578

Bacon CR., Bniggman P.E., Christiansen R.L., Clynne M.A., Donnelly-Nolan J.M. and Hildreth W.

(1997) Primitive magmas at five Cascade volcanic fields: melts from hot, heterogeneous sub-arc

mantle. Canad. Mineral. 35, 397-423

Baker J.A. and Jensen K.K. (2004) Coupled 18COs-187Os enrichments in the Harm's mantle - core-

mantle interaction or recycling of ferromanganese crusts and nodules? Earth Planet. Sei. Lett.

220, 277-286

Baker M.B., Grove T.L., Kinzler R.J., Donnelly-Nolan J.M. and Wandless G.A. (1991) Origin of

compositional zonation (high-alumina basalt to basaltic andésite) in the Giant Crater lava field,Medicine Lake Volcano, Northern California. J. Geophys. Res. 96 (B13), 21819-21842

Baker M.B., Grove T.L. and Price R.C. (1994) Primitive basalts and andésites form the Mt.Shasta

region, N.California: products of varying melt fraction and water content. Contrib. Mineral.

Petrol. 118,111-129

Barrat J.A., Chaussidon M., Bohn M., Gillet P., Göpel C. and Lesourd M. (2005) Lithium behavior

during cooling of a dry basalt: An ion-microprobe study of the lunar meteorite Northwest Africa

479 (NWA 479). Geochim. Cosmochim. Acta 69 (23), 5597-5609

Bartels K.S., Kinzler R.J. and Grove T.L. (1991) High pressure phase relations of primitive high-alumina basalts from Medicine Lake, northern California. Contrib. Mineral. Petrol. 108, 253-270

Beck P., Barrât J.A., Chaussidon M., Gillet P. and Bohn M. (2004) Li isotopic variations in single

pyroxenes from the Northwest Africa 480 shergottite (NWA 480): a record of degassing of

Martian magmas'? Geochim. Cosmochim. Acta 68 (13), 2925-2933

Beck P., Chaussidon M., Barrat J.A., Gillet P., and Bohn M. (2006) Diffusion induced Li isotopicfractionation during the cooling of magmatic rocks: The case of pyroxene phenocrysts form

nakhlite meteorites. Geochim. Cosmochim. Acta 70, 4813-4825

Benton L.D., Ryan J.G. and Savov LP. (2004) Lithium abundance and isotope systematics of forearc

serpentinites, Conical Seamount, Mariana forearc: Insights into the mechanics of slab-mantle

exchange during subduction. G-cubed 5 (8), paper number Q08J12, doi: 10.1029/2004GC000708

119

References

Benz H.M., Zandt G. and Oppenheimer D.H. (1992) Lithospheric structure of northern California

from teleseismic images of the upper mantle. J. Geophys. Res. 97, 4791-4807

Bickle M.J., Chapman H.J. and You C.-F. (2000) Measurement of lithium isotopic ratios as lithium

tetraborate ions. Int. J. Mass Spectrom. 202, 273-282

Bigeleisen J. (1965) Chemistry of isotopes. Science 147, 463-471

Bindeman I.N., Davis A.M. and Drake M.J. (1998) Ion-microprobe study of plagioclase-basalt

partition experiments at natural concentration levels of trace elements. Geochim. Cosmochim.

Acta 62 (7), 1175-1193

Blichert-Toft J„ Agranier A., Andres M., Kingsley R., Schilling J.-G., and Albarede F. (2005)Geochemical segmentation of the Mid-Atlantic Ridge north of Iceland and ridge-hot spotinteraction in the North Atlantic. G-cuhed 6 (1), paper number Q01E19,doi : 10.1029/2004GC000788.

Blichert-Toft J., Boyet M., Télouk P. and Albarede F. (2002) M7Sm-143Nd and 176Lu-176Iirin eucrites

and the differentiation of the HED parent body. Earth Planet. Sei. Lett. 204, 167-181

Blundy J. and Cashman K. (2001) Ascent-driven crystallization of dacite magmas at Mount St.

Helens, 1980-1986. Contrib. Mineral. Petrol. 140, 631-650

Borg L.E., Clynne M.A. and Bullen T.D. (1997) The variable role of slab-derived fluids in the

generation of a suite of primitive calc-alkaline lavas from the southernmost Cascades, California.

Canad. Mineral, 35, 425-452

Borg L.E., Brandon A.D., Clynne M.A. and Walker R.J. (2000) Re-Os isotopic systematics of

primitive lavas from the Lassen region of the Cascade arc, California. Earth Planet. Sei. Lett. 177,

301-317

Bouman C, Vroon P.Z., Elliott T.R., Schwieters J. and Hamester M. (2002) High precisionmeasurements of lithium isotope ratios by multi-collector ICMPS (ThermoFinniganMATNeptune). Plasma Spectrom. Winter Conf, Scottsdale, AZ

Bouman C, Elliott T. and Vroon P.Z. (2004) Lithium inputs to subduction zones. Chem. Geol. 212 (1 -

2), 59-79

Breddam K. (2002) Kistufell: primitive melts from the Iceland mantle plume. J. Petrol. 43 (2), 345-

373

Brooker R., Blundy J. and James R.H. (2000) Subduction-related mantle pyroxenites from Zabargadisland, Red Sea. J. Conf. Abstracts 5, 249

Brooker R.A., James R.H. and Blundy J.D. (2004) Trace elements and Li isotope systematics in

Zabargad peridotites: evidence of ancient subduction processes in the Red Sea mantle. Chem.

Geol. 212, 179-204

Bryant C.J., Chappell B.W., Bennett V.C. and McCulloch M.T. (2003a) Li isotopic variations in

Eastern Australian granites. Geochim. Cosmochim. Acta 67 (18), A47

Bryant C.J., McCulloch M.T. and Bennett V.C. (2003b) Impact of matrix effects on the accurate

measurement of Li isotope ratios by inductively coupled plasma mass spectrometry (MC-ICP¬

MS) under "cold" plasma conditions. J. Anal. Atom. Spectrom. 18, 734-737

Bryant C.J., Chappell B.W., Bennett V.C, and McCulloch M.T. (2004) Lithium isotopic compositionsof the New England Batholith: correlations with inferred source rock compositions. Trans. R. Soc.

Edinburgh - Earth Sei. 95, 199-214

120

References

Bunch T.E. and Reid A.M. (1975) The nakhlites, part 1: Petrography and mineral chemistry.Meteoritics 10,303-315

Burnard P., Graham D., and Farley K. (2004) Fractionation of noble gases (He, Ar) during MORB

mantle melting: a case study on the Southeast Indian Ridge. Earth Planet. Sei. Lett. 227 (3-4),457-472

Burnard P. and Harrison D. (2005) Argon isotope constraints on modification of oxygen isotopes in

Iceland basalts by surficial processes. Chem. Geol. 216 (1-2), 143-156

Canup R.M. and Agnor CB. (2000) Accretion of the terrestrial planets and the Earth-Moon system. In

The origin of the Earth and Moon (ed. R.M. Canup and K. Righter), pp. 1 13-129. Univ. Arizona

Press

Chambers J.E. (2004) Planetary accretion in the inner Solar System. Earth Planet. Sei. Lett. 223, 241-

252

Chan L.H. (1987) Lithium isotope analysis by thermal ionization mass spectrometry of lithium

tetraborate. Anal. Chem. 59, 2662-2665

Chan L.H., Edmond J.M., Thompson G. and Gillis K. (1992) Lithium isotopic composition of

submarine basalts: implication for the lithium cycle in the oceans. Earth Planet. Sei. Lett. 108 (l-

3), 151-160

Chan L.H., Edmond J.M. and Thompson G. (1993) A lithium isotope study of hot springs and

metabasalts from mid-ocean ridge hydrothermal systems. J. Geophys. Res. 98, 9653-9659

Chan L.H., Leeman W.P. and You C.-F. (1999) Lithium isotopic composition of Central American

Volcanic Arc lavas: implications for modification of subarc mantle by slab-derived fluids. Chem.

Geol. 160, 255-280

Chan L.H. and Kastner M. (2000) Lithium isotopic composition of pore fluids and sediments in the

Costa Rica subduction zone: implications for fluid processes and sediment contribution to the arc

volcanoes. Earth Planet. Sei. Lett. 183, 275-290

Chan L.H., Alt J.C. and Teagle D.A.H. (2002a) Lithium and lithium isotope profiles through the upper

oceanic crust: a study of seawater-basalt exchange at ODP sites 504B and 896A. Earth Planet.

Sei. Lett. 201 (1), 187-201

Chan L.H., Leeman W.P. and You C.-F. (2002b) Lithium isotopic composition of Central American

Volcanic Arc lavas: implications for modification of subarc mantle by slab-derived fluids:

correction. Chem. Geol. 182 (3-4), 293-300

Chan L.H. (2003) Variation of lithium isotopic composition in terrestrial systems. Geochim.

Cosmochim. Acta 67, A57

Chan L.H. and Frey A.F. (2003) Lithium isotope geochemistry of the Hawaiian plume: Results from

the Hawaii Scientific Drilling Project and Koolau volcano. G-cubed 4 (3), paper number 8707,

doi : 10.1029/2002GC000365

Chan L.H., Leeman W.P., and Plank T. (2006) Lithium isotopic composition of marine sediments. G-

cubedl (6), paper number Q06005, doi:10.1029/2005GC001202

Chapman J.F. and Dale L.S. (1976) The determination of lithium isotope abundances with a dual-

beam atomic absorption spectrometer. Anal. Chim. Acta 87, 91-95

Chaussidon M. and Robert F. (1996) Boron and lithium isotope variations in chondmles: the signatureof presolar nucleosynthesis. Lunar Planet. Sei. Conf. XXV11, 207-208

121

References

Chaussidon M. and Robert F. (1998) 7Li/6Li and nB/H,B variations in chondrules from Semarkona

unequilibrated chondrite. Earth Planet. Set. Lett. 164, 577-589

Chaussidon M. and Robert F. (1999) Lithium nucleosynthesis in the Sun inferred from the solar-wind

7Li/6Li ratio. Nature 402, 270-273

Chaussidon M., Robert F., McKeegan K..D. and Krot A.N. (2001) Lithium and boron isotopic

compositions of refractory inclusions from primitive chondrites: a record of irradiation in the

early Solar system. Lunar Planet. Sei. Conf. XXXII, #1862

Chaussidon M., Robert F. and McKeegan K.D. (2006) Li and B isotopic variations in an Allende CAL

Evidence for the in situ decay of short-lived 10Bc and for the possible presence of the short-lived

nuclide 7Be in the early solar system. Geochim. Cosmochim. Acta 70 (1), 224-245

Chauvel C. and llémond C (2000) Melting of a complete section of recycled oceanic crust: Trace

clement and Pb isotopic evidence from Iceland. G-cubed 1, paper number

doi: 10.1029/1999GC000002

Christiansen R.L., Kleinhampl F.J., Blakely R.J., Tuchek E.T., Johnson F.L. and Conyac M.D. (1977)Resource appraisal of the Mt. Shasta wilderness study area. Open-file report 77-250, US Geol,

Surv.

Clynne M.A. and Borg L.E. (1997) Olivine and chromian spinel in primitive calc-alkaline and

tholeiitic lavas from the southernmost Cascade range, California: A reflection of relative fertilityof the source. Canad. Mineral. 35, 453-472

Cohen A.S., O'Nions R.K., Siegenthaler R. and Griffin W.L. (1988) Chronology of the pressure-

temperature history recorded by a granulite terrain. Contrib. Mineral. Petrol. 98, 303-311

Conrey R.M., Hooper P.R., Larson P.B., Chelsey J. and Ruiz J. (2001) Trace clement and isotopicevidence for two types of crustal melting beneath a High Cascade volcanic center, Mt. Jefferson,

Oregon. Contrib. Mineral. Petrol. 141, 710-732

Coogan L.A., Kasemann S.A. and Chakraborty S. (2005) Rates of hydrothermal cooling of new

oceanic upper crust derived from lithium-geospeedometry. Earth Planet. Sei. Lett. 240, 415-424

Coplen T.B., Böhlke J.K., De Bièvre P., Ding T., Holden N.E., Hopple J.A., Krause H.R., LambcrtyA., Peiser H.S., Révész K, Rieder S.E., Rosman K.J.R., Roth E., Taylor P.D.P., Vocke Jr. R.D.

and Xiao Y.K. (2002) Isotope-abundance variations of selected elements. Pure Appl. Chem. 74

(10), 1987-2017

David K., Schiano P. and Allègre C.J. (2000) Assessment of the Zr/Hf fractionation in oceanic basalts

and continental materials during petrogenetic processes. Earth Planet. Sei. Lett. 178, 285-301

Decitre S., Deloule E., Reisberg L., James R.H., Agrinier P. and Mével C. (2002) Behavior of Li and

its isotopes during serpentinization of oceanic peridotites. G-cubed 3 (1),doi: 10.1029/2001GC000178

Delbourgo-Salvador P., Malinie G. and Adouze J. (1985) Thermal nuclear destruction of lithium,

beryllium and boron inside a 1 solar mass star. In Isotopic ratios in the Solar System (ed. Centre

national d'études spatiales, Tolouse: Cepadues-ßditions), pp. 267-272

Dempster A.J. (1921) Positive ray analysis oflithium and magnesium. Phys. Rev. 18 (6), 415-422

Donnelly-Nolan J.M., Champion D.E., Grove T.L., Baker M.B., Taggart Jr. J.E. and Bruggman P.E.

(1991) The Giant Crater lava field: Geology and geochemistry of a compositionally zoned, high-alumina basalt to basaltic andésite eruption at Medicine Lake Volcano, California. J. Geophys.Res. 96 (B13), 21843-21863

122

References

Dorendorf F., Wiechert U. and Wörner G. (2000) Hydrated sub-arc mantle: a source for die

Klyuchevskoy volcano, Kamchatka/Russia. Earth Planet. Sei. Lett. 175, 69-86

Dreibus G., Huisl W., Spettel B. and Haubold R. (2003) The chemical composition of the lunar

meteorite EET 96008 and the lithium distribution on Moon. Annu. Meteorit. Soc. Meet. 66th,

#5108

Drozd R.J., Hohenberg CM., Morgan C.J., Podosek F.A. and Wroge M.L. (1977) Cosmic-ray

exposure history of Taurus-Littrow. Proc. Lunar Sei. Conf. 8th, 3027-3043

Dutrow B.L., Holdaway M.J. and Hinton R.W. (1986) Lithium in staurolitc and its petrologic

significance. Contrib. Mineral. Petrol. 94, 496-506

Ebihara M., Wolf R., Warren P.H. and Anders E. (1992) Trace elements in 59 mostly highland Moon

rocks. Proc. Lunar Planet. Sei. Conf. 22nd, 417-426

Edwards C.M.H., Morris J.D. and Thirlwall M.F. ( 1993) Separating mantle from slab signatures in arc

lavas using B/Bc and radiogenic isotope systematics. Nature 362, 530-533

Eiler J.M. (2001) Oxygen isotope variations of basaltic lavas and upper mantle rocks. Rev Mineral,

Geochem. 43,319-364

Elkins Tanton L.T., Grove T.L. and Donnelly-Nolan J.M. (2001) Hot, shallow mantle melting under

the Cascades volcanic arc. Geology 29 (7), 631-634

Ellam R.M. and Stuart F.M. (2004) Coherent He-Nd-Sr isotope trends in high 3He/4He basalts:

implications for a common reservoir, mantle heterogeneity and convection. Earth Planet. Sei.

Lett. 228(3-4), 511-523

Elliott T.R., Hawkcsworth C.J., and Grönvold K. (1991) Dynamic melting of the Iceland plume.Nature 351, 201-206

Elliott T., Zindler A., and Bourdon B. (1999) Exploring the kappa conundrum: the role of recycling in

the lead isotope evolution of the mantle. Earth Planet. Sei. Lett. 169 (1-2), 129-145

Elliott T., Jeffcoate A. and Bouman C. (2004) The terrestrial Li isotope cycle: light-weight constraints

on mantle convection. Earth Planet. Sei. Lett. 220 (3-4), 231-245

Elliott T., Thomas A., Jeffcoate A., and Niu Y. (2006) Lithium isotope evidence for subduction-

enriched mantle in the source of mid-ocean-ridge basalts. Nature 443, 565-568

Eugster O. and Bernas R. (1971) Li, B, Mg and Ti isotopic abundances and search for trapped solar

wind Li in Apollo 11 and Apollo 12 material. Proc. Lunar Sei. Conf. 2nd, 1641-1469

Eugster O., Eberhardt P., Geiss J., Grögler N., Jungck M. and Mörgeli M. (1977) The cosmic-ray

exposure history of Shorty Crater samples; The age of Shorty Crater. Proc. Lunar Sei. Conf. 8th,

3059-3082

Fahre C, Boiron M.C, Dubessy J., Chabiron A., Charoy B., and Crespo T.M. (2002) Advances in

lithium analysis in solids by means of laser-induced breakdown spectroscopy: An exploratory

study. Geochim. Cosmochim. Acta 66 (8), 1401-1407

Fitton J.G., Saunders A.D., Norry M.J., Hardarson B.S., and Taylor R.N. (1997) Thermal and

chemical structure of the Icelandic plume. Earth Planet. Sei. Lett. 153 (3-4), 197-208

Fitton J.G., Saunders A.D., Kempton P.D., and Hardarson B.S. (2003) Does depleted mantle form an

intrinsic part of the Iceland plume? G-cubed 4 (3), paper number 1032,

doi : 10.1029/2002GC000424

123

Referenc es

Flesch G.D., Anderson A.R. and Svec H.J. (1973) A secondary isotopic standard for 6Li/7Li

determinations. Int. J. Mass Spectrom. Ion Phys. 12, 265-272

Freedman P.A. (2002) Mass bias in ICP mass spectrometers. Geochim. Cosmochim. Acta 66, A245

Fuis G.S., Zucca J.J., Mooney W.D. and Milkereit B. (1987) A geologic interpretation of seismic-

reflection results in northern California. Geol, Soc. Am. Bull. 98, 53-65

Galy A., Yoffe O., Janney P.E., Williams R.W., Cloquet C, Alard O., Halicz L., Wadhwa M.,

Hutcheon I.D., Ramon K. and Carignan J. (2003) Magnesium isotope heterogeneity of the

isotopic standard SRM980 and new reference materials for magnesium-isotope-ratiomeasurements./ Anal. Atom. Spectrom. 18, 1352-1356

Gautason B, and Muehlenbachs K. (1998) Oxygen isotopic fluxes associated with high-temperature

processes in the rift zones of Iceland. Chem. Geol. 145 (3-4), 275-286

Gill J.B. (1981) Orogenic andésites and plate tectonics. Springer-Verlag, Berlin. 391 pp.

Graham D.W. (2003) Noble gas isotope geochemistry of mid-ocean ridge and ocean island basalts:

Characterization of mantle source reservoirs. Rev. Mineral. Geochem. 47, 247-318

Green N.L. and Harry D.L. (1999) On the relationship between the subducted slab age and arc basalt

petrogenesis, Cascadia subduction system, North America. Earth Planet. Sei. Lett. 171, 367-381

Grove T.L. and Kinzler R.J. (1986) Petrogenesis of andésites. Annu. Rev. Earth Planet. Sei. 14, 417-

454

Grove T.L., Parman S.W., Bowring S.A., Price R.C. and Baker M.B. (2002) The role of an H20-rich

fluid component in the generation of primitive basaltic andésites and andésites from the Mt.

Shasta region, N California. Contrib. Mineral. Petrol. 142, 375-396

Grove T.L., Elkins Tanton L.T., Parman S.W., Chatterjec N., Müntener O. and Gaetani G.A. (2003)Fractional crystallization and mantle-melting controls on calc-alkaline differentiation trends.

Contrib. Mineral. Petrol. 145, 515-533

Grove T.L., Baker M.B., Price R.C, Parman S.W., Elkins Tanton L.T., Chatterjee N. and Müntener O.

(2005) Magnesian andésite and dacite lavas from Mt. Shasta, northern California: products of

fractional crystallization of H20-rich mantle melts. Contrib. Mineral. Petrol. 148, 542-565

Gurenko A.A. and Chaussidon M. (1997) Boron concentrations and isotopic composition of the

Icelandic mantle: evidence from glass inclusions in olivine. Chem. Geol. 135(1 -2), 21 -34

Gurenko A.A. and Chaussidon M. (2002) Oxygen isotope variations in primitive tholeiites of Iceland:

evidence from a SIMS study of glass inclusions, olivine phenocrysts and pillow rim glasses.Earth Planet. Sei. Lett. 205 (1-2), 63-79

Gurenko A.A. and Schmincke H.-U. (2002) Orthopyroxenc-bearing tholeiites of the Iblean Plateau

(Sicily): constraints on magma origin and evolution from glass inclusions in olivine and

orthopyroxene. Chem. Geol. 183, 305-331

Haase K., Devey C.W., Mertz D.F., Stoffers P., and Garbe-Schönberg D. (1996) Geochemistry of

lavas from Mohns Ridge, Norwegian-Greenland Sea: implications for melting conditions and

magma sources near Jan Mayen. Contrib. Mineral. Petrol. 123 (3), 223-237

Haase K.M. and Devey C.W. (1994) The petrology and geochemistry of Vesteris seamount,

Greenland basin - an intrapalte alkaline volcano of non-plume origin. J. Petrol. 35 (2), 295-328

124

References

Halliday A.N., Lee D.-C, Tommasini S., Davies GR., Paslick CR., Fitton J.G., and James D.E.

(1995) Incompatible trace elements in OIB and MORB and source enrichment in the sub-oceanic

mantle. Earth Planet, Sei, Lett. 133 (3-4), 379-395

Halliday A.N. and Porcelli D. (2001) In search of lost planets - the paleocosmochemistry of the inner

solar system. Earth Planet. Sei. Lett. 192 (4), 544-559

Halliday A.N., Freedman P.A., Oberli F., Baur H., Hollins S., Levasseur S., Leya L, Poitrasson F.,

Quitté G, Teutsch N., Wiechert U., Williams H. and Williams J. (2002) Nu 1700: a new large-

geometry high-resolution multiple collector ICP-MS. Geochim. Cosmochim. Acta 66, A303

Halliday A.N. (2003) The origin and the earliest history of the Earth. In Treatise on geochemistry,Vol. 1 (ed. A.M. Davis), pp. 509-557. Elsevier Ltd.

Hanan B.B. and Graham D.W. (1996) Lead and helium isotope evidence from oceanic basalts for a

common deep source of mantle plumes. Science 272, 991-995

Hansen H.-E., Grönvold K, Kosler J., and Pedersen R.-B. (2005) Isotopic composition of Li in

volcanic rocks from the northern Icelandic rift zone. AGU Fall Meet,, V41C-1464

Hansteen T.H. (1991) Multi-stage evolution of the picritic Maelifell rocks, SW Iceland: constraints

from mineralogy and inclusions of glass and fluid in olivine. Contrib. Mineral. Petrol. 109 (2),225-239

Barris R.A., Iyer H.M. and Dawson P.B. (1991) Imaging the Juan de Fuca plate beneath southern

Oregon using teleseistnic P wave residuals. J. Geophys, Res. 96, 19879-19889

Harrison D., Burnard P., and Turner G. (1999) Noble gas behaviour and composition in the mantle:

constraints from the Icelandic plume. Earth Planet. Sei. Lett. 171 (2), 199-207

Harrison D., Burnard P.G., Trieloff M., and Turner G. (2003) Resolving atmospheric contaminants in

mantle noble gas analyses. G-cubed4 (3), 1023, doi: 10.1029/2002GC000325

Harrison D., Barry T., and Turner G. (2004) Possible diffusive fractionation of helium isotopes in

olivine and clinopyroxene phenocrysts. Eur. J. Mineral. 16 (2), 213-220

Hart W.K. (1985) Chemical and isotopic evidence for mixing between depleted and enriched mantle,

northwestern U.S.A. Geochim. Cosmochim. Acta 49, 131-144

Harvey R.P. and McSween H.Y. (1992) The parent magma of the nakhlite meteorites: Clues from

melt inclusions. Earth Planet. Sei. Lett. Ill, 467-482

Heber V., Brooker R.A., Kelley S.P., and Wood B.J. (2007) Crystal-melt partitioning of noble gases

(helium, neon, argon, krypton, and xenon) for olivine and clinopyroxene. Geochim. Cosmochim.

Acta 71 (4), 1041-1061

Heiken Ci. and McKay D.S. (1974) Petrography of Apollo 17 soils. Proc. Lunar Sei. Conf 5th, 843-

860

Hemond C, Arndt N.T., Lichtenstein U., Hofmann A.W., Oskarsson N., and Steinthorsson S. (1993)The heterogeneous Icelandic plume - Nd-Sr-0 isotopes and trace-element constraints. J. Geophys.Res. 98 (B9), 15833-15850

Herd C.D.K., Treiman A.H., McKay G.A. and Shearer CK. (2004) The behavior of Li and B during

planetary basalt crystallization. Am. Mineral. 89, 832-840

Herd C.D.K., Treiman A.H., McKay G.A. and Shearer CK. (2005) Light lithophile elements in

martian basalts: Evaluating the evidence for magmatic water degassing. Geochim, Cosmochim.

Acta 69 (6), 2431-2440

125

References

Hilton D.R., Grönvold K., Macpherson CG., and Castillo P.R. (1999) Extreme 3He/4He ratios in

northwest Iceland: constraining the common component in mantle plumes. Earth Planet. Sei.

Lett. 173(1-2), 53-60

Hochstaedter A., Gill J., Peters R., Broughton P., Holden P. and Taylor B. (2001) Across-arc

geochemical trends in the Izu-Bonin arc: Contributions from the subducting slab. G-cubed 2,

paper number 2000GC000105

Iloefs J. and Sywall M. (1997) Lithium isotope composition of Quarternary and Tertiary biogenecarbonates and a global lithium isotope balance. Geochim. Cosmochim. Acta 61 (13), 2679-2690

Hofmann A.W. and White W.M. (1982) Mantle plumes from ancient oceanic crust. Earth Planet. Sei.

Le». 57 (2), 421-436

Hofmann A.W. and Jochum K.P. (1996) Source characteristics derived from very incompatible trace

elements in Mauna Loa and Mauna Kea basalts, Hawaii Scientific Drilling Project. J. Geophys.Res. 101 (B5), 11831-11839

Hofmann A.W. (1997) Mantle geochemistry: the message from oceanic volcanism. Nature 385, 218-

229

Hörz F., Gibbons R.V., Gault D.E., Härtung J.B. and Brownlee D.E. (1975) Some correlation of rock

exposure ages and regolith dynamics. Proc. Lunar Sei. Conf. 6th, 3495-3508

Ilorwitz E.P., Chiarizia R. and Dietz M.L. (1992) A novel strontium-selective extraction

chromatographic resin. Solv. Extract. Ion Exchange 10 (2), 313-336

Huh Y., Chan L.H., Zhang L., and Edmond J.M. (1998) Lithium and its isotopes in major world

rivers: implications for weathering and the oceanic budget. Geochim. Cosmochim. Acta 62 (12),2039-2051

Huh Y., Chan L.H. and Edmond J.M. (2001) Lithium isotopes as a probe of weathering processes:

Orinoco River. Earth Planet. Sei. Lett. 194, 189-199

Humayun M. and Clayton R.N. (1995) Potassium isotope cosmochcmistry: Genetic implications of

volatile clement depletion. Geochim. Cosmochim. Acta 59 ( 10), 2131 -2148

Huneke J.C, Jessberger E.K., Podosek P.A. and Wasserburg G.J. (1973) 40Ar/3l)Ar measurements in

Apollo 16 and 17 samples and chronology of metamorphic and volcanic activity in the Taurus-

Littrow region. Proc. Lunar Sei. Conf. 4th, 1725-1756

lonov D.A. (2004) Chemical variations in peridotite xenoliths from Vitim, Siberia: Inferences for REE

and Hf behaviour in the garnct-facies upper mantle. J. Petrol. 45 (2), 343-367

lonov D.A., Ashchepkov I. and Jagoutz E. (2005) The provenance of fertile off-craton lithosphericmantle: Sr-Nd isotope and chemical composition of garnet and spinel peridotite xenoliths from

Vitim, Siberia. Chem. Geol. 217, 41-75

Ishikawa T. and Tera F. (1997) Source, composition and distribution of the fluid in the Kurile mantle

wedge: Constraints from across-arc variations of B/Nb and B isotopes. Earth Planet. Sei. Lett.

152,123-138

Jagoutz E., Palme H., Baddenhausen H., Blum K., Cendales M., Dreibus G, Spettel B., Lorenz V. and

Wanke H. (1979) The abundances of major, minor and trace elements in the earth's mantle as

derived from primitive ultramafic nodules. Proc. Lunar Planet. Sei. Conf. 10th, 2031-2050

James R.H. and Palmer M.R. (2000) The lithium isotope composition of international rock standards.

Chem. Geol. 166,319-326

126

References

James R.H., Allen D.E. and Seyfried Jr. W.E. (2003) An experimental study of alteration of oceanic

crust and terrigenous sediments at moderate temperatures (51 to 350°C): Insights as to chemical

processes in near-shore ridge-flank hydrothermal systems. Geochim. Cosmochim. Acta 67 (4),681-691

Jeffcoate A.B., Elliott T., Thomas A. and Bouman C. (2004) Precise, small sample size determinations

of lithium isotopic compositions of geological reference materials and modern seawater by MC-

ICP-MS. Geostandards Geoanal. Res. 28 (1), 161-172

Jeffcoate A.B., Elliott T., Kasemann S.A., lonov D. and Cooper K.M. (2007) Li isotope fractionation

in peridotites and mafic melts. Geochim. Cosmochim. Acta 71 (1), 202-218

Jonasson K. (1994) Rhyolite volcanism in the Krafla central volcano, north-east Iceland. Bull.

Volcanol. 56(6-7), 516-528

Jönasson K. (2007) Silicic volcanism in Iceland: Composition and distribution within the active

volcanic zones. J. Geodynam. 43 (1), 101-117

Kasemann S.A., Jeffcoate A.B. and Elliott T. (2005) Lithium isotope composition of basalt glassreference material. Anal. Chem. 11, 5251-5257

Kay R.W. (1978) Aleutian magnesian andésites: melts from subducted Pacific Ocean crust. J.

Volcanol. Geotherm. Res. 4, 117-132

Kelemen P.B. (1995) Genesis of high Mg andésites and the continental crust. Contrib. Mineral.

Petrol. 120(1), 1-19

Kelley K.A., Plank T., Farr L., Ludden J., and Staudigel H. (2005) Subduction cycling of U, Th, and

Pb. Earth Planet. Sei. Lett. 234 (3-4), 369-383

Kempton P.D., Fitton J.G., Saunders A.D., Nowell CM., Taylor R.N., Hardarson B.S., and Pearson

G. (2000) The Iceland plume in space and time: a Sr-Nd-Pb-Hf study of the North Atlantic rifted

margin. Earth Planet, Sei. Lett, 111 (3-4), 255-271

Kinzler R.J., Donnelly-Nolan J.M. and Grove T.L. (2000) Late Holocene hydrous mafic magmatismat the Paint Pot Crater and Callahan flows, Medicine Lake Volcano, N. California and the

influence of H20 in the generation of silicic magma. Contrib. Mineral. Petrol. 138, 1-16

Kirsten T., Horn P. and Kiko J. (1973) 39Ar-40Ar dating and rare gas analysis of Apollo 16 rocks and

soils. Proc. Lunar Sei. Conf 4th, 1757-1784

Kleine T., Mezger K. and Münker C. (2003) Constraints on the age of the Moon from 182W. Annu.

Meteorit. Soc. Meet. 66th, #5212

Knauth D.C., Federman S.R., Lambert D.L. and Crane P. (2000) Newly synthesized lithium in the

interstellar medium. Nature 405, 656-658

Knauth D.C., Federman S.R. and Lambert D.L. (2003) An ultra-high-resolution survey of the

interstellar 7Li/6Li isotope ratio in the solar neighborhood. Astrophys. J. 586, 268-285

Kobayashi K., Tanaka R., Moriguti T., Shimizu K, and Nakamura E. (2004) Lithium, boron and lead

isotope systematics of glass inclusions in olivines from Hawaiian lavas: evidence for recycledcomponents in the Hawaiian plume. Chem. Geol. 212, 143-161

Köhler T.P. and Brey G.P. (1990) Calcium exchange between olivine and clinopyroxene calibrated as

a geothermobarometer for natural peridotites from 2 to 30 kb with applications. Geochim.

Cosmochim. Acta 54, 2375-2388

127

References

Korotev R.L. and Kremser D.T. (1992) Compositional variations in Apollo 17 soils and their

relationship to the geology of the Taurus-Littrow site. Proc. Lunar Planet. Sei. Conf. 22nd, 275-

301

Kosler J„ Kucera M. and Sylvester P. (2001) Precise measurement of Li isotopes in planktonicforaminiferal tests by quadrupole ICPMS. Chem. Geol. 181, 169-179

Kosler J., Magna T., Sylvester P. and Pedersen R.-B. (2003) Precision and reproducibility of Li

isotopic analyses by TCPMS - comparison of single- and multi-collector measurements.

Geoanalysis 2003. Abstract volume, 48. Rovaniemi, Finland.

Kurz M.D., Curtice J., Dempsey E.L.I., and Solow A. (2004) Rapid helium isotope variability in

Mauna Kea shield lavas from the Hawaiian Scientific Drilling Project. G-cubed 5 (4), paper

number Q04G14,doi: 10.1029/2002GC000439

Leeman W.P., Tonarini S., Chan L.H. and Borg L.E. (2004) Boron and lithium isotopic variations in a

hot subduction zone - the southern Washington Cascades. Chem. Geol. 212, 101-124

Lemoine M., Vangioni-Flam E. and Cassé M. (1998) Galactic cosmic rays and the evolution of lightelements. Astrophys. J. 499, 735-745

Lentz R.C.F., McSween H.Y., Ryan J. and Riciputi L.R. (2001) Water in martian magmas: Clues from

light lithophile elements in shergottite and nakhlite pyroxenes. Geochim. Cosmochim. Acta 65

(24), 4551-4565

Longhi J. (1992) Origin of picritic green glass magmas by polybaric fractional fusion. Proc. Lunar

Planet. Sei. Conf. 22nd, 343-353

Lundstrom C.C., Gill J. and Williams Q. (2000) A geochcmically consistent hypothesis for MORB

generation. Chem. Geol. 162, 105-126

Lundstrom C.C., Chaussidon M., Hsui A.T., Kelemen P. and Zimmerman M. (2005) Observations of

Li isotopic variations in the Trinity Ophiolite: Evidence for isotopic fractionation by diffusion

during mantle melting. Geochim. Cosmochim. Acta 69 (3), 735-751

Lynton S.J., Walker R.J., and Candela P.A. (2005) Lithium isotopes in the system Qz-Ms-fluid:An

experimental study. Geochim. Cosmochim. Acta 69 (13), 3337-3347

Maaloe S., Sorensen I., and Hertogen J. (1986) The trachybasaltic suite of Jan Mayen. J. Petrol. 27

(2), 439-466

Macpherson CG, Hilton D.R., Day J.M.D., Lowry D., and Grönvold K. (2005) High-3He/4He,depleted mantle and low-5180, recycled oceanic lithosphère in the source of central Iceland

magmatism. Earth Planet. Sei. Lett. 233 (3-4), 411-427

Magna T., Harrison D., Wiechert U.H. and Halliday A.N. (2004a) Coupled Li-He isotope systematicsin Icelandic basalt glasses. Geochim. Cosmochim. Acta 68, A48

Magna T., Wiechert U.H. and Halliday A.N. (2004b) Low-blank separation and isotope ratio

measurement of small samples of lithium using multiple-collector ICPMS. Int. J. Mass Spectrom.239(l),67-76

Magna T., Wiechert U., Grove T.L., and Halliday A.N. (2006a) Lithium isotope fractionation in the

Southern Cascadia subduction zone. Earth Planet. Set. Lett. 250 (3-4), 428-443

Magna T., Wiechert U. and Halliday A.N. (2006b) New constraints on the lithium isotope

compositions of the Moon and terrestrial planets. Earth Planet. Sei. Lett. 243 (3-4), 336-353

128

Référença

Marschall H.R., Altherr R., Ludwig T., Kalt A., Gmeling K., and Kasztovszky Z. (2006) Partitioningand budget of Li, Be and B in high-pressure metamorphic rocks. Geochim. Cosmochim. Acta 70

(18), 4750-4769

Marschall H., Altherr R., and Lüpkc L. (2007) Squeezing out the slab - modelling the release of Li,

Be, and B during progressive high-pressure metamorphism. Chem. Geol. 239 (3-4), 323-335

Martin E. and Sigmarsson O. (2007) Crustal thermal state and origin of silicic magma in Iceland: the

case of Torfajökull, Ljösfjöll and Snœfellsjokull volcanoes. Contrib. Mineral. Petrol. 153 (5),593-605

McDonough W.F., Teng F.-Z., Tomascak P.B., Ash R.D., Grossman J.N. and Rudnick R.L. (2003)Lithium isotopic composition of chondritic meteorites. Lunar Planet. Sei. Conf. XXXIV, #1931

McDonough W.F., Teng F.-Z., Rudnick R.L., and Ash R.D. (2006) Lithium isotopic analyses of

chondrites and chondmles. Lunar Planet. Sei. Conf. XXXVII, #2416

Meier A.L. (1982) Determination of lithium isotopes at natural abundance levels by atomic absorption

spectrometry. Anal. Chem. 54, 2158-2162

Melosh H.J. (1984) Impact ejection, spallation and the origin of meteorites. Icarus 59, 234-260

Mertz D.F., Devey C.W., Todt W., Stoffers P., and Hofmann A.W. (1991) Sr-Nd-Pb isotope evidence

against plume-asthenosphere mixing north of Iceland. Earth Planet. Sei. Lett. 107 (2), 243-255

Mertz D.F., Sharp W.D., and Haase K.M. (2004) Volcanism on the Eggvin Bank (Central Norwegian-Greenland Sea, latitude -71 °N): age, source, and relationship to the Iceland and putative Jan

Mayen plumes. J. Geodynam, 38 (1), 57-83

Millot R., Guerrot C and Vigier N. (2004) Accurate and high-precision measurement of lithium

isotopes in two reference materials by MC-ICP-MS. Geostandards Geoanal. Res. 28 ( 1 ), 153-159

Montalban J. and Rebolo R. (2002) Planet accretion and the abundances of lithium isotopes. Aslron.

Astrophys. 386, 1039-1043

Moriguti T. and Nakamura E. (1998a) High-yield lithium separation and the precise isotopic analysisfor natural rock and aqueous samples. Chem. Geol. 145 (1-2), 91-104

Moriguti T. and Nakamura E. (1998b) Across-arc variation of Li isotopes in lavas and implication for

crust/mantle recycling at subduction zones. Earth Planet. Sei. Lett. 163 (1-4), 167-174

Moriguti T., Shibata T. and Nakamura E. (2004) Lithium, boron and lead isotope and trace element

systematics of Quarternary basaltic volcanic rocks in northeastern Japan: mineralogical controls

on slab-derived fluid composition. Chem. Geol. 212, 81-100

Muehlenbachs K., Fumes II., Fonneland U.C., and Hellevang B. (2004) Ophiolites as faithful records

of ancient seawater: the Solund-Stavfjord Ophiolite Complex as a Late Ordovician example. In

Ophiolites in Earth history (ed. Y. Dilek and P.T. Robinson). Geol. Soc. Lond. Spec. Publ. 218, 1-

14

Newman S., Macdougall J.D. and Finkel R.C. (1986) Petrogenesis and 230Th-23SU disequilibrium at

Mt.Shasta, California, and in the Cascades. Contrib. Mineral. Petrol. 93, 195-206

Nishio Y. and Nakai S. (2002) Accurate and precise lithium isotopic determinations of igneous rock

samples using multi-collector inductively coupled plasma mass spectormetry. Anal. Chim. Acta

456,271-281

129

Referenc es

Nishio Y., Nakai S., Yamamoto J., Sumino H., Matsumoto T., Prikhod'ko V.S. and Arai S. (2004)Lithium isotopic systematics of the mantle-derived ultramafic xenoliths: implications for EMI

origin. Earth Planet. Sei. Lett. 217, 245-261

Nishio Y., Nakai S., Kogiso T., and Barsczus H.G. (2005) Lithium, strontium, and neodymium

isotopic compositions of oceanic island basalts in the Polynesian region: constraints on a

Polynesian HIMU origin. Geochem../. 39 (1), 91-103

Niu Y. and O'Hara M.J. (2003) Origin of ocean island basalts: a new perspective from petrology,

geochemistry and mineral physics considerations. J. Geophys. Res. 108 (B4), paper number 2209,

doi : 10.1029/2002JB002048

Nyquist L.E., Wiesmann H., Bansal B., Shih C-Y., Keith J.E. and Harper C.L. (1995) '46Sm-,42Nd

formation interval for the lunar mantle. Geochim. Cosmochim. Acta 59 (13), 2817-2837

Oi T., Odagiri T. and Nomura M. (1997) Extraction of lithium from GSJ rock reference samples and

determination of their isotopic compositions. Anal. Chim. Acta 340, 221-225

Olive K.A. and Schramm D.N. (1992) Astrophysical 7Li as a product of Big Bang nucleosynthesis and

galactic cosmic-ray spallation. Nature 360, 439-442

O'Neill Il.S.C (1991) The origin of the Moon and the early history of the Earth - A chemical model.

Part 1: The Moon. Geochim. Cosmochim. Acta 55, 1135-1157

Papike J.J., Bence A.E. and Lindsley D.H. (1974) Mare basalts from the Taurus-Littrow region of the

moon. Proc. Lunar Sei. Conf. 5th, 471-504

Paquin J. and Altherr R. (2002) Subduction-related lithium metasomatism during exhumation of the

Alpe Arami ultrahigh-pressure garnet peridotite (Central Alps, Switzerland). Contrib. Mineral.

Petrol. 143, 632-640

Paquin J., Altherr R. and Ludwig T. (2004) Li-Be-B systematics in the ultrahigh-pressure garnet

peridotite from Alpe Arami (Central Swiss Alps): implications for slab-to-mantle wedge transfer.

Earth Planet. Sei. Lett. 218, 507-519

Parman S.W. (2001) Petrology and geochemistry of high-degree mantle melts. PhD thesis, MIT

Cambridge, Massachusetts, USA.

Pearce N.J.G., Perkins W.T., Westgate J.A., Gorton M.P., Jackson S.E., Neal C.R. and Chenery S.P.

(1997) A compilation of new and published major and trace element data for N1ST SRM 610 and

N1ST SRM 612 glass reference materials. Geostandards Newslett. 21 (1), 115-144

Pistincr J.S. and Henderson G.M. (2003) Lithium-isotope fractionation during continental weathering

processes. Earth Planet. Sei. Lett. 214 (1-2), 327-339

Pogge von Strandmann P.A.E., Burton K.W., James R.H., van Calstercn P., Gislason S.R., and

Mokadem F. (2006) Riverine behaviour of uranium and lithium isotopes in an actively glaciatedbasaltic terrain. Earth Planet. Sei. Lett. 251 (1-2), 134-147

Poitrasson F., Halliday A.N., Lee D.-C, Levasseur S. and Teutsch N. (2004) Iron isotope differences

between Earth, Moon, Mars and Vesta as possible records of contrasted accretion mechanisms.

Earth Planet. Sei. Lett. 223, 253-266

Press S., Witt G., Seek H.A., Eonov D. and Kovalenko V.l. (1986) Spinel peridotite xenoliths from

the Tariat depression, Mongolia. 1. Major element chemistry and mineralogy of a primitivemantle xenolith suite. Geochim. Cosmochim. Acta 50 (12), 2587-2599

Qi H.P., Coplen T.B., Wang Q.Z. and Wang Y.H. (1997) Unnatural isotopic composition of lithium

reagents. Anal. Chem. 69, 4076-4078

130

References

Ramaty R., Kozlovsky B. and Lingenfelter R.E. (1996) Light isotopes, extinct radioisotopes and

gamma-ray lines from low-energy cosmic-ray interactions. Astrophys. J. 456, 525-540

Rhodes J.M. and Hubbard N.J. (1973) Chemistry, classification and petrogenesis of Apollo 15 mare

basalts. Proc. Lunar Sei. Conf. 4th, 1127-1148

Rhodes J.M., Rodgers K.V., Shih C.-Y., Bansal B.M., Nyquist L.E., Wiesmann H. and Hubbard N.J.

(1974) The relationship between geology and soil chemistry at the Apollo 17 landing site. Proc.

Lunar Sei. Conf 5th, 1097-1117

Richter F.M., Davis A.M., DePaolo D.J. and Watson E.B. (2003) Isotope fractionation by chemical

diffusion between molten basalt and rhyolite. Geochim. Cosmochim. Acta 67 (20), 3905-3923

Ritzenhoff S., Schröter E.H. and Schmidt W. (1997) The lithium abundances in sunspots. Astron.

Astrophys. 328, 695-701

Rose E.F., Shimizu N., Layne G.D. and Grove T.L. (2001) Melt production beneath Mt. Shasta from

boron data in primitive melt inclusions. Science 293, 281-283

Rose H.J., Cuttitta F„ Berman S., Brown F.W., Carron M.K., Christian R.P., Dwornik E.J. and

Greenland L.P. (1974) Chemical composition of rocks and soils at Taurus-Littrow. Proc. Lunar

Sei. Conf. 5th, 1119-1133

Rose H.J., Baedecker P.A., Berman S., Christian R.P., Dwornik E.J., Finkelman R.B. and SchnepfeM.M. (1975) Chemical composition of rocks and soils returned by the Apollo 15, 16 and 17

missions. Proc. Lunar Sei. Conf. 6th, 1363-1373

Rudnick R.L., Barth M.G., Horn L, and McDonough W.F. (2000) Rutile-bearing refractory eclogites:

Missing link between continents and depleted mantle. Science 287, 278-281

Rudnick R.L., Tomascak P.B., Njo H.B., and Gardner L.R. (2004) Extreme lithium isotopicfractionation during continental weathering revealed in saprolites from South Carolina. Chem.

Geol. 212 (1-2), 45-57

Rudnick R. and lonov D.A. (2007) Lithium elemental and isotopic disequilibrium in minerals from

peridotitic xenoliths from far-east Russia: product of recent melt-fluid/rock reaction. Earth

Planet. Sei. Lett. 256 (1-2), 278-293

Ryan J.G. and Langmuir C.H. (1987) The systematics of lithium abundances in young volcanic rocks.

Geochim. Cosmochim. Acta 51, 1727-1741

Ryan J.G. and Kyle P.R. (2004) Lithium abundance and lithium isotope variations in the mantle

sources: insights from intraplate volcanic rocks from Ross Island and Marie Byrd Land

(Antarctica) and other oceanic islands. Chem. Geol. 212 (1-2), 125-142

Sahoo S.K. and Masuda A. (1995) High precision isotopic measurement of lithium by thermal

ionization mass spectrometry. Int. J. Mass Spectrom. Ion Process. 151 (2-3), 189-196

Scambelluri M., Müntener O., Ottolini L., Pettke T.T. and Vannucci R. (2004) The fate of B, CI and

Li in the subducted oceanic mantle and in the antigorite breakdown fluids. Earth Planet. Sei. Lett.

222(l),217-234

Schilling J.-G, Kingsley R., Fontignie D., Poreda R., and Xuc S. (1999) Dispersion of the Jan Mayenand Iceland mantle plumes in the Arctic: A He-Pb-Nd-Sr isotope tracer study of basalts from the

Kolbeinsey, Mohns and Knipovich Ridges. J. Geophys. Res. 104 (B5), 10543-10569

Schou M. (1988) Lithium treatment of manic-depressive illness - Past, present and perspectives. J.

Am, Med. Assoc. 259, 1834-1836

131

Referenc es

Schramm A. (1993) Geochemische Charakterisierung ausgewählter Eklogite (MünchbergerGneismasse, Beilsbank, Roberts Victor, Zagadochnaya). MSc. thesis, Georg-August Univ. 33 pp.

Seitz H.-M. and Woodland A.B. (2000) The distribution of lithium in peridotitic and pyroxeniticmantle lithologies - an indicator ofmagmatic and metasomatic processes. Chem. Geol. 166, 47-64

Seitz H.-M., Brey G.P., Lahaye Y., Durali S. and Weyer S. (2004) Lithium isotopic signatures of

peridotite xenoliths and isotopic fractionation at high temperature between olivine and pyroxenes.

Chem. Geol. 212 (1-2), 163-177

Seitz H.-M., Brey G.P., Harris J.W., and Ludwig T. (2006a) Lithium isotope composition of lower

mantle ferropericlase inclusions in diamonds from Sao Luiz, Brazil. Geochim. Cosmochim. Acta

70(18), A17

Seitz H.-M., Brey G.P., Weyer S., Zipfel J, Ott U., and Durah S. (2006b) Lithium isotope

composition of ordinary and carbonaceous chondrites, and differentiated planetary bodies: Bulk

solar system and solar reservoirs. Lunar Planet. Sei. Conf. XXXVII, #1708

Sephton M.A., James R.H. and Bland P.A. (2004) Lithium isotope analyses of inorganic constituents

from the Murchison meteorite. Astrophys. J. 612, 588-591

Seyfried W.E.J., Chen X. and Chan L.H. (1998) Trace element mobility and lithium isotope exchange

during hydrothermal alteration of seafloor weathered basalt: An experimental study at 350°C, 500

bars. Geochim. Cosmochim. Acta 62 (6), 949-960

Shearer CK. and Papike J.J. (1989) Is plagioclase removal responsible for the negative Eu anomaly in

the source regions of mare basalts? Geochim. Cosmochim. Acta 53, 3331-3336

Shearer CK. and Papike J.J. (1993) Basaltic magmatism on the Moon: A perspective from volcanic

picritic glass beads. Geochim. Cosmochim. Acta 57, 4785-4812

Shearer C.K. and Papike J.J. (1999) Magmatic evolution of the Moon. Am. Mineral. 84, 1469-1494

Sigmarsson O., Hémond C, Condomines M., Fourcade S., and Oskarsson N. (1991) Origin of silicic

magma in Iceland revealed by Th isotopes. Geology 19 (6), 621-624.

Sigmarsson O., Condomines M., and Fourcade S. (1992) A detailed Th, Sr and O isotope study of

Hekla: differentiation processes in an Icelandic volcano. Contrib. Mineral. Petrol. 112(1), 20-34

Sirbescu M.-L.C and Nabelek P.I. (2003) Crustal melts below 400°C Geology 31 (8), 685-688

Sisson T.W. and Layne G.D. (1993) H20 in basalt and basaltic andésite glass inclusions from four

subduction-related volcanoes. Earth Planet. Sei. Lett. 117, 619-635

Skovgaard A.C., Storey M., Baker J., Bluzstajn J., and Hart S.R. (2001) Osmium-oxygen isotopicevidence for a recycled and strongly depleted component in the Iceland mantle plume. Earth

Planet. Sei. Lett. 194 (1-2), 259-275

Smales A.A. and Webster R.K. (1958) The determination of lithium in rocks by the method of stable

isotope dilution. Anal, Chim, Acta 18, 587-596

Smith B.W., Gornuskhin LB., King L.A. and Winefordner J.D. (1998) A laser ablation-atomic

fluorescence technique for isotopically selective determination of lithium in solids. Spectrochim.Acta(B) 53, 1131-1138

Snyder G.A., Taylor L.A. and Neal CR. (1990) The sources of mare basalts: A model involving lunar

magma ocean crystallization, plagioclase flotation and trapped instantaneous residual liquid.

Workshop on mare basalt petrogenesis, LPI Houston, pp. 45-46

132

Referent, es

Stettier A., Eberhardt P., Geiss J., Grögler N. and Maurer P. (1973) Ar^-Ar40 ages and Ar^-Ar38

exposure ages of lunar rocks. Proc. Lunar Sei. Conf. 4th, 1865-1888

Stracke A., Zindler A., Salters V.J.M., McKenzie D., Blichert-Toft J., Albarèdc F., and Grönvold K.

(2003) Theistareykir revisited. G-cubed 4 (2), paper number 8507, doi:10.1029/2001GC000201

Strelow F.W.E., Weinert C.H.S.W. and van der Walt T.N. (1974) Separation of lithium from sodium,

beryllium and other elements by cation-exchange chromatography in nitric acid-methanol. Anal.

Chim. Acta 11, 123-132

Stuart F.M., Lass-Evans S., Fitton J.G., and Ellam R.M. (2003) High \He/4He ratios in picritic basalts

from Baffin Island and the role of a mixed reservoir in the mantle plumes. Nature 424, 57-59

Stuart F.M., Ellam R.M., and Fitton J.G. Constraints on mantle enrichments in the North Atlantic

ocean from the geochemistry of basalts from Jan Mayen. J, Petrol., submitted manuscript

Sun S.-S. and McDonough W.F. (1989) Chemical and isotopic systematics of oceanic basalts:

implications for mantle composition and processes. In Magmatism in the ocean basins, Vol. 42

(cd. A.D. Saunders and M.J. Norry), pp. 313-345. Blackwell Scientific Publ.

Suzuki T.K., Yuzuru Y. and Beers T.C (2000) Primordial lithium abundance as a stringent constraint

on the baryonic content of the universe. Astrophys. J. 540, 99-103

àulcek Z., Povondra P. and Stangl R. (1965) Chromatographische Trennung der Lithium- und

Natriumionen. Collect. Czechoslovak Chem. Commun. 30, 380-387

èulcck Z. and Rubeäka J. (1969) Bestimmung des Lithiumspurengehaltes in Gesteinen. Collect.

Czechoslovak Chem. Commun. 34, 2048-2056

Tatsumi Y. and Maruyama S. (1989) Boninites and high-Mg andésites: tectonics and petrogenesis. In

Boninites (ed. A.J. Crawford), pp. 50-71. Chapman and Hall.

Taylor S.R., Gorton M.P., Muir P., Nance W., Rudowski R. and Ware N. (1973) Lunar highlandscomposition: Apennine Front. Proc. Lunar Sei. Conf. 4th, 1445-1459

Taylor S.R. and Jakes' P. (1974) The geochemical evolution of the moon. Proc. Lunar Sei. Conf. 5th,

1287-1305

Taylor T.I. and Urey H.C (1938) Fractionation of the lithium and potassium isotopes by chemical

exchange with zeolithes. J. Chem. Phys. 6, 429-438

Teng F.-Z., McDonough W.F., Rudnick R.L., Dalpé C, Tomascak P.B., Chappell B.W. and Gao S.

(2004) Lithium isotopic composition and concentration of the upper continental crust. Geochim.

Cosmochim. Acta 68 (20), 4167-4178

Teng F.-Z., McDonough W.F., Rudnick R.L., and Walker R.J. (2006a) Diffusion-driven extreme

lithium isotopic fractionation in country rocks of the Tin Mountain pegmatite. Earth Planet. Sei.

Lett. 243(3-4), 701-710

Teng F.-Z., McDonough W.F., Rudnick R.L., Walker R.J., and Sirbescu M.-L.C. (2006b) Lithium

isotopic systematics of granites and pegmatites from the Black Hills, South Dakota. Am. Mineral.

91(10), 1488-1498

Thirlwall M.F. (1997) Pb isotopic and elemental evidence for OIB derivation from young HIMU

mantle. Chem. Geol. 139 (1-4), 51-74

Thirlwall M.F., Gee M.A.M., Taylor R.N., and Murton B.J. (2004) Mantle components in Iceland and

adjacent ridges investigated using double-spike Pb isotope ratios. Geochim. Cosmochim. Acta 68

(2), 361-386

133

References

Thirlwall M.F., Gee M.A.M., Lowry D., Mattey D.P., Murton B.J., and Taylor R.N. (2006) Low 8li!0

in the Icelandic mantle and its origin: Evidence from Reykjanes Ridge and Icelandic lavas.

Geochim. Cosmochim. Acta 70 (4), 993-1019

Tomascak P.B. and Tera F. (1997) An analytical method for precise lithium isotope determination,

with applications to basaltic rocks. EOS 78, F 737

Tomascak P.B. and Langmuir CH. (1999) Lithium isotope variability in MORB. EOS 80, F 1086-

1087

Tomascak P.B., Carlson R.W. and Shirey S.B. (1999a) Accurate and precise determination of Li

isotopic compositions by multi-collector sector ICP-MS. Chem. Geol. 158(1-2), 145-154

Tomascak P.B., Tera F., Heiz R.T. and Walker R.J. (1999b) The absence of lithium isotopefractionation during basalt differentiation: new measurements by multicollector sector ICP-MS.

Geochim. Cosmochim. Acta 63 (6), 907-910

Tomascak P.B., Ryan J.G. and Defant M,J. (2000) Lithium isotope evidence for light element

decoupling in the Panama subarc mantle. Geology 28 (6), 507-510

Tomascak P.B., Widom E., Benton L.D., Goldstein S.L. and Ryan J.G. (2002) The control of lithium

budgets in island arcs. Earth Planet. Sei. Lett. 196, 227-238

Tomascak P.B., Hemming N.G., and Hemming S.R. (2003) The lithium isotopic composition of

waters of the Mono Basin, California. Geochim. Cosmochim. Acta 67 (4), 601-611

Tomascak P.B. (2004) Developments in the understanding and application of lithium isotopes in the

earth and planetary sciences. Rev. Mineral. Geochem, 55, 153-195

Tomascak P.B., Langmuir C.H., LeRoux P., and Shirey S.B. Lithium isotopes in global mid-ocean

ridge basalts. Geochim. Cosmochim. Acta, in press

Tronnes R.G. (1990) Basaltic melt evolution of the Hengill volcanic system, SW Iceland, and

evidence for clinopyroxene assimilation in primitive tholeiitic magmas. J. Geophys. Res. 95 (10),

15893-15910

Tronnes R.G, Planke S., Sundvoll B., and Imsland P. (1999) Recent volcanic rocks from Jan Mayen:

Low-degree melt fractions of enriched northeast Atlantic mantle. J. Geophys. Res. 104 (B4),

7153-7168

Ulmer P. and Trommsdorff V. (1995) Serpentine stability to mantle depths and subduction-related

magmatism. Science 268, 858-861

Vigier N., Burton K.W., Gislason S.R., Roger N.W., Schaefer B.F., and James R.H. (2002)Constraints on basalt erosion from Li isotopes and U-series nuclides measured in Iceland rivers.

Geochim, Cosmochim. Acta 66, A806

Volpe A.M. (1992) 238U-230Th-226Ra disequilibrium in young Mt. Shasta andésites and dacites. J.

Volcanol. Geotherm. Res. 53, 227-238

Wagner T.P., Donnelly-Nolan J.M. and Grove T.L. (1995) Evidence of hydrous differentiation and

crystal accumulation in the low-MgO, high-Al203 Lake Basalt from Medicine Lake volcano,

California. Contrib. Mineral. Petrol. 121, 201-216

Weyer S., Münker C. and Mezger K. (2003) Nb/Ta, Zr/Hf and REE in the depleted mantle:

implications for the differentiation history of the crust-mantle system. Earth Planet. Sei. Lett.

205, 309-324

134

Références

Wiechert U., lonov D.A. and Wedepohl K.H. (1997) Spinel peridotite xenoliths from the Atsagin-Dush volcano, Dariganga lava plateau, Mongolia: a record of partial melting and crypticmetasomatism in the upper mantle. Contrib. Mineral. Petrol. 126, 345-364

Wiechert U., Halliday A.N., Lee D.-C, Snyder G.A., Taylor L.A. and Rumble D. (2001) Oxygen

isotopes and the Moon-forming giant impact. Science 294, 345-348

Wiechert U.H., Halliday A.N., Palme H. and Rumble D. (2004) Oxygen isotope evidence for rapid

mixing of the HED meteorite parent body. Earth Planet. Sei. Lett. 221, 373-382

Wieler R. and Baur H. (1995) Fractionation of Xe, Kr, and Ar in the Solar corpuscular radiation

deduced by closed system etching of lunar soils. Astrophys. J. 453, 987-997

Willbold M. and Stracke A. (2006) Trace element composition of mantle end-members: Implicationsfor recycling of oceanic and upper and lower continental crust. G-cubed 7 (4), paper number

Q04004, doi: 10.1029/2005GC001005

Wilson D. (1988) Tectonic history of the Juan de Fuca ridge over the last 40 million years. J.

Geophys. Res. 93, 11863-11876

Wolfe C.J., Bjarnason I.T., VanDecar J.C., and Solomon S.C (1997) Seismic structure of the Iceland

mantle plume. Nature 385, 245-247

Woodhead J.D. and Hergt J.M. (2000) Pb-isotope analyses of USGS reference materials.

Geostandards Newslett. 24 (1), 33-38

Woodland A.B., Seitz H.-M., Altherr R., Marschall H., Olker B., and Ludwig T. (2002) Li abundances

in eclogite minerals: a clue to a crustal or mantle origin? Contrib. Mineral. Petrol. 143, 587-601

Workman R.K., Hart S.R., Jackson M., Regelous M., Farley K.A., Bluzstajn J., Kurz M., and

Staudigel H. (2004) Recycled metasomatised lithosphère as the origin of the Enriched Mantle II

(EM2) end-member: Evidence from the Samoan Volcanic Chain. G-cubed 5 (4), paper number

Q04008,doi : 10.1029/2003GC000623

Wunder B., Meixner A., Romer R.L., and Heinrich W. (2006) Temperature-dependent isotopicfractionation of lithium between clinopyroxene and high-pressure hydrous fluids. Contrib.

Mineral. Petrol. 151, 1 12-120

Yogodzinski G.M., Kay R.W., Volynets O.N., Koloskov A.V. and Kay S.M. (1995) Magnesianandésite in the western Aleutian Komandorsky region: Implications for slab melting and

processes in the mantle wedge. Geol. Soc. Am. Bull. 107 (5), 505-519

Yokoyama Y., Reyss J.L. and Giuchard F. (1974) 22Na-2f'Al chronology of lunar surface processes.

Proc. Lunar Sei, Conf. 5th, 2231-2247

You C.-F., Chan L.H., Spivack A.J. and Gieskes J.M. (1995) Lithium, boron and their isotopes in

sediments and pore waters of Ocean Drilling Program Site 808, Nankai trough: Implications for

fluid expulsion in accretionary prisms. Geology 23 (1), 37-40

You C.-F. and Chan L.H. (1996) Precise determination of lithium isotopic composition in low

concentration natural samples. Geochim. Cosmochim. Acta 60 (5), 909-915

Zack T., Tomascak P.B., Rudnick R.L., Dalpé C. and McDonough W.F. (2003) Extremely light Li in

orogenic eclogites: The role of isotope fractionation during dehydration in subducted oceanic

crust. Earth Planet. Sei. Lett. 208 (3-4), 279-290

Zhang L., Chan L.H. and Gieskes J.M. (1998) Lithium isotope geochemistry of pore waters from

Ocean Drilling Program sites 918 and 919, Irminger Basin. Geochim. Cosmochim, Acta 62 (14),2437-2450

135

References

Zhu X.K., Makishima A., Guo Y., Belshaw N.S. and O'Nions R.K. (2002) High precisionmeasurements of titanium isotope ratios by plasma source mass spectrometry. Int. J. Mass

Spectrom. 220,21-29

Zucca J.J., Fuis GS., Milkereit B., Mooney W.D. and Catchings R.D. (1986) Crustal structure of

Northeastern California. J. Geophys. Res. 91 (B7), 7359-7382

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Tomas Magna

Curriculum Vitae

Date ofbirth:

Place ofbirth:

Nationality:

March 3, 1977

Chomutov, Czech Republic

Czech

Education

1984-1991 Primary school, Chomutov, Czech Republic

1991-1995 State Gymnasium, Chomutov, Czech Republic

1995-1998 Bachelor study - Geology, Charles University in Prague, Czech Republic

1998-2000 Diploma study - Geochemistry, Charles University in Prague, Czech Republic

Diploma Thesis: Determination of Sm and Nd concentrations using isotope

dilution inductively coupled plasma quadrupole mass spectrometry

2001-2006 Doctoral study, Institute of Isotope Geochemistry and Mineral Resources,

ETH Zürich, Switzerland

PhD Thesis: Lithium isotopes in the Earth and terrestrial planets

137

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