Experimental Constraints on Lithium Exchange between ...

Experimental Constraints on Lithium Exchange between Clinopyroxene, Olivine and Aqueous Fluid at High Pressures and Temperatures

by

Natalie Caciagli Warman

A thesis submitted in conformity with the requirements for the degree of Doctor of Philosophy Department of Geology University of Toronto

© Copyright by Natalie Caciagli Warman 2010

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Experimental Constraints of Lithium Exchange between Clinopyroxene, Olivine and Aqueous Fluid at High Pressures and Temperatures Natalie Caciagli Warman Doctor of Philosophy Department of Geology University of Toronto 2009

Abstract Clinopyroxene, olivine, plagioclase and hydrous fluid lithium partition coefficients have been measured between 800-1100oC at 1 GPa. Clinopyroxene-fluid partitioning is a function of temperature (ln DLicpx/fluid = -7.3 (+0.5) + 7.0 (+0.7) * 1000/T) and appears to increase with increasing pyroxene Al2O3 content. Olivine-fluid partitioning of lithium is a function of temperature (ln DLiol/fluid = -6.0 (+2.0) + 6.5 (+2.0) * 1000/T) and appears to be sensitive to olivine Mg/Fe content. Anorthite-fluid lithium partitioning is a function of feldspar composition, similar to the partitioning of other cations in the feldspar-fluid system. Isotopic fractionation between clinopyroxene and fluid, Licpx-fluid, has been measured between 900-1100oC and ranges from 0.3 to -3.4 ‰ (±1.4 ‰). Lithium diffusion has been measured in clinopyroxene at 800-1000oC and in olivine at 1000oC. The lithium diffusion coefficient is independent of the diffusion gradient as values are the same if the flux of lithium is into or out of the crystal and ranges from -15.19 ± 2.86 m2/s at 800oC to 11.97 ± 0.86 m2/s at 1000oC. Lithium diffusion in olivine was found to be two orders of magnitude slower than for clinopyroxene at similar conditions. Closure temperatures calculated

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for lithium diffusion in clinopyroxene range from ~400 to ~600oC. These results demonstrate that lithium equilibration between fluids and minerals is instantaneous, on a geological timescales. The confirmation of instantaneous equilibration, combined with min-fluid partition coefficients and values for Licpx-fluid, permits quantitative modeling of the evolution of lithium concentration and isotopic composition in slab-derived fluids during transport to the arc melt source. Our results indicate that fluids migrating by porous flow will rapidly exchange lithium with the mantle, effectively buffering the fluid composition close to ambient mantle values, and rapidly attenuating the slab lithium signature. Fluid transport mechanisms involving fracture flow are required to maintain a slab-like lithium signature (both elemental and isotopic) from the slab to the melt source of island arc basalts. This study demonstrates that mineral-fluid equilibration is rapid, and as a result the lithium content of minerals can only reliably represent chemical exchange in the very latest stages of the sample’s history.

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Acknowledgments This thesis dissertation marks the conclusion of work that began in 2001. There were many who helped, supported, and cheered me on my way and I fear that to be able to acknowledge everyone who assisted me would be an impossible task. I am certain that once this work has been submitted I will realize that I have left out several important people. First, I must thank Dr. James Brenan, my mentor and supervisor, who never gave up on me despite everything. I am indebted to Dr. Lesley Rose Weston (my lab partner in crime) and Dr. Boris Foursenko for all their time and invaluable mechanical, technical, and moral support. I am grateful to my collaborators at Lawrence Livermore National Laboratory, Dr. Doug Phinney, Dr. Ian Hutcheon and Dr. Rick Ryerson for access to and assistance with the SIMS. Thanks to Dr. Bill McDonough and his lab at University of Maryland for the isotopic analyses. I also wish to thank Dr. Grey Bebout at Lehigh University, for his patience and understanding when, to quote Edison, “I found 10,000 ways that won’t work.” before I found one way that did work. I would also like to thank Dr. Paul Tomascak, who pointed me in the direction that would bear the most fruit and Dr. Jon Davidson for his unwavering encouragement for always having time for a ‘little chat’. Mainly I am indebted to my husband Tim, who never once let me give up on myself. This work must be dedicated to my children, Olivia and Henry, whose existence shaped this project more than any other thing.

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Table of Contents

Acknowledgments.......................................................................................................................... iv Table of Contents............................................................................................................................ v List of Tables ................................................................................................................................. ix List of Figures ................................................................................................................................. x List of Appendices ....................................................................................................................... xiii 1 Introduction ................................................................................................................................ 1 1.1 Chemical Properties of Lithium.......................................................................................... 1 1.2 Sources and Concentrations of Lithium in the Earth .......................................................... 3 1.2.1

Subducted Materials (AOC and Seafloor Sediments) ............................................ 3

1.2.2

Eclogites.................................................................................................................. 4

1.2.3

Upper Mantle .......................................................................................................... 5

1.2.4

Convergent Margin Magmas .................................................................................. 6

1.3 The scale of lithium heterogenities ..................................................................................... 7 1.4 Previous experimental work ............................................................................................... 8 1.5 Focus of Thesis and Distribution of Work.......................................................................... 9 2 Lithium Partitioning and Isotopic Fractionation ...................................................................... 14 2.1 Introduction....................................................................................................................... 14 2.2 Methods............................................................................................................................. 16 2.3 Analytical Techniques ...................................................................................................... 18 2.3.1

Major Element Analyses....................................................................................... 18

2.3.2

Lithium Analyses .................................................................................................. 19 2.3.2.1 MC-ICPMS............................................................................................. 19 2.3.2.2 SIMS....................................................................................................... 19

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2.3.2.3 LA-ICPMS ............................................................................................. 20 2.4 Results............................................................................................................................... 21 2.4.1

Major Element Chemistry..................................................................................... 21

2.4.2

Mineral – Fluid Lithium Partitioning.................................................................... 21 2.4.2.1 Clinopyroxene ........................................................................................ 23 2.4.2.2 Olivine .................................................................................................... 24 2.4.2.3 Plagioclase .............................................................................................. 25

2.4.3

Olivine-Clinopyroxene Pair Experiments............................................................. 26 2.4.3.1 Experiments at variable fO2 .................................................................... 27 2.4.3.2 Experiments with added REE................................................................. 27

2.4.4

Lithium Isotope Fractionation............................................................................... 28 2.4.4.1 Clinopyroxene- fluid .............................................................................. 28 2.4.4.2 Olivine-Clinopyroxene Lithium Isotope Fractionation .......................... 29

2.5 Discussion ......................................................................................................................... 29 2.5.1

Controls on Partitioning........................................................................................ 29 2.5.1.1 Clinopyroxene ........................................................................................ 29 2.5.1.2 Olivine .................................................................................................... 31 2.5.1.3 Plagioclase .............................................................................................. 31 2.5.1.4 Intermineral Partitioning......................................................................... 32

2.5.2

Controls on Isotopic Fractionation........................................................................ 33

2.5.3

Lithium Incorporation into the Mantle ................................................................. 34

2.5.4

The Mantle Wedge as a Chromatograph .............................................................. 35

2.5.5

Isotopic Evolution of Lithium-Bearing Fluids in the Mantle ............................... 39 2.5.5.1 Percolation and Rayleigh Distillation..................................................... 39 2.5.5.2 Generation of 6Li-rich fluids................................................................... 41 2.5.5.3 Generation of 6Li-rich zones in the mantle............................................. 42

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2.6 Conclusions....................................................................................................................... 44 3 Lithium Diffusion..................................................................................................................... 66 3.1 Introduction....................................................................................................................... 66 3.2 Experimental Methods ...................................................................................................... 68 3.3 Analytical Techniques ...................................................................................................... 69 3.3.1

Major Element Analyses....................................................................................... 69

3.3.2

Lithium Analyses .................................................................................................. 70 3.3.2.1 LA-ICPMS ............................................................................................. 70 3.3.2.2 Secondary Ion Mass Spectrometry (SIMS) ............................................ 70

3.3.3

Data Reduction...................................................................................................... 71

3.4 Results............................................................................................................................... 72 3.4.1

Diffusion in Clinopyroxene .................................................................................. 72

3.4.2

fO2 Series Experiments ......................................................................................... 72

3.4.3

Diffusion in Olivine .............................................................................................. 73

3.4.4

Kinetic Fractionation of 7Li/6Li ............................................................................ 73

3.5 Discussion ......................................................................................................................... 74 3.5.1 Effect of fO2 on lithium diffusion in clinopyroxene ............................................. 74 3.5.2

Comparison with other lithium diffusion studies.................................................. 76

3.5.3

Comparison with diffusion of other cations in clinopyroxene.............................. 76

3.5.4

Geological Implications ........................................................................................ 77 3.5.4.1 Preservation of Lithium Signatures ........................................................ 77 3.5.4.2 Closure Temperature .............................................................................. 79 3.5.4.3 The Potential for Re-Equilibration of Lithium Composition ................. 79 3.5.4.4 Diffusion-Induced Isotopic Fractionation .............................................. 82

3.6 Conclusions....................................................................................................................... 84 4 Technique Development to Study Muscovite-Fluid Partitioning of Nitrogen....................... 104

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4.1 Introduction..................................................................................................................... 104 4.2 Theoretical Considerations ............................................................................................. 106 4.2.1 N-speciation and Isotopic Fractionation ............................................................. 106 4.2.2

Buffering pH ....................................................................................................... 108

4.3 Experimental Methodology ............................................................................................ 109 4.4 Analytical Methods......................................................................................................... 109 4.5 Results............................................................................................................................. 110 4.5.1

Nitrogen Partitioning .......................................................................................... 110

4.5.2

Nitrogen Isotopic Fractionation .......................................................................... 111

4.6 Discussion ....................................................................................................................... 112 4.6.1 Utility of NH4Cl as Nitrogen Source .................................................................. 112 4.6.2

Analytical Considerations................................................................................... 112

4.6.3

Experimental Considerations .............................................................................. 113

4.6.4

Isotopic Fractionation Experiments and Atmospheric Contamination ............... 114

4.7 Suggestions for Future Work .......................................................................................... 114 5 Summary of Results and Conclusions.................................................................................... 125 References................................................................................................................................... 128 Appendix 1.................................................................................................................................. 140 6 Summary of Boron Work....................................................................................................... 140

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List of Tables

Table 2.1 Composition of Starting Material ................................................................................. 46 Table 2.2 Experimental Conditions .............................................................................................. 47 Table 2.3 Standards and Reference Material ................................................................................ 48 Table 2.4 Run Product Composition............................................................................................. 49 Table 2.5 Run Product Lithium Concentration............................................................................. 50 Table 2.6 Isotopic Composition of Starting Materials and Run Products .................................... 51 Table 3.1 Composition of Starting Material ................................................................................. 85 Table 3.2 Measurements of Standards .......................................................................................... 86 Table 3.3 Summary of Experiments ............................................................................................. 87 Table 4.1 Experiments and Results............................................................................................. 116 Table 4.2 Percentage of Nitrogen Contribution from Air........................................................... 116

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List of Figures Figure 1.1 Sources and Concentration of Lithium in the Earth .................................................... 11 Figure 1.2 Li/Y ratio and 7Li in Arc Lavas ................................................................................. 12 Figure 1.3 Lithium Diffusion Coefficients ................................................................................... 13 Figure 2.1 Internal Reference Materials ....................................................................................... 52 Figure 2.2 Standards and Reference Material............................................................................... 53 Figure 2.3 Photomicrographs of Starting Material and Run Products.......................................... 54 Figure 2.4 Time Resolved Spectra................................................................................................ 55 Figure 2.5 lnDLi cpx/fluid vs 1000/T ........................................................................................... 56 Figure 2.6 lnDLi ol/fluid vs 1000/T.............................................................................................. 57 Figure 2.7 Anorthite/Fluid Lithium Partitioning .......................................................................... 58 Figure 2.8 Olivine/Clinopyroxene Lithium Partitioning .............................................................. 59 Figure 2.9 Lithium Partitioning From Mantle Xenoliths and Experimental Studies.................... 60 Figure 2.10 Mineral/Fluid Isotopic Fractionation of Lithium ...................................................... 61 Figure 2.11 Time for Li and B Transport to Top of Column........................................................ 62 Figure 2.12 Evolution of the Slab Derived Fluid by due to Rayleigh Distillation ....................... 63 Figure 2.13 Lithium Coordination and P-T Paths......................................................................... 64 Figure 2.14 Evolution of 7Li of Mantle Wedge due to Hydrofractures ...................................... 65 Figure 3.1 Li Elemental and Isotopic Gradients in San Carlos Opx............................................. 88 Figure 3.2 Effect of fO2 Anneal .................................................................................................... 89

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Figure 3.3 Zero time Experiment.................................................................................................. 90 Figure 3.4 Results for Experiment Kcpx-12 ................................................................................. 91 Figure 3.5 X-Ray Maps of Run Product ....................................................................................... 92 Figure 3.6 Time Series.................................................................................................................. 93 Figure 3.7 Measured Lithium Diffusion Coefficients................................................................... 94 Figure 3.8 fO2 Experiment Series ................................................................................................. 95 Figure 3.9 Lithium Diffusion Profile in San Carlos Olivine ........................................................ 96 Figure 3.10 Lithium Diffusion and Isotopic Fractionation in Kcpx-2.......................................... 97 Figure 3.11 Comparison of Lithium Diffusion Coefficients ........................................................ 98 Figure 3.12 Comparison of Diffusivities Measured in Pyroxene ................................................. 99 Figure 3.13 Retention of Lithium Composition.......................................................................... 100 Figure 3.14 Comparison of Closure Temperature of Li and Sr in Clinopyroxene ..................... 101 Figure 3.15 Lithium Isotopic Compositions of Kilauea Iki Lava Lake Rocks ........................... 102 Figure 3.16 Li Isotopic Gradient in San Carlos Opx and Modeled Profile ................................ 103 Figure 4.1 Summary of N Concentration and Isotopic Composition ......................................... 117 Figure 4.2 Calculated N2-, and NH3-NH4+ Fractionation ........................................................... 118 Figure 4.3 Relationship of fH2, fN2, and fNH3 ............................................................................ 119 Figure 4.4 Scanning Electron Micrograph of Muscovite Texture .............................................. 120 Figure 4.5 Nitrogen contents of run products ............................................................................. 121 Figure 4.6 Nitrogen isotopic compositions of run products ....................................................... 122

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Figure 4.7 Isotopic shifts of run products ................................................................................... 123 Figure 4.8 Puncturing Device ..................................................................................................... 124

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List of Appendices 6.1 11B notation ................................................................................................................... 140 6.2 Evidence of Boron Mobility from Arc Lavas ................................................................. 140 6.3 Evidence of Boron Mobility from Eclogites................................................................... 142 6.4 Summary of Experimental Methodology........................................................................ 143 6.5 Details of Boron Study.................................................................................................... 143 6.6 Boron Analyses............................................................................................................... 146 6.7 References for Boron Study............................................................................................ 148

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Introduction

According to Davidson (1996), three fundamental questions remain unanswered in the study of island arc magmagenesis: “1. What is the (presubduction) composition of the mantle wedge source of arc magmas? 2. To what extent does it melt, and by what process? 3. What is the composition and amount of slab-derived component added to the wedge?” Despite the significant strides that have been made in both our knowledge of earth processes and our technical ability to analyze earth materials with greater precision and accuracy, these questions remain unsatisfactorily answered today. Low abundance, or trace elements, can provide essential information to address these issues. For example, both convergent margin basalts and mid ocean ridge basalts (MORB) have similar major element compositions; however, convergent margin basalts are differentiated by high LILE/REE; (large ion lithophile element - Rb, K, Cs, Ba, Sr to rare earth element - actinides; La through Lu) and high LILE/HFSE (high field strength element – Nb, Ta, Zr, Hf, Ti) signature (Davidson, 1996). This pattern is interpreted to mean that arc magmas are products of the overlying mantle wedge melt plus a LILE-rich fluid or melt originating from the subducted slab. Boron, lithium and nitrogen have often been employed to identify the composition and amount of the slab component in island arc magmas. These elements are relatively fluid mobile and somewhat incompatible in mantle minerals and are easily released from the slab and concentrated into the magmas at the arc front (Ishikawa and Tera, 1999; Leeman, 1996; Ishikawa and Nakamura, 1994; Ishikawa et al., 2001; Moriguti and Nakamura, 1998).

1.1 Chemical Properties of Lithium Lithium belongs to the Group 1 elements of the periodic table. Like the other alkali metals it is characterized by low ionization energy and low electronegativity, and commonly forms hydroxides, nitrides, carbonates and chlorides. When octahedrally coordinated it has an effective ionic radius of 0.59 Å which is comparable to octahedrally coordinated Mg2+ (0.72 Å) and Fe2+

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(0.78 Å), which allows it to substitute for these elements in olivine, pyroxenes, amphiboles and clays (Brenan et al., 1998; Wenger and Armbruster, 1991). The oxygen co-ordination of lithium can vary from 3 to 8; although lithium has a preference for tetrahedral co-ordination in melts and fluids (Cormier et al., 1998; Majérus et al., 2003), it can be accommodated by octahedral coordination, as is the case in many silicate minerals (Wenger and Armbruster, 1991). Lithium has a single valence electron with a very low ionization potential, which makes it easily solvated. Lithium has two stable isotopes, 6Li and 7Li, with a relative mass difference of ~16 % and abundances of ~7.52 % and ~92.48 %, respectively. Enrichments in lithium isotopes are described as either:

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  6 Li/ 7 Li smp     1  1000  6 Li    6 7   Li/ Li std     

or

  7 Li/ 6 Li smp     1  1000  7 Li    7 6   Li/ Li std     

where ‘smp’ refers to the sample and ‘std’ refers to the standard, typically the NBS lithium carbonate L-SVEC (SRM #8545). The δ7Li notation is recommended by the International Union for Pure and Applied Chemistry (Coplen et al., 1996) and will be used here. Figure 1.1 shows the lithium abundance and isotopic composition of various geochemical reservoirs. As with other isotopes, lithium isotopic fractionation between minerals and fluids depends on the difference in the zero point potential energy (ZPE) between the phases of interest. Heavier isotopes have lower vibrational frequencies, and therefore a lower ZPE than lighter isotopes (Chacko et al., 2001). The molecule or phase that will undergo the greatest reduction in ZPE with the substitution of the heavy isotope will become enriched in the heavier isotope (Chacko et al., 2001). Ab initio calculations have demonstrated that during mineral-solution reactions 6Li should be preferentially incorporated into octahedrally coordinated sites in the solid (Yamaji et al., 2001). This appears to be the case in the formation of secondary minerals produced during alteration of crustal rocks (Huh et al., 2001; Pistiner and Henderson, 2003; Seyfried et al., 1998). During the formation of clays, 6Li is concentrated in the solid phase while the resulting fluid becomes enriched in 7Li (Huh et al., 2001; Pistiner and Henderson, 2003). Experimental measurements of lithium isotopic fractionation between spodumene and fluids also confirm this behavior (Wunder et al., 2006). Interestingly, experiments measuring lithium isotopic fraction between staurolite and fluids found that 6Li was preferentially enriched in the

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fluids, and 7Li was enriched in the solids (Wunder et al., 2007). Considering that lithium is in tetrahedral coordination in staurolite, this result also appears to confirm the ab initio calculations.

1.2 Sources and Concentrations of Lithium in the Earth The various components of the convergent margin system have significant differences in lithium concentrations and lithium isotopic compositions (see Figure 1.1). Inputs such as continental crust (as pelagic clays) and altered oceanic crust tend to be enriched in lithium with respect to the mantle, but the isotopic composition can vary considerably depending on the type and degree of alteration or weathering. Mantle inputs tend to be more uniform in terms of lithium concentration and isotopic composition; however, the MORB-source mantle can potentially contain both elemental and isotopic heterogeneities. Very low δ7Li (-11 ‰ to +5 ‰) values are found in samples of metamorphosed oceanic crust, possibly reflecting low temperature dehydration of the slab during subduction (Zack et al., 2003). Conversely some pyroxenites from the Zabargad peridotite, which is considered to be a fragment of exhumed mantle wedge, have high δ7Li values (+8.4 ‰ to +11.8 ‰) suggesting that 7Li-rich fluids or melts derived from subducted altered oceanic crust (AOC) may also be transferred to the mantle (Brooker et al., 2004). The variability exhibited by the components of the arc magmagenesis system is not always reflected in the output. Some island arc magmas display higher ratios of fluid mobile elements, such as lithium, to relatively immobile elements, such as yttrium in the front-arc regions, which decrease towards the back-arc. However, few arc lavas have δ7Li significantly greater than MORB, and correlations between δ7Li and fluid enrichment are not always clear or consistent (Chan et al., 2002; Tomascak et al., 2002; Tomascak, 2004; Leeman et al., 2004).

1.2.1

Subducted Materials (AOC and Seafloor Sediments)

Lithium in seafloor sediments ranges from 5 to 80 ppm with δ7Li ranging from -5 ‰ to +20 ‰ (Marschall et al., 2007 and references therein). In oceanic crust lithium can range from 5 to 6 ppm with a δ7Li of +1.5 ‰ to +5.6 ‰ in fresh mid-ocean ridge basalts (Moriguti and Nakamura, 1998; Tomascak et al., 2008; Chan et al., 1992) whereas altered oceanic crust may have >75 ppm lithium with δ7Li of +14.2 ‰ (Moriguti and Nakamura, 1998; Tomascak et al., 2008; Chan et al., 1992). Low temperature alteration of basalt results in concentration of lithium, and preferentially

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Li, into secondary minerals and enrichment of 7Li in seawater with a fractionation of up to +19

‰ with respect to the residual solid (Chan et al., 1992; Chan et al., 2002; Seyfried et al., 1998; James et al., 2003). At temperatures greater than 350 oC, lithium is mobilized by saline fluids and the extent of the isotopic fractionation decreases (Chan et al., 1992; Chan et al., 2002; Seyfried et al., 1998; James et al., 2003). The extent of solid-fluid fractionation has been inferred by various authors by measuring lithium in sediment-derived pore fluids and serpentine diapirs (Chan et al., 1992; Chan et al., 2002; Chan and Kastner, 2000; Benton et al., 2004). Pore fluids from the Costa Rican trench show a ~11 ‰ enrichment of 7Li with respect to the down-going sediments (Chan and Kastner, 2000). A larger range of isotopic compositions, 7Li of -0.5 ‰ to +10 ‰ has been measured in so called serpentinite diapers, which are super-hydrated mantle wedge extruded in the fore arc of the Mariana trench (Benton et al., 2001; Benton et al., 2004). The variability in the lithium isotopic composition of vent fluids (δ7Li ranging from +5 ‰ to +43.8 ‰) likely reflects differences in temperature, reaction paths and fluid - rock ratios as well as source rock composition (Zhang et al., 1998; Chan and Kastner, 2000; Foustoukos et al., 2004).

1.2.2

Eclogites

Alpine eclogites are thought to be exhumed remnants of subducted oceanic crust. The eclogites at Trescolmen, Switzerland investigated by Zack et al. (2003) displayed extremely light 7Li values ranging from -11 ‰ to +5 ‰. This study speculated that during subduction the slab was progressively depleted in 7Li, via Rayleigh distillation during dehydration, and that the resulting 6

Li enriched material was recycled into the mantle. A more comprehensive study of eclogites,

blueschists and other high pressure metamorphic rocks from classic European (Swiss Alps, Münchberg, Aldalen and Greek Islands) and Asian localities (Qaidam, Dabieshan and Tianshan) by Marschall et al. (2007) found an even larger range of lithium concentrations and isotopic compositions. Lithium concentrations ranged from ~1 to ~50 ppm and 7Li values ranged from 21.9 ‰ to > +6 ‰ (Marschall et al., 2007). More significantly, no correlation was found between lithium content and 7Li within any single locality or within the full population. In fact, many of the lightest samples had > 30 ppm lithium, contrary to what would be expected from a simple Rayleigh-type distillation of the subducted slab. Marschall et al. (2007) concluded that many exhumed eclogites have lithium compositions (both elemental and isotopic) that have been

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influenced by influx of retrograde fluids and kinetically induced isotopic fractionation during exhumation, and therefore primary subduction-induced fractionation has been modified by more recent processes.

1.2.3

Upper Mantle

From analyses of pristine peridotite xenoliths, the upper mantle is estimated to contain 1.5 ppm lithium with an average δ7Li of +4 ‰ (Jagoutz et al., 1979; Tomascak, 2004; Jeffcoate et al., 2007). However, the fact that OIB and MORB can sometimes display a range of values has led to the suggestion that variable amounts of recycled crustal material are sometimes present in the mantle sources of these lavas (Tomascak et al., 2008). A study of MORB lavas from different ridge systems found δ7Li ranging from +1.6 ‰ to +5.6 ‰. This represents a 5 ‰ heterogeneity in the samples with no consistent correlation of δ7Li with major and trace elements or radiogenic isotopes (Tomascak et al., 2008). Analysis of lithium in glass inclusions from Hawaii showed δ7Li varying from –10.2 ‰ to +8.4 ‰ (Kobayashi et al., 2004). Similarly, low δ7Li (-3.3 ‰ to +1.2 ‰) from glass inclusions in Iblean (Sicilian) Plateau tholeiites are thought to reflect melting of an isotopically light region in the mantle (Gurenko and Schmincke, 2002). In rare cases, δ7Li correlations with abundances of other trace elements or isotopes can be found. For example, samples from the East Pacific Rise (EPR) show a weak correlation of increasing δ7Li with increasing Cl/K, which was interpreted to reflect mixing or assimilation of recycled crustal material (Tomascak et al., 2008). Other EPR samples show a positive correlation between δ7Li and 143Nd/144Nd, (Elliot, 2004), again reflecting a possible recycled component. That reservoirs with variable δ7Li exist in the mantle is also suggested by studies of peridotite massifs and ultramafic xenoliths. Nishio et al. (2004) report the lithium isotopic composition of clinopyroxene from xenoliths from Japan, SE Australia and eastern Russia. The δ7Li values from NE Japan and SE Australia were high (+4 to +7 ‰), whereas Russian and SW Japan samples were significantly lower (-17 to -3 ‰). In some of the samples from eastern Russia, δ7Li could be positively correlated with 143Nd/144Nd but those correlations did not apply to the other sample populations or even to all the eastern Russia samples. Ionov and Seitz (2008) also reported lithium concentrations and isotopic data from xenoliths from the Kamchatka arc and the Vitim (Siberia) volcanic field (an intra-plate continental volcanic setting). They found a relatively small range of lithium concentrations, ~1 to ~2 ppm, and lithium isotopic composition of -3.6 to +6 ‰.

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These studies demonstrate significant variations in the isotopic compositions of lithium in the upper mantle, but suggest that such variation is fairly localized. Reports of significant correlation between δ7Li and other elemental or isotopic tracer elements are rare within any given sample locality and no global correlation has been found within a given tectonic setting.

1.2.4

Convergent Margin Magmas

Magmatism at convergent margins is commonly believed to be due to hydrous fluids from the subducting slab fluxing the mantle wedge and lowering the mantle solidus. As mentioned previously, the resulting magmas are very often characterized by higher ratios of fluid mobile elements to relatively insoluble elements, such as the high field strength elements (HFSE). In order to distinguish between crystal/melt fractionation effects and fluid involvement, fluid mobile elements (such as lithium and boron) are measured against relatively insoluble elements with similar solid/melt partitioning (i.e. B/Be, Li/Yb or Li/Y). The Izu arc in Japan is the locality showing the clearest indication of a Li-bearing, slab-derived component. This is shown by the correlated decrease in both Li/Y and δ7Li from the arc front lavas to those erupted in the back arc region (Figure 1.2a, data from Moriguti and Nakamura, 1998). This is suggestive of continuing mobilization of lithium into the arc source region by fluids derived from dehydration reactions in the down going slab (Leeman, 1996; Ishikawa and Tera, 1999; Ishikawa and Nakamura, 1994; Ishikawa et al., 2001). However, the trend displayed at Izu appears to be the exception and not the rule (Figure 1.2b, data from Tomascak et al., 2002). Variations in slab age and angle of subduction, which would influence the thermal regime, and therefore the extent of dehydration, have been cited to explain these differences (Moriguti et al., 2004). Yet, a difference in the Li/Y and 7Li behavior between the Kurile arc and the Izu arc, or even between the Izu arc and the Japan arc where age of the subducted slab is similar, still persists, despite the similarity of subduction regime (Morguti et al., 2004; Tomascak et al., 2002). Another suggestion is that slab-derived fluids are significantly modified during transport through the mantle wedge to the melt source, and that the lithium signal is attenuated by interaction with mantle minerals (Tomascak et al., 2002). The extent of modification of the slab-derived component during transport through, and interaction with, the overlying mantle wedge is unknown, but it remains unanswered as to why more arc magmas do not display the same trends as clearly as the Izu arc.

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1.3 The scale of lithium heterogenities Recent in situ micron scale analyses of lithium isotopes in geological samples have yielded unexpected results. High temperature equilibrium is expected to impose minimal differences in the lithium isotopic composition of individual minerals; however, samples from some magmatic environments have revealed significant isotopic heterogeneity at the grain-scale. For example, Rudnick and Ionov (2007) reported highly variable δ7Li in clinopyroxene and olivine grains in peridotite xenoliths from eastern Russia. δ7Li values ranged from -0.8 to -14.6 ‰ for clinopyroxene and -1.7 to +11.9 ‰ for corresponding olivine, and olivine/clinopyroxene distribution coefficients varied from 0.2 to 1.0, which is lower than previously estimated for equilibrium partitioning. Analyses of olivine and clinopyroxene pairs from a xenolith from the Vitim volcanic field found δ7Li to range from -17 to -18 ‰ in the pyroxenes with a δ7Li of +6 ‰ in the corresponding olivine (Ionov and Seitz, 2008). Bulk measurements of olivine phenocrysts in primitive magmas from a variety of localities found a relatively uniform δ7Li of +3.2 to +4.9 ‰; however, measurements of clinopyroxene yielded highly variable δ7Li (+6.6 ‰ to -8.1 ‰; Jeffcoate et al., 2007). Both the olivine and clinopyroxene phenocrysts from Solomon Island lavas are zoned with respect to lithium and δ7Li (Parkinson et al., 2007). Rims of phenocrysts are enriched in lithium compared to the cores and the δ7Li decreases from core to rim by as much as 20 ‰ in a W-shaped profile (Parkinson et al., 2007). This pattern was also observed by Jeffcoate et al. (2007) who measured a 40‰ variation in a single orthropyroxene crystal from a San Carlos xenolith. The extreme grain-scale variability exhibited by lithium and lithium isotopes is not limited to terrestrial samples. The basaltic lunar meteorite NWA 479 examined by Barrat et al. (2005) contains olivine and pyroxene phenocrysts that also display a wide range of δ7Li values (+2.4 to +15.1 ‰ in olivine and -0.2 to 16.1 ‰ in pyroxene). Beck et al. (2004) examined pyroxenes in the shergottite meteorite NWA 480 and found extreme zoning of δ7Li from -17 ‰ in the cores to +10 ‰ in the rims and an absence of lithium compositional variation. The extreme fractionation of lithium isotopic values documented in these high temperature samples suggests kinetic rather than equilibrium processes (Lundstrom et al., 2005; Beck, 2006; Jeffcoate et al., 2007; Parkinson et al., 2007; Rudnick and Ionov, 2007; Marchall et al., 2007). This kinetic effect has been experimentally demonstrated by Richter et al. (2003) for diffusion of

8

lithium in molten silicate where fractionation occurs due to slightly faster transport of 6Li than 7

Li. The time scale for the development of diffusion-controlled isotopic fractionation is likely to

be quite short as documented by the rapid lithium exchange in natural samples. Berlo et al. (2004) reported rapid mobilization of lithium in a study of plagioclase phenocrysts from the 1980 eruption of Mount St. Helens in Washington, USA. Plagioclase phenocrysts erupted prior to the degassing event contained ~14 ppm lithium, whereas those erupted immediately after contained ~5 ppm. The implication is that the magma lost a significant amount of lithium in a seven-day period, which was recorded in the lithium content of the plagioclase phenocrysts. Kent et al. (2007) also found that the lithium contents of plagioclase phenocrysts from the Mount St. Helens 2004 dome lavas had increased due to the addition of a pre-eruptive lithium rich vapour phase. Based on the lithium contents of plagioclase phenocrysts, melt inclusions, and plagioclase encapsulated within gabbroic inclusions Kent et al. (2007) were able to estimate that the influx of the lithium-rich volatile phase occurred within ~1 yr of the dome lava eruptions.

1.4 Previous experimental work Previous experimental work has found lithium to be moderately incompatible in clinopyroxene co-existing with either fluid or melt, and that partitioning is a function of clinopyroxene major element composition, DLicpx/melt increasing with increasing Ca/Al ratio (Hart and Dunn, 1993; Brenan et al., 1998a; Blundy et al., 1998; Blundy and Dalton, 2000; Hill et al., 2000; Bennett et al., 2004). Lithium was also found to be moderately incompatible in olivine relative to melt (Brenan et al., 1998a; Brenan et al., 1998b, Taura et al., 1998; Zanetti et al., 2004). To date, lithium partitioning between olivine and fluids has not been measured. Coogan et al. (2005) measured lithium partitioning between plagioclase and clinopyroxene and found DLiplag/cpx increases with increasing temperature (900oC to 1200oC). In this case partitioning was not investigated with respect to plagioclase or clinopyroxene major element chemistry. Little work has been done to determine lithium isotope partitioning and diffusion at pressures and temperatures corresponding to crustal and mantle processes. A study examining isotopic fractionation between spodumene and hydrous fluids measured an enrichment of 7Li in the fluid from +3.5 ‰ at 500oC to ~ +1.0 ‰ at 900oC and 2.0 GPa (Wunder et al., 2006). Isotopic fractionation between fluids and mica (from 300oC to 500oC, 2.0 GPa) and staurolite (from

9

670oC and 880oC, 3.5 GPa) found fluids to be preferentially enriched in 7Li relative to the mica, and staurolite to be slightly enriched in 7Li relative to the fluid (Wunder et al., 2007). These studies are consistent with ab initio calculations where 6Li is preferentially fractionated in sites with octahedral coordinations (e.g. mica, spodumene; Yamaji et al., 2001, Wunder et al. 2006) and 7Li is preferentially fractionated in to sites with tetrahedral coordination (e.g. staurolite; Wunder et al., 2007). Fractionation of lithium isotopes was also measured in the quartzmuscovite-fluid system, from 400-500oC (Lynton et al., 2005). Lynton et al. (2005) found both the quartz and the mica to be preferentially enriched in 7Li, with fractionation factors ranging from +8 to +12 ‰ for quartz and +18 to +20 ‰ for mica. For reasons that are unclear, these results are inconsistent with the subsequent studies of Wunder et al. (2006, 2007) or the results of this study. Because mica has lithium in octahedral coordination the expectation is that the fluids would be preferentially enriched in 7Li with respect to the solid. Experimental studies of kinetic isotopic fractionation (Richter et al., 2003) found that 7Li could be fractionated from 6Li by tens of per mil during diffusion between molten basalt and rhyolite or when diffusing through fluids. Although estimates of lithium diffusion coefficients have been made from natural samples, to date there have been few laboratory measurements of lithium diffusion in rock forming minerals (Figure 1.3). Giletti and Shanahan (1997) measured the diffusion rates of various alkali elements in plagioclase feldspar. They found that diffusion in feldspars is a function of ionic radius and cation charge, and as a result of its small size, lithium diffusion rates are very rapid. Coogan et al. (2005) measured the diffusion coefficient for 6Li in clinopyroxene between 800oC and 1100oC by using SIMS analysis and found lithium to be similarly rapid.

1.5 Focus of Thesis and Distribution of Work Knowledge of lithium elemental partitioning and isotopic fractionation between fluids and common rock forming minerals is essential to evaluate the variations seen in natural samples adequately. Information on lithium diffusion in minerals can be used to more accurately assess the time-scales of magmatic and hydrothermal processes and account for intermineral isotopic differences.

10

This study examines the lithium, nitrogen and boron isotope fractionation that occurs in mineralfluid reactions during slab and mantle interaction. The existing experimental database is insufficient to properly evaluate the degree of isotopic fractionation that occurs during fluidmineral partitioning of lithium, boron, and nitrogen. The results of this work provide the essential input for modeling the behavior of lithium in the mantle. Chapter 2 of this study examines the partitioning and isotopic fractionation that occurs between clinopyroxene, olivine, plagioclase and aqueous fluids and the intermineral fractionation between olivine and clinopyroxene. All experiments were conducted by N. Caciagli in the High Pressure Laboratory at the University of Toronto. The elemental lithium was analyzed by laser ablation inductively coupled plasma mass spectroscopy (LA-ICPMS) at the University of Toronto by N. Caciagli. The multi collector inductively coupled plasma mass spectroscopy (MC-ICPMS) analyses of the clinopyroxene starting material and bulk analyses of the run product clinopyroxene were done at University of Maryland by W. F. McDonough. The lithium isotopic composition of starting material anorthite and olivine, and in situ lithium isotopic compositions were analyzed by secondary ionization mass spectroscopy (SIMS) at Lawrence Livermore National Laboratory by N. Caciagli with the assistance of D. Phinney. Chapter 3 of this study measures the diffusion coefficient of lithium in clinopyroxene and olivine. All experiments were conducted by N. Caciagli in the High Pressure Laboratory at the University of Toronto. The elemental lithium was analyzed by LA-ICPMS at the University of Toronto by N. Caciagli and lithium isotopic analyses were done by SIMS at Lawrence Livermore National Laboratory by N. Caciagli with the assistance of D. Phinney. Chapter 4 of this study outlines the technique development for experimental measurements of nitrogen partitioning and isotopic fractionation between fluids and muscovite. All experiments were conducted by N. Caciagli in the High Pressure Laboratory at the University of Toronto. The nitrogen elemental and isotopic analyses were done at Lehigh University with the assistance of G. Bebout. The exploratory work on techniques to measure boron partitioning and isotopic fractionation between muscovite and fluid is summarized in Appendix 1.

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est mean mantle

A

xenoliths eclogites IAB OIB MORB AOC marine seds est mean c.c. sea water ave river water 0

0.01

0.1

1 ppm

10

100

est mean mantle xenoliths eclogites IAB OIB MORB AOC marine seds est mean c.c. sea water river water -30

1000

B

-20

-10

0  Li

10

20

30

40

7

Figure 1.1 Sources and Concentration of Lithium in the Earth Lithium concentration (A) and isotopic composition (B) of various terrestrial reservoirs. River water: Huh et al. (1998); seawater: Millot et al. (2004); estimated continental crust (c.c): Teng et al. (2004); marine sediments: Bouman et al. (2004); AOC: Chan et al. (1992); MORB: Moriguti and Nakamura (1998); Tomascak et al. (2008); Nishio et al. (2002); OIB: Kobayashi et al. (2004); IAB: Moriguti and Nakamura (1998); Tomascak et al. (2002); eclogites: Marschall et al. (2007); xenoliths: Nishio et al. (2004); est mean mantle: Jagoutz et al., (1979) and Tomascak (2004).

12

8

A 7

6

7Li

5

4

3

2

1 0

0.2

0.4

0.6

0.8

1

1.2

1.4

Li/Y

8

B 7 6

7

 Li

5 4 3 2 1 0 0

0.2

0.4

0.6

0.8

1

1.2

1.4

Li/Y

Figure 1.2 Li/Y ratio and 7Li in Arc Lavas A plot of 7Li as a function of Li/Y ratio in (a) Izu arc basalts and (b) basalts from other Sunda and Aleutian arcs. Lavas from the Izu arc display a trend of increasing 7Li with increasing Li/Y ratio, and show an inverse relationship with depth to the slab (Benioff zone) suggestive of progressively decreasing amounts of fluid being mobilized during subduction. The trend of increasing 7Li with increasing Li/Y ratio is not consistently observed in other island arcs. Izu arc data from Moriguti and Nakamura (1998), Sunda and Aleutian arc data from Tomascak et al. (2002).

13

o

T ( C) 600 400

1400 1000 800

-6.0

200

-8.0 Si-crystal -10.0

2

D (m /s)

-12.0

Li

-14.0 cpx, Coogan et al. 2005 -16.0 -18.0 albite & anorthite

-20.0 -22.0 4

8

12

16

20

10,000/T (K)

Figure 1.3 Lithium Diffusion Coefficients Plot of log DLi vs. 10,000/T (K) for lithium diffusion in geologically significant minerals. Lithium diffusion in a p-type Sicrystal data are from Pell (1960), feldspar data are from Giletti and Shanahan (1997), and cpx data are from Coogan et al. (2005)

14

2

Lithium Partitioning and Isotopic Fractionation

2.1 Introduction Lithium and lithium isotopes are increasingly used as tracers of surface inputs to the mantle during subduction. With a strong affinity for fluids, an incompatible nature during mantle melting and a high relative mass difference (~16 %) between the two stable isotopes, (6Li and 7

Li), lithium has the potential to serve as a robust indicator of fluid-rock interaction in a variety

of geological settings. Enrichments in lithium isotopes are described as either:

(2)

  6 Li/ 7 Li smp     1  1000  Li    6 7   Li/ Li std      6

or

  7 Li/ 6 Li smp     1  1000  Li    7 6   Li/ Li std      7

where ‘smp’ refers to the sample and ‘std’ refers to the standard, typically the NIST lithium carbonate L-SVEC. The δ7Li notation is recommended by the International Union for Pure and Applied Chemistry (Coplen et al., 1996) and will be used here. Several reservoirs of lithium, which are isotopically distinct from the mantle and each other, are present within Earth. Seawater has 0.18 ppm lithium with δ7Li of +32 ‰ (James and Palmer, 2000) and the continental crust contains an average of 35 + 11 ppm lithium with a δ7Li that ranges from –5 to +5 ‰ (Teng et al., 2004). Lithium in fresh mid-ocean ridge basalts (MORB) can range from 5 to 6 ppm, with a δ7Li of +1.5 ‰ to +5.6 ‰ (Moriguti and Nakamura, 1998; Tomascak et al., 2008; Chan et al., 1992). Altered oceanic crust (AOC) has a significantly greater concentration of lithium, >75 ppm lithium, and is isotopically heavier than pristine MORB with a δ7Li of up to +14.2 ‰ in the most altered oceanic crust (Moriguti and Nakamura, 1998; Tomascak et al., 2008; Chan et al., 1992). The mantle is estimated to contain 1.6 ppm lithium with an average δ7Li of +4 ‰ (Jagoutz et al., 1979; Moriguti and Nakamura, 1998; Tomascak, 2004; Tomascak et al., 2008; Teng et al., 2004). However, studies of mantle xenoliths suggest that reservoirs with variable δ7Li exist in the mantle (Seitz et al., 2004; Nishio et al., 2004; Brooker at al., 2004; Lundstrom et al., 2005). Very low δ7Li (-11 ‰ to +5 ‰) values are found in orogenic eclogites which are thought to reflect low temperature dehydration of the slab

15

during subduction (Zack et al., 2003). Conversely some pyroxenites from the Zabargad peridotite, which is considered to be a fragment of an exhumed mantle wedge, have high δ7Li values (+8.4 ‰ to +11.8 ‰) suggesting that 7Li-rich fluids or melts derived from subducted AOC are also transferred to the mantle (Brooker et al., 2004). In an investigation of lithium isotopes from the Kilauea Iki lava lake, Tomascak et al. (1999) demonstrated that neither partial melting nor low pressure differentiation results in significant (> +2 ‰) variations in δ7Li. This has led to the interpretation that the variability evident in some mantle-derived lavas is due to melting of a heterogeneous source. The Izu arc shows a trend with the greatest lithium and 7Li enrichment occurring at the arc front where δ7Li in lavas ranges from +7.6 ‰ to +1.1 ‰ (see Figure 1.2a) suggesting enrichment of the arc melt source by fluids derived from the down going slab (Moriguti and Nakamura, 1998). A study of lithium in glass inclusions from Hawaii showed the lithium isotopic composition to vary from –10.2 ‰ to +8.4 ‰ (Kobayashi et al., 2004). Similarly, low δ7Li (-3.3 ‰ to +1.2 ‰) has been measured in glass inclusions from the Iblean (Sicilian) Plateau tholeiites (Gurenko and Schmincke, 2002). A study of δ7Li values in OIBs from Antarctica did not show any significant deviations from MORB values (Ryan and Kyle, 2004); however, a study of MORB lavas from different ridge systems found a 5 ‰ heterogeneity in the samples, and no significant correlation of δ7Li with major elements, trace elements or radiogenic isotopes (a slight apparent correlation with Cl/K was observed; Tomascak et al., 2008). The fact that both OIB and MORB can sometimes display a range of values has prompted many researchers to suggest that variable amounts of recycled material with modified δ7Li is transported into the mantle sources of these lavas (Tomascak et al., 2008). But the origin and scale of these mantle heterogeneities are not well defined. The extent of the modification of the down going slab by fluid-mineral exchange during subduction remains ambiguous, as is the extent of modification of the slab derived fluid during transport though the mantle to the arc source. Previous experimental work has found lithium to be moderately incompatible in clinopyroxene relative to fluid or melt with DLicpx/fluid increasing with increasing Ca/Al ratio (Hart and Dunn, 1993; Brenan et al., 1998a; Blundy et al., 1998; Blundy and Dalton, 2000; Hill et al., 2000; Bennett et al., 2004). Lithium was also found to be moderately incompatible in olivine relative to melt (Brenan et al., 1998a; Brenan et al., 1998b; Taura et al., 1998; Zanetti et al., 2004) and fluid (Blundy and Dalton, 2000). Coogan et al. (2005) measured lithium partitioning between plagioclase and clinopyroxene and found DLiplag/cpx

16

increasing with increasing temperature (900oC to 1200oC). In this case, partitioning was not correlated with plagioclase or clinopyroxene major element composition. Little experimental work has been done to examine lithium isotope fractionation at pressures and temperatures relevant to crustal and mantle processes. Comparisons of δ7Li from metasomatized and pristine peridotite xenoliths suggest that some olivine – clinopyroxene fractionation, (up to 3.5 ‰ enrichment in 7Li), may occur at mantle temperatures (950oC; Seitz et al., 2004). Fluidmineral partitioning will also fractionate lithium isotopes as documented in a single study which measured an enrichment of 7Li in fluids relative to a Li-pyroxene (spodumene) by +3.5 ‰ at 500 o

C to ~ +1.0 ‰ at 900oC and 2.0 GPa (Wunder et al., 2006).

To date, a systematic investigation of the degree of isotopic fractionation and the extent of partitioning that occurs during mantle processes has been lacking. Both the isotopic fractionation, Li, and the lithium partitioning between major mantle phases need to be known to determine the extent to which a slab signal can propagate to the IAB source. This study presents lithium partitioning and isotopic fractionation measurements between fluids and common rock forming minerals. With this information, more accurate models can be developed to constrain the origins of lithium anomalies in the mantle.

2.2 Methods This study measured the partitioning of lithium between aqueous fluids and clinopyroxene, olivine and plagioclase at pressures and temperatures corresponding to lower crustal and upper mantle conditions (800oC to 1200oC; 1 GPa). Additional experiments were done to measure olivine – clinopyroxene isotopic fractionation at similar conditions. Starting materials were natural single crystals of: olivine (Fo82) from San Carlos, Arizona; plagioclase (bytownite) from Crystal Bay, Minnesota; clinopyroxene (diopside) from Dekalb, New York; and plagioclase (albite) from Mont St. Hilare, Quebec. Table 2.1 gives the composition of the starting materials. In all cases, the mineral samples were first crushed to 1-3 mm grain size, after which grains free of inclusions and alteration were hand picked and cleaned in dilute HNO3 and rinsed with ultrapure water in an ultrasonic cleaner. For the olivine and plagioclase experiments the minerals were ground to a fine powder under ethanol with an additional SiO2 + Al2O3 (1:1) mixture (~3

17

wt. % of total) added to approximate natural mantle fluid compositions (Brenan et al., 1998b; Holloway, 1971). Experiments containing olivine were not buffered for Fe loss to the noble metal capsule; therefore, run product compositions are shifted to more magnesium rich compositions (from Fo82 to Fo97 – 99). Experiments with bytownite as a starting mineral composition were not buffered for Na2O loss to the fluid and as a result run product compositions are shifted from bytownite (An75) to anorthite (An98-99). A few plagioclase experiments contained additional albite to stabilize more sodic compositions of plagioclase. With one exception all experiments with Dekalb diopside as starting material had ~3 wt. % SiO2 added (no Al2O3) since clinopyroxene dissolution buffers the aluminum content of the fluid. To minimize compositional zoning, a large fluid to solid ratio (4:1 by mass) was utilized for all experiments. One clinopyroxene experiment was carried out with the addition of 3 wt. % albite (+3 wt. % SiO2) to encourage compositional zoning with respect to the aluminum content of the clinopyroxene. A series of mineral pair experiments were run with clinopyroxene and olivine to constrain their inter-mineral isotopic fractionation. These experiments used the same starting materials prepared as above and mixed 80:20 cpx-olivine by mass. Isotopically labeled solutions were made from ultra pure water with lithium added as either Li2CO3 (LSVEC, SRM#8545) or some combination of LSVEC and a 6Li spike. For each experiment a sample of either clinopyroxene, olivine or plagioclase powder (+/- SiO2, Al2O3 or albite) and lithium bearing solution were added to a large volume Ni capsule with a Pt insert (Ayers et al., 1992). To promote crystal dissolution and re-precipitation, the bottom of the capsule was centered in the hotspot of the furnace. The temperature gradient over the length of the capsule is less than 10oC (Ayers et al., 1992). The experiments were conducted in an end-loaded piston–cylinder apparatus (Boyd and England, 1960) using a 1.9 cm bore pressure vessel, employing a cylindrical graphite heater and pressure cells consisting of crushable MgO, Pyrex and NaCl. Samples were initially cold pressurized to ~0.5 GPa and then heated to 300oC to generate sufficient internal pressure to prevent capsule deformation (Brenan et al. 1995). Temperature and pressure were then increased simultaneously with the maximum pressure being achieved by the time the sample reached 600oC. Temperature was monitored with W26% Re-W5% Re thermocouples uncorrected for the effect of pressure on

18

EMF. Experiments were run for 72 to 144 hours and quenched by cutting power to the sample heater which resulted in temperatures dropping to < 300oC in 20 seconds. The capsules were then recovered, punctured and dried. Fluid masses determined by weight loss were usually > 70 % of the initial fluid mass; however, during the course of the experiment the capsule material became work hardened; consequently, an undetermined amount of capsule material was sheared off during puncturing. As a result, the fluid masses used in the mass balance calculations are the initial fluid masses. In the case where a watertight seal was not maintained throughout an experiment, a drop in pressure and a collapsed capsule would result. One additional experiment was carried out in a cold seal vessel at 0.2 GPa and 800oC. In this case the sample powder and lithium solution were loaded into a 5 mm O.D. Au capsule. The capsule was then crimped, weighed, sealed by arc welding and reweighed to check for fluid loss. The sample was loaded into a vertically mounted pressure vessel, first pressured to 0.2 GPa and then externally heated. Temperatures were monitored with an internal type K thermocouple. Experiments were quenched by removing the furnace and cooling the pressure vessel with compressed air, which resulted in temperatures dropping to < 300oC in ~3 minutes. Table 2.2 provides a summary of experimental conditions.

2.3 Analytical Techniques 2.3.1

Major Element Analyses

Samples of starting material and splits of run products were mounted in epoxy, ground, polished to 0.3 m and carbon coated. The major element compositions of the starting materials and run product clinopyroxene, olivine and plagioclase were then obtained using the University of Toronto’s Cameca SX50 Electron Probe X-ray Microanalyzer (EPMA). An accelerating voltage of 15 kV and a focused 20 nA beam was used for all samples. The standards were albite for Na, anorthite for Al, diopside for Ca, Mg, Si, basalt for Fe, and bustamite, (Mn,Ca)3Si3O9, for Mn. X-ray intensities were converted to concentrations using modified ZAF or Phi-Rho-Z correction schemes. The reported errors are the 1σ variations of (n) analyses.

19

2.3.2

Lithium Analyses

2.3.2.1 MC-ICPMS Bulk lithium isotopic composition and element concentrations were determined for run product clinopyroxene using a Multi-Collector Inductively Coupled Plasma Mass Spectrometry (MCICPMS) at the University of Maryland. Samples were first rinsed in ultra-pure water to remove any water-soluble Li-bearing residue and then analyzed following the method described in Teng et al. (2004). The samples, 4-10 mg of run-product clinopyroxene, were digested in a mixture of HF and HNO3 and dissolved in a 4M HCl solution. The lithium from the samples was then separated from the dissolved matrix by thrice processing the solutions in cation exchange columns. The solutions were then introduced to the Nu-Plasma MC-ICP-MS in a 2% HNO3 solution, and isotopic compositions were obtained by measurement of 7Li and 6Li simultaneously on two high and low mass Faraday cups. Each sample analysis was bracketed by measurement of the L-SVEC standard. The isotopic values are reported as δ7Li (equation 2.1) in which the lithium isotopic standard is NIST L-SVEC Li2CO3. The 2σ precision of each analysis is ±1 ‰. Table 2.3 lists measurements of standards and reference materials.

2.3.2.2 SIMS In situ analyses of the lithium abundance and isotopic composition of the starting materials, run product clinopyroxene, olivine, and plagioclase were obtained using the Cameca IMS 3f ion microprobe at Lawrence Livermore National Laboratory. Secondary ions were generated by bombardment with a 5-12 nA negatively charged 16O primary beam, accelerated through –12.5 kV and focused to ~20 μm. The positive secondary ions were accelerated through 4.5 kV. 6Li and 7Li were measured with a mass resolving power of 1011, and no energy offset was applied. The background, mass 5.8, 6Li and 7Li were counted on an electron multiplier for 2 s, 10 s and 2 s respectively over 120-400 counting cycles, depending on count rate. Figure 2.1 shows the ‘uncorrected’ δ7Li values measured by SIMS of the internal reference materials; Dekalb diopside (measured in this study by MCICP-MS), San Carlos olivine (Magna et al., 2006), and clinopyroxene from experiment NCDL6 (analysed by MC-ICPMS), plotted against the δ7Li values measured by MC-ICPMS. The δ7Li values measured by MC-ICPMS of the internal reference materials range from -3 ‰ to +10 ‰ whereas the corresponding ‘uncorrected’ δ7Li values measured by SIMS range from +15 ‰ to +30 ‰. The discrepancy between values is due

20

to mass fractionations that are the result of both instrumental parameters and matrix effects (Decitre et al., 2002, Tomascak, 2004). It is important to note that all the values plot on a single line with a slope of ~1, suggesting that any matrix effect on the lithium instrumental isotopic fractionation is still within the 2σ precision of the analysis (± 4‰). The 7Li/6Li ratios can be corrected for this fractionation using the instrumental correction factor, ∆i (Decitre et al., 2002); (3)

∆i = δ7LiSIMS – δ7LiMC-ICPMS

Because all of the fractionation measured is relative to the same starting material, the 7Li/6Li ratio of Dekalb diopside was used as the internal reference material to determine ∆i. As a check on the calibration of this correction factor, ∆i, was determined using San Carlos olivine, ∆i = -22 ‰, and was found to be within 2σ (± 4 ‰) error of that for the Dekalb diopside, ∆i = -20 ‰. However, the concentration of lithium in Dekalb diopside and the NCDL6 experiment are both greater than that of San Carlos olivine (i.e. 8.6 ppm and 67 ppm vs. 2.5 ppm), therefore the Dekalb diopside and NCDL6 were primarily used as internal reference materials.

2.3.2.3 LA-ICPMS In situ analyses of lithium abundances in clinopyroxene, olivine, and plagioclase were also determined by laser ablation inductively coupled plasma mass spectrometry (LA–ICP–MS) at the University of Toronto. The system employs a frequency quintupled Nd:YAG laser operating at 213 nm, coupled to a VG PQExcell quadrupole ICP-MS. The laser was operated at 10 Hz, with He flushing the ablation cell to enhance sensitivity (Eggins et al. 1998), and produced spot sizes ~50 μm in diameter and ~ 50 μm deep. The torch position and lens settings were adjusted prior to each analytical session to optimize the signal intensity while ablating NIST 610 with a spot size of approximately 75 μm and a laser beam energy of less than 3 mJ, so that the sensitivity to 7Li was maximized. Data were collected as time-resolved spectra with background levels determined by counting for 20 s prior to the 60 s of sampling by laser ablation. Analyses were collected in blocks of >20, with the first and last two spectra acquired on standard reference materials. Count rates were collected and exported as CSV (comma-delimited values) files by “ThermoElectron Plasmalab” (TRA) software. All subsequent data reduction was performed off-line using the GLITTER 5.3 software package, supplied by Macquarie Research, Ltd. Ablation yields were corrected by referencing to the

21

known concentration of 43Ca or 55Mn as determined previously by electron microprobe analyses. Lithium concentrations in clinopyroxene, olivine and anorthite were quantified using the “in house” standard Kunlun diopside, which contains 42.6 ppm lithium. This standard was used routinely because it generated a lower lithium background over the course of the analysis compared to that produced by NIST 610. Kunlun diopside was, in turn, characterized using NIST 610 silicate glass, which contains 470.5 ppm lithium (Pearce et al., 1997). Table 2.3 compares the accepted values for several standard and reference materials (fused into glass in sealed Pt capsules at 1 GPa and 1200oC) which were also measured by LA-ICPMS using NIST 610 as a standard. The precision for concentration measurements is better than ± 10 %. Figure 2.2 shows lithium abundance of the cross-referenced samples plotted against lithium abundance determined by LA-ICP-MS.

2.4 Results 2.4.1

Major Element Chemistry

Figure 2.3 shows the size and morphology of run product clinopyroxene, olivine and plagioclase. The run products display considerable coarsening compared to starting materials, and welldeveloped crystal faces. Crystals grew as aggregates of grains nucleating on the lid and/or upper walls of the capsule. Table 2.4 lists the major element composition of the clinopyroxene, olivine and plagioclase run products produced in this study. The minimum detection limits are less than the lowest value cited for each element and the number in parentheses refers to 1σ of the standard deviation for n analyses and reflects the degree of sample heterogeneity.

2.4.2

Mineral – Fluid Lithium Partitioning

Table 2.5 lists the lithium contents of experimentally produced crystals as analyzed by LAICPMS as well as the calculated distribution coefficient (DLi) values for mineral/fluid partitioning or where applicable the D-values for mineral/melt partitioning. Mineral/fluid distribution coefficients (Dmin/fluid) were determined from final mass balance calculations: (4)

CLitotal = Cinitialmin Xinitialmin + Cinitialfluid Xinitialfluid

(5)

CLitotal = Cfinalmin Xfinalmin + Cfinalfluid Xfinalfluid

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Where Cmin is the lithium concentration of the mineral in ppm, Cfluid is the lithium concentration of the fluid in ppm and Xmin is the mass fraction of the mineral, and Xfluid is the mass fraction of the fluid. The initial concentration of fluid, Cinitialfluid, and the initial mass fraction of fluid, Xinitialfluid, is assumed to be equal to the final concentration and mass fraction, Cfinalfluid and Xfinalfluid. Mineral solubility is also assumed to be negligible such that Xfinalmin is equal to Xinitialmin. (6)

CLitotal/ Cfinalmin = Xfinalmin + (Cfinalfluid Xfinalfluid)/ Cfinalmin

The Nernst distribution coefficient is defined as, (7)

Dmin/fluid = Cfinalmin/ Cfinalfluid

Then, (8)

[CLitotal/ Cfinalmin] – Xfinalmin = 1/ Dmin/fluid Xfinalfluid

(9)

Dmin/fluid = Xfinalfluid/( [CLitotal/ Cfinalmin] – Xfinalmin)

The fluid-mineral ratio was such that the volume of the solution would serve as an infinite reservoir and the lithium concentration would remain constant throughout the experiment. Mineral/mineral distribution coefficients (Dmin/min) were calculated from: (10)

Dmin/min = CfinalminA/CfinalminB

where Cfinal is the lithium concentration of mineral A or mineral B in ppm. The range of lithium content in the run products was 5 ppm to 7 ppm in the clinopyroxene, 13 ppm to 466 ppm in the olivines, and 20 ppm to 70 ppm in the plagioclase. Mineral – fluid equilibrium was assessed in terms of run-product homogeneity. The concentration of lithium within a single experiment typically varied by 14 % relative to the mean concentration from grain to grain and in a few cases, lithium contents varied as much as 30 %. The 50 μm spot size and 80-second ablation time often meant that the laser analysis consumed the whole grain. The run product olivine grains typically had diameters < 75 m, the clinopyroxene > 75 m, and the anorthite grains > 100 m. Figure 2.4 shows the time resolved spectra for 7Li and 43Ca, measured by LA-ICPMS, for clinopyroxene produced in the lowest temperature partitioning experiment

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(800oC). Time resolved spectra for all experiments, with a few exceptions thought to be the result of fluid inclusions, display level spectra, which are interpreted as homogenous equilibrated grains. The slight sloping of the spectra is due to signal decay as a result of the geometry of the ablated pit and does not reflect sample heterogeneity as both the 7Li and 43Ca signals remain consistent with respect to each other. An additional test for equilibration between mineral samples and fluids was attempted by measuring partition coefficients in reversal experiments where crystals previously equilibrated with solution ‘A’ were re-equilibrated with solution ‘B’, containing a lower concentration of lithium and a differing isotopic composition. These reversal experiments confirm isotopic equilibrium; however, changes in mineral assemblages (zoisite in anorthite reversal, Mghydroxides in olivine reversal, monticellite in olivine + clinopyroxene reversal, undetermined phase in clinopyroxene reversal) make the mass balance calculations, used to determine the distribution coefficients, impossible to resolve. A simple calculation using the diffusion rates of lithium in clinopyroxene measured by Coogan et al. (2005) determines that at 800oC lithium diffusion should penetrate to a distance of 150 μm in 72 hrs, which is greater than the radius of the largest run product crystal (Figure 2.3). An experiment was attempted at 0.2 GPa and 800oC to ascertain the effect of pressure on the partitioning behavior of lithium, with no significant effect of decreasing pressure noted. Lithium values were always shifted with respect to the starting material. Any change in the lithium composition of the fluid due to uptake by the growing mineral is insignificant compared to the total amount of lithium in the fluid even when the DLimin/fluid is greater than one. Generally, the lithium mineral/fluid distribution coefficients decrease in the order: olivine (2.51 – 0.17), plagioclase (0.32 – 0.090) and clinopyroxene (0.32 – 0.07).

2.4.2.1 Clinopyroxene Run products produced in these experiments typically contained no phases other than clinopyroxene. One experiment, NCDL3, consisted of approximately 50 % (by volume) olivine crystals and 50 % clinopyroxene crystals, which was most likely a result of magnesium contamination from the ceramic pressure cell during sample assembly or loading. Electron microprobe traverses across individual clinopyroxene grains shows that the resulting crystals are homogenous with respect to major elements. As shown in Table 2.4, these clinopyroxene grains

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have much lower Al2O3 (~0.5 wt. % to ~0.2 wt. %) and FeO (~0.9 wt. % to ~0.4 wt. %) contents than natural upper mantle clinopyroxene (~3 wt. % Al2O3, ~2 wt. % FeO; Lundstrom et al., 2005). However, their MgO concentrations (~17 wt. % to ~21 wt. %) and Na2O concentrations (~0.02 wt. % to ~0.30 wt. %) are similar to those in clinopyroxene from mantle xenoliths (~17 wt. % MgO and ~0.30 wt. % Na2O; Lundstrom et al., 2005). The FeO content of all the experimentally produced clinopyroxene is low (< 1 wt. %) due to loss of Fe to the platinum capsule. Additionally because the starting clinopyroxene material contained very low (below detection limits) abundances of chromium and titanium these elements are absent in the run products. The range of DLi (0.07 to 0.613) for clinopyroxene/fluid measured in this study is similar to the clinopyroxene/fluid DLi (0.08 to 0.25) measured by Brenan et al. (1998b) and the clinopyroxene/silicate melt DLi (0.14 to 0.27) measured by Brenan et al. (1998a). Lithium partitioning between clinopyroxene and fluid decreases from 0.32 to 0.09 with increasing temperature from 800oC to 1100oC at 1 GPa. The temperature dependence of lithium partitioning between clinopyroxene and hydrous fluids can be demonstrated on a plot of ln DLicpx/fluid versus 1000/T, (Figure 2.5). A linear regression of the data yields the relationship: (11)

ln DLicpx/fluid = -7.3 (+0.5) + 7.0 (+0.7) * 1000/T

(R2=0.98)

where T is temperature in Kelvins.

2.4.2.2 Olivine Olivine partitioning experiments occasionally produced some oxide grains and in the reversal experiment (NCOR), an unidentified magnesian phase. All are likely due to incongruent dissolution of olivine. Electron microprobe traverses of individual crystals show the run product olivines to be homogenous with respect to major element chemistry. Iron loss to the platinum capsule resulted in considerably more magnesian (Fo# 98 to 99) olivines than those naturally occurring in the mantle. Reversal experiments were attempted; however, re-equilibrating run product material, from experiment NCOL2, caused Fe and water soluble elements (i.e. Na, Ca, etc.) to become further depleted, such that the final solid composition was no longer in the stability field of olivine. The same effect occurred with the plagioclase reversal run, NCAR, which resulted in the crystallization of zoisite.

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The range of DLi (0.17 to 0.57) measured in this study for olivine/fluid is similar to the olivine/silicate melt DLi (0.13 to 0.35) measured by Brenan et al. (1998a) with the exception of the experiment with added albite (OlAb10) which produced a value of 1.34. The temperature dependence of lithium partitioning between olivine and hydrous fluids can be demonstrated on a plot of ln DLiol/fluid versus 1000/T, (Figure 2.6). A linear regression of the data (excluding the reversal, NCOLR) yields the relationship: (12)

ln DLiol/fluid = -6.0 (+2) + 6.5 (+2) * 1000/T

(R2=0.82)

where T is temperature in Kelvins.

2.4.2.3 Plagioclase The starting material for the anorthite experiments was bytownite (An76), but the experiments that were not buffered for Na loss to the solution resulted in run product compositions very close to end member anorthite (An96-99). The reversal run, NCAR, which consisted of re-equlibrating material from NCA3, resulted in zoisite + unidentified Al-rich phase. At 1000oC run products consisted of 30 % melt, 70 % anorthite crystals, at 900oC the amount of melt was negligible, and runs at 800 oC were melt free. Electron microprobe traverses of these grains show that they are homogenous with respect to major elements. In two of the experiments, AnAb10 and AnAb20, the Na content of the fluid was buffered by addition of albite. This resulted in homogenous anorthite crystals with only slightly more sodic compositions (see Table 2.4). The range of DLi for anorthite – fluid partitioning measured in this study is 0.09 to 0.32. The range of DSr and DBa for plagioclase with similar An content is 1 -3 and 0.1 to 0.2 respectively (Blundy and Wood, 1991). Similar to the results from a study of Sr and Ba partitioning in plagioclase (Bludy and Wood, 1991), the data show a linear relationship with a negative slope on a plot of ln DLi versus XAn suggesting that lithium is more compatible in albite than in anorthite (Figure 2.7). Linear regression of the six partitioning experiments yields the relationship, in Jmol-1: (13)

RTlnDLi = 162,000 (+26,000)– 188,000 (+28,000) (XAn)

(R2=0.96)

26

where R is the gas constant, T is temperature in Kelvins, and XAn is the anorthite content of the plagioclase. Following Blundy and Wood (1991), RTlnDLi is used rather than lnDLi to minimize the effect of temperature in the linear regression.

2.4.3

Olivine-Clinopyroxene Pair Experiments

A series of experiments were done with olivine + clinopyroxene + fluid to investigate intermineral partitioning and isotopic fractionation. The run products from these experiments typically consisted of coarse-grained intergrowths of olivine and clinopyroxene and in the case of experiments 2m-lo and Yb-1, molybdenum oxide (from outer capsule material) and ytterbium oxide, respectively. Due to the uncertainties in the mass balance of each phase after equilibration, mineral-fluid D values have not been calculated for these runs. Experiment 2m-1 at 900oC contained only clinopyroxene; however, NCDL3 at 900oC stabilized both clinopyroxene and olivine. Experiment 2m-hi, investigating partitioning at log fO2 of –5, stabilized only enstatite. The reversal experiment, 2m-R, resulted in olivine + clinopyroxene + monticellite (CaMgSiO4; see NCOR and NCAR above). In all the experiments run at Ni-NiO, with the exception of the reversal experiment, 2m-R, olivine preferentially incorporated lithium relative to clinopyroxene. The range of DLi for olivine/clinopyroxene measured in this study is 1.2 to 6.7. Figure 2.8 shows how olivine/clinopyroxene partitioning increases with increasing temperature between 800oC and 900oC. The lithium content of the run product olivine in the 800oC experiment, 2m-2, is 101 ± 59 ppm; this large standard deviation suggests that the run product olivine may be heterogeneous with respect to lithium and may have formed lithium rich fluid inclusions. The reversal experiment, which equilibrated a split of the 2m-2 experiment at 800oC with a solution of 96 ppm lithium, did result in a lower lithium content in the run products than in the starting material. This experiment is complicated by the formation of monticellite, which contains 39 ppm lithium, more than either the forsteritic olivine or clinopyroxene (11 ppm and 26 ppm, respectively). The constant Dol/cpxLi versus temperature further suggests that olivine/clinopyroxene partitioning of lithium is independent of temperature. Also shown in Figure 2.8 is the ratio of DLiolivine/fluid/ DLicpx/fluid calculated from single-phase experiments, which at 1000oC is the same as that determined from the two-phase experiment.

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2.4.3.1 Experiments at variable fO2 Three experiments were conducted to investigate the effect of oxygen fugacity on the lithium partitioning behavior between olivine and clinopyroxene. Experiment, 2m-hi, which was run at 1000oC in a Pt + Re lined nickel capsule to generate an oxygen fugacity of log fO2 of –5, stabilized enstatite. The enstatite had a lithium content of 5.28 ppm and resulted in a DLient/fluid of 0.02, which is much lower than the DLimin/fluid for either clinopyroxene (0.17) or olivine (0.48 at Fo#98, or 0.17 at Fo#63) at the same temperature. Experiment 2m-lo was run at 1000oC in a Pt lined Mo capsule to generate an oxygen fugacity of log fO2 of –15 and resulted in olivine with 30 ppm lithium and clinopyroxene with 42 ppm lithium which results in a Dol/cpxLi of 0.7. The oxygen fugacity generated by the Ni lined Pt capsule at 1000oC is log fO2 = -10.3 and resulted in olivine with 52 ppm lithium and clinopyroxene with 13 ppm lithium which results in a Dol/cpxLi of 4.0. Higher oxygen fugacity results in higher abundances of Fe3+ relative to Fe2+. Incorporating a higher proportion of Fe3+ into the olivine structure should generate more charge balancing opportunities for the Li+ ion in the crystal structure; resulting in a coupled substitution where: (14)

Li1+ M1 + X3+ M2  (Mg2+, Fe2+)M1 + (Mg2+, Fe2+)M2

This is consistent with the results between the NNO run 2m-3, (higher fO2 and Dol/cpxLi = 4.0) and the 2m-lo run (lower fO2 and Dol/cpxLi = 0.7).

2.4.3.2 Experiments with added REE Experiment Yb-1, carried out at 1000oC with 0.25 mg of Yb2O3, was an attempt to determine the effect of rare earth elements (REE) on the relative partitioning of lithium between olivine and clinopyroxene. This experiment resulted in olivine with 175 ppm lithium and clinopyroxene with 17 ppm lithium, and resulted in a Dol/cpxLi of 10, which is significantly higher than the Dol/cpxLi of 4.0 at 1000oC that results from no addition of REE. The increased partitioning of lithium into the olivine is most likely a result of a coupled substitution with Yb3+, analogous to that in Equation 14 for Fe3+.

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2.4.4

Lithium Isotope Fractionation

2.4.4.1 Clinopyroxene- fluid Table 2.6 gives the isotopic ratios of the starting materials and run products from the isotopic fractionation experiments as well as the calculated ∆7Licpx-fluid, where; (15)

∆7Licpx-fluid = δ7Licpx (‰)- δ7Lifluid (‰)

With the exception of the reversal runs, all experiments had Dekalb diopside (δ7Li of +9.7 ‰) as starting material. Two sets of solutions were used: two L-SVEC based solutions, (A) with 243 ppm lithium and δ7Li of 0 ‰ and (B) 96 ppm lithium and δ7Li -2.7 ‰ for the reversal and two 6

Li doped solutions one (C) with 306 ppm lithium and δ7Li of -88.4 ‰ and (D) 180 ppm lithium

and δ7Li of –46.1 ‰ for the reversal. In all the experiments, the crystals were preferentially enriched in 6Li with respect to the fluid. Duplicate experiments at 900oC and run times of 72 hrs and 142 hrs had ∆7Licpx-fluid within + 2 ‰, which is within the precision of the analysis, indicating that run times were sufficient for isotopic equilibrium. As a further test, a reversal experiment, Ldi-12, was conducted in which a split of sample Ldi-10 with δ7Li of –90.9 ‰ was reacted with solution B (δ7Li –46.1 ‰). The run product clinopyroxene from Ldi-12 had a δ7Li of –49.5 ‰, an enrichment of 45 ‰ in the heavier isotope from its initial value, and resulted in a ∆7Licpx-fluid of –3.4 ‰, which is within the precision of the other experiments. The lithium isotopic fractionation at high temperatures (T>900oC) is +2.5 ‰, which is just at the limit of the analytical precision of this study. The range of ∆7Licpx-fluid measured in this study is from –0.3 ‰ to –3.4 ‰, and follows the trend of ∆7Licpx-fluid decreasing with increasing temperature. When the run products were not rinsed prior to sample digestion and analysis, the data produced scattered results, most likely due to the precipitation of lithium as the remainder of the solution was dried down after sample recovery. Figure 2.10 is a plot of the ∆7Licpx-fluid (‰) from this study, as well as the ∆7Lispodumene-fluid (‰) from Wunder et al (2006), the ∆7Libasalt-fluid (‰) measured between basalt and seawater at 350oC (Chan et al., 1993) and 2oC (Chan et al., 1992) versus 1000/T (K). Despite the fact that Wunder et al. (2006) used both OH- and Cl-bearing fluids and this study was chlorine free, with the lithium introduced as Li2CO3, all of the data plot on the same regression line empirically

29

determined by Wunder et al. (2006). There appears to be no difference in fractionation behavior with pressure (seafloor to 2 GPa) or complexing anion. It should be noted that measurements of lithium isotopic fractionation in the quartz-muscovitefluid system, from 400-500oC, found the quartz and the mica to be preferentially enriched in 7Li (Lynton et al., 2005), which is inconsistent with this study.

2.4.4.2 Olivine-Clinopyroxene Lithium Isotope Fractionation Table 2.6 gives the isotopic ratios of the starting materials and run products from the isotopic fractionation experiments as well as the calculated ∆7Liol-cpx, where; (16)

∆7Liol-cpx = δ7Liol (‰) - δ7Licpx (‰)

All experiments, with the exception of the reversal runs, had 82 wt. % Dekalb diopside (δ7Li of +9.7 ‰) + 18 wt. % San Carlos olivine (δ7Li of +1.0 ‰) as starting material. Two L-SVEC based solutions were used: (A) with 243 ppm lithium and δ7Li of 0 ‰ and (B) 96 ppm lithium and δ7Li -2.7 ‰ for the reversal. In all the experiments, except the reversal run, the isotopic composition of the olivine grains did not shift significantly; whereas, the isotopic composition of the clinopyroxene became as much as 15 ‰ lighter (i.e. from δ7Li of +9.7 ‰ to ~-5 ‰). ∆7Liolcpx

measured in this study is ~5 ±5 ‰, which is not resolvable with the precision of this study.

2.5 Discussion 2.5.1

Controls on Partitioning

2.5.1.1 Clinopyroxene Clinopyroxene has two sites, M1 and M2; M1 has six-fold coordination with respect to oxygen and M2 has eight-fold coordination. Given that the M1 site is slightly smaller than M2 (optimal site radius (ro) of ~0.7 Å versus M2 ro of ~1.1 Å) and has lower defect energies for univalent cations, it is likely that the primary site for lithium in the clinopyroxene structure is M1 where it can exchange for Mg2+ (Purton et al., 1997). Such an exchange should be coupled by a trivalent cation such as Al in jadeite (NaAlSi2O6) component or spodumene (LiAlSi2O6) or Fe3+ in a aegirine-like molecule, NaFe3+Si2O6. (17)

Li1+M1 + X3+M2  (Mg2+)M1 + (Ca2+) M2

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Due to Fe loss to Pt capsule the total iron content of the run product clinopyroxene is low, less than 2 wt. % and does not vary systematically with temperature. A Mössbauer study of natural diopside crystals by De Grave (2003) found all to contain some component of Fe3+ in either M1 or M2 sites. Furthermore, increasing Fe3+ at the expense of Fe2+ also serves to increase the Mg/(Mg+Fe2+) ratio (Luth and Canil, 1993) which should then increase the availability of sites for lithium exchange. Experiments examining the effect of fO2 on lithium diffusion in clinopyroxene appear to confirm this; lithium diffusion in clinopyroxene appears to increase with decreasing oxygen fugacity (Caciagli, Chapter 3). Previous work has shown lithium partitioning to increase slightly with increasing Al/Si ratio (Brenan et al., 1998b). A compilation of the data collected in this study and other experimental data displays a similar trend. In Figure 2.5, the partition coefficients determined from cpx/fluid experiments in this study, and Brenan et al. (1998b), and the cpx/melt experiments of Hart and Dunn (1993), Blundy (1998), Brenan et al. (1998a), and Blundy and Dalton (2000) are plotted as a function of temperature and the data points are labeled with the wt. % Al2O3 content of the run product clinopyroxene. It is important to note that the alumina contents of run product clinopyroxene in this study are approximately constant. Generally, lithium-partitioning at a given temperature increases with increasing Al2O3 content of the run product clinopyroxene, regardless if the partitioning measured is min/melt or min/fluid. For example, the run product clinopyroxene from the cpx/melt experiment of Blundy and Dalton (2000) at 1375oC (and 0.8 GPa) has a similarly low alumina content as the clinopyroxene from this study, and plots on the regression line determined in this study. As the alumina content of the clinopyroxene increases, lithium partitioning increases and the data fall above the regression line. Other elements, such as Fe and Na, appear to influence the lithium partitioning as well. For example, the experiment of Brenan et al. (1998b) labeled 1.5* was conducted at 900oC with 0.5 M aq NaCl, and has run product clinopyroxene with 1.5 wt. % Al2O3, but resulted in a lower cpx/fluid partition coefficient than measured in another 900oC experiment from that study containing 0.2 wt. % Al2O3 and the 900oC experiment containing 0.3 wt. % Al2O3 from this study. This relationship is consistent with Equation 17 where lithium substitution is coupled with Al3+. The correlation that is observed here may exist because the total Al2O3 content is correlated with, or somehow serving as an indicator of the amount of Al3+.

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2.5.1.2 Olivine The solution energies calculated by Purton et al. (1997) for forsterite demonstrate that the lowest energy pairing is a 3+ cation in the M2 site coupled with Li+ in the M1 site. Previous studies have shown lithium partitioning between olivine and silicate melt to be coupled with Al3+ (Suzuki and Akaogi, 1995; Taura et al., 1998), suggesting the following mechanism. (18)

Li1+M1 + X3+M2  (Mg2+, Fe2+)M1 + (Mg2+, Fe2+) M2

Comparison with other experimental data suggests that Fe may be affecting the partitioning of lithium into olivine. In Figure 2.6, the partition coefficients measured from ol/fluid experiments of this study, and the ol/melt experiments of Brenan et al. (1998a), Taura et al. (1998), Blundy and Dalton (2000) and Zanetti et al. (2004) are plotted as a function of temperature and the data points are labeled with the wt. % FeO in the run product olivine. The olivine-melt partitioning experiments with low or no FeO plot on the same regression line as those determined in this study. Experiments with run product olivine containing 8-10 wt. % FeO have higher lithium partition coefficients than experiments with lower FeO contents in run product olivine, at a given temperature. Assuming that at a given fO2, and temperature the total Fe3+ content of olivine scales with total Fe content, the trend of increasing lithium partitioning in olivine with increasing FeO, suggests that lithium may be coupling with Fe3+ as substitution mechanism (see Equation 18). The experiment of Zanetti et al. (2004), labeled 19*, further supports this suggestion. The Zanetti et al. (2004) experiment was conducted at an fO2 equivalent to QFM-2, which is almost four orders of magnitude more reducing than the experimental conditions of this study (NNO), and resulted in a much lower lithium partition coefficient, despite its high (19 wt. %) FeO content. More reducing experimental conditions in the Zanetti et al. (2004) experiment would result in lower Fe3+ contents, and therefore less favorable conditions for Li+ substitution, than in the experiments from this study, despite the fact that the olivine has a high FeO content.

2.5.1.3 Plagioclase The primary control on lithium partitioning between plagioclase and hydrous fluids is the composition of the feldspar (Figure 2.7). This is similar to the Sr and Ba partitioning behavior observed between plagioclase and silicate melts or hydrothermal solutions (Blundy and Wood,

32

1991; Lagache and Dujon, 1987). In the case of Sr or Ba, both these cations are divalent and based on size and charge balance considerations should be accepted more readily into the anorthite structure in exchange for Ca2+ rather than Na+ in the albite structure. This apparent discrepancy is explained by the highly elastic nature of the albite structure (Blundy and Wood, 1991). Albite has a lower bulk modulus and a lower shear modulus than anorthite, which results in an increased “flexibility” of the albite crystal structure (Angel et al., 1988; Blundy and Wood, 1991). These results suggest that the albite crystal lattice would better accommodate Li+, despite the fact that Na+ is larger than Ca2+, than the more rigid anorthite structure.

2.5.1.4 Intermineral Partitioning Previous studies of the lithium content in mantle xenoliths have made a correlation between lithium contents of olivine and clinopyroxene pairs and xenolith paragenesis, e.g. equilibrated, metasomatised, etc., (Figure 2.9; Seitz and Woodland, 2000; Paquin and Altherr, 2002; Woodland et al., 2002; Woodland et al., 2004). Specifically, olivine-clinopyroxene pairs that are apparently equilibrated (are in chemical equilibration, have no major inhomogeneities or mineral zoning; Seitz and Woodland, 2000) tend to fall on a linear ~1:1 trendline when lithium abundance of the olivine is plotted against the lithium abundance of the clinopyroxene. The xenoliths apparently metasomatised by silicate melt and hydrous fluids fall below the trend line (depleted in olivine relative to clinopyroxene), and those altered by carbonatite melt fall on and above the trend line (enriched in olivine relative to clinopyroxene). Figure 2.9 also includes experimental data corresponding to similar metasomatic regimes. The higher concentrations of lithium are due to experimental requirements and analytical detection limits. All the experimental data fall in a relatively restricted range of Dol/cpxLi of ~1 or > 1. A silicate melt equilibrated olivine-clinopyroxene pair, from Brenan et al. (1998b), plots slightly below the line projected from the equilibrated mantle xenoliths; and carbonatite melt olivine-clinopyroxene pairs, from Blundy and Dalton (2000), fall on the projected line, above the silicate melt experiment. Olivine-clinopyroxene pairs equilibrated with hydrous fluids in this study fall above the equilibrated mantle xenoliths trend, not below, where the hydrous fluid metasomatised samples plot. Despite the diversity of experimental methods, the range of Dol/cpxLi exhibited in natural samples is not reflected in experimental studies, as no experimental studies have shown Dol/cpxLi < 1.

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One possible explanation is that the clinopyroxene compositions resulting from the hydrous fluid experiments are very low in Al2O3, (0.2 wt. % to 0.3 wt. %); much lower than the Al2O3 content of clinopyroxene found in mantle most xenoliths (2 – 6 wt. %, Seitz and Woodland, 2000). Clinopyroxene-fluid partitioning experiments of Brenan et al. (1998b) have shown that lithium partitioning increases with increasing Al content of the pyroxene. The high Dol/cpxLi values from this study may be a reflection of the low Al2O3 content of the pyroxene. In the case of carbonatite melt equilibrated olivine-clinopyroxene pairs from Blundy and Dalton (2000), their experiments contain similar Al2O3 compositions as those found in carbonatite melt metasomatised xenoliths, and the discrepancy between experimentally determined Dol/cpxLi,, (values of ~1 or > 1) and those measured in mantle xenoliths (D < 1) still exists. Recent studies of lithium and lithium isotopes in mantle xenoliths and other rocks have found that intermineral partitioning and fractionation of lithium often will not correlate with expected equilibrium values. These apparent disequilibrium signatures are attributed to remobilization of lithium with differing rates of diffusion between olivine and clinopyroxene (Parkinson et al. 2007; Rudnick and Ionov, 2007; Jeffcoate et al., 2007).

2.5.2

Controls on Isotopic Fractionation

Lithium isotopic fractionation between minerals and fluids depends on the difference in the zero point potential energy (ZPE) between the phases of interest. 7Li is heavier and has a lower vibrational frequency, and therefore a lower ZPE than 6Li (Chacko et al., 2001). The phase that will undergo the greatest reduction in ZPE will preferentially take 7Li over 6Li (Chacko et al., 2001). This has been demonstrated by Ab initio calculations, which have predicted that during mineral-solution reactions 6Li, is preferentially incorporated into octahedrally coordinated sites in the solid, and 7Li is preferentially incorporated into the dominantly tetrahedrally coordinated sites in the fluid (Yamaji et al., 2001). The coordination state and bonding environment of lithium in both spodumene, a clinopyroxene with up to 7 % lithium content, and Ca-clinopyroxene, in this case diopside with 6 to 60 ppm lithium, is octahedral. The consistent fractionation between spodumene and aqueous fluids at 2 GPa (Wunder at al., 2006); clinopyroxene and aqueous fluids at 1 GPa (measured in this study); and altered seafloor basalts (Chan et al., 1992; Chan et al., 1993) suggests that the coordination state of lithium in aqueous fluids does not change throughout this range of conditions.

34

2.5.3

Lithium Incorporation into the Mantle

Pristine mid-ocean ridge basalts (MORB) contain 5-6 ppm lithium, with an average δ7Li of +4 ‰, and resemble the mantle with respect to lithium content and composition (Jagoutz et al., 1979; Moriguti and Nakamura, 1998; Tomascak, 2004; Tomascak et al., 2008). The mantle source that produces MORB is thought to also provide the source for IAB, after modification by a slab-derived flux. Given that neither partial melting nor differentiation and crystallization will cause significant fractionation of δ7Li (Tomascak et al., 1999), variations in lithium content and isotopic composition in some arc lavas are believed to arise from the slab inputs to the melt source regions. During prograde metamorphism the mineral assemblages in the subducted slab become more anhydrous with increasing pressure and temperature. Fluids produced by dehydration reactions in the subducting slab add fluid mobile elements to the overlying mantle wedge; a signature that is believed to be reflected in the lavas derived from this re-hydrated mantle. For example, Kamchatka arc lavas are most enriched in boron relative to Nb or Zr at the arc front and the enrichment decreases to MORB values with increasing slab depth (Ishikawa et al., 2001). This is suggestive of continuing mobilization of fluid mobile elements into the arc source region by fluids derived from dehydration reactions in the down going slab (Leeman, 1996; Ishikawa and Tera, 1999; Ishikawa and Nakamura, 1994; Ishikawa et al., 2001). The Izu arc is one of the few localities where a clear correlation between lithium content, boron content, δ7Li, and distance to the arc front can be made. The δ7Li in the Izu lavas range from +7.6 ‰ in the arc front, to +1.1 ‰ in the back arc; this is thought to reflect enrichment of the arc melt source by fluids derived from the down going slab (Moriguti and Nakamura, 1998). Typically the lithium isotopic composition of most arc lavas ranges from δ7Li approximately +5 ‰ to +1 ‰ and shows a slight negative correlation with Li/Y ratio (Tomascak et al., 2002). It is not clear why correlations between lithium content and δ7Li of arc lavas are not consistent among all arcs. Variations in the extent of slab dehydration due to slab age, angle of subduction and overall thermal regime have been suggested as possible factors (Moriguti et al., 2004). However, differences in the Li/Y content and 7Li composition between the Kurile arc lavas and the Izu arc lavas, or even the Japan arc lavas, persist despite similarities in the subduction regime and age of the slab (Morguti et al., 2004; Tomascak et al., 2002). Another suggestion is that slab-derived fluids are significantly modified during transport through the mantle wedge to the melt source, and that the lithium signal is attenuated by interaction with mantle minerals (Tomascak et al.,

35

2002). Subtle differences in the mode of fluid transport through the mantle wedge could lead to significant differences in the overall behavior of lithium in subduction zones.

2.5.4

The Mantle Wedge as a Chromatograph

The fluids derived from the dehydrating slab during subduction are potentially very lithium rich depending on the nature of the subducted sediments, in some cases containing as much as 2000 ppm lithium or more (Chan et al., 2002). However, consistent and clear correlations of lithium with other fluid mobile elements, such as boron, are rare. More commonly, a slab-like lithium signal cannot be correlated with other indicators of fluid involvement such as B/Be ratios or depth to slab. For example, some of the calc-alkaline lavas belonging to the Panamanian Old Group lavas have high B/Be contents, suggesting high fluid input, and MORB-like δ7Li (+4.7 to +5.6 ‰; Tomascak et al., 2000). Similar behavior is found in other Central American lavas (Chan et al., 2002), as well as lavas from the Aluetian and Kurile arcs (Tomascak et al., 2002). It has been suggested that the relatively compatible mineral-fluid partition coefficients for lithium, the rapid diffusion of lithium into mantle minerals, and the high rock/fluid ratio experienced by the fluids in the mantle wedge can provide a mechanism by which the lithium signal is decoupled from other fluid mobile trace elements in slab-derived fluids (Tomascak et al., 2000; Tomascak, 2004; Wunder et al., 2006). The diffusion coefficients measured in this study provide constraints as to the time required for fluid-mineral equilibrium. This information, coupled with measurements of lithium partitioning and isotopic fractionation between fluids and mantle minerals, allows for quantitative modeling of the interaction between slab-derived fluids and the mantle wedge during fluid transport from the slab to the arc melt source. The effect of mantle wedge and fluid interaction can be evaluated following the method of Navon and Stolper (1987), who modeled the distance traversed by various elements flowing through an ideal mantle column of fixed porosity. This ideal column contains solid rock with an interconnected fluid network along grain edges. Assuming that partition coefficients are constant, the fluid fraction is uniform, and the densities and diffusivities of the solid and fluid do not vary across the column length then:

36

(19)

C f t

 X fVf

C f z

0

where Cf is the trace element concentration in the fluid, Xf is the mass fraction of the trace element in the fluid, Vf is the velocity of the fluid, t is time and, z is the distance traversed along the column. Assuming solid-fluid equilibrium is maintained, the most incompatible trace elements have fronts that travel farther than trace elements that are more compatible for any given time. The transport velocity of a trace element relative to the transport velocity of the fluid, (Vtr/Vfl), is equal to the mass fraction of the trace element in the fluid (Xf): (20)

Xf = f / (f +(1-)sD)

Where  is the volume fraction of fluid in the column (assumed to be 0.03 by Navon and Stolper, 1987), f and s are the fluid density and mantle wedge density; (assumed to be 1 g/cm3 and 3 g/cm3 respectively) and, D is the bulk partition coefficient for the element of interest. A bulk DLi of 0.42 was calculated, assuming 80 % olivine and 20 % clinopyroxene, using the DLi of olivine-fluid and cpx-fluid measured in this study. The olivine-cpx boron data (from Brenan et al., 1998a) and cpx-fluid boron (from Brenan et al., 1998b) were combined to estimate the bulk DB for the same lherzolite assemblage and XfB was also calculated as above for comparison. From Navon and Stolper (1987) the rate at which a point of constant concentration moves through the column (Vtr) is: (21)

 z     Vtr  X f V f  t  Cf

So for a given column length, at the time that the fluid front reaches the melt source, the boron front will be 91 % of the column length, and the lithium front will only be 2 % of the column length. The maximum capacity of the column for each element can be determined by calculating at what time the trace element front reaches the top of the column, relative to the time the fluid reaches the top of the column. (22)

L X fVf t 1   tc L Vf Xf

37

where tc is the time for the fluid front to reach the top of the column, t is the time for the trace element front to reach the top of the column, L is the column length, Vf is the fluid velocity, and Xf is the mass fraction of the trace element in the fluid. For a given column length and a fluid velocity, the boron front will reach the melt source at approximately the same time as the fluid front (1.08tc); however, for the lithium front to reach the top of the column requires the column to be filled ~42 times. This is a strikingly large volume of fluid, requiring a total of 1.26 cm3 of fluid for every 1 cm3 of rock for a column with 3 % porosity, which is equivalent to ~40 wt. % fluid. The highest estimates of fluid involved in arc magmatism from the literature is ~20 wt. % (Ayers, 1998) and most estimates range from 1-5 wt. % (Stolper and Newman, 1994). Even if the subducting slab has the capacity to generate such a large quantity of fluid, the time required to deliver this amount of fluid to the melt zone needs to be considered. Measurements of U-series disequilibria provide constraints on the timescales of fluid-mobile element transport from the slab to the melt source (Elliott et al., 1997). Young lavas from subduction zones often contain an excess of 238U relative to 230Th, or [238U]/[230Th] >1 (Elliott et al., 1997). Unlike Th, U readily partitions into oxidized fluids, therefore a [238U]/[230Th] ratio >1 is believed to be the result of the addition of a slab derived fluid containing both 238U and 234U (the parent of 230Th; half-life ( ~250 kyr), to the melt source within the last 30, 000 years (Elliott et al., 1997). Fluid velocities predicted from U-series disequilibria range from 4 to 10 m/yr. Given these velocities, the time required for the lithium signal to reach the melt source can be compared with that for the boron signal. Assuming that partial melting of peridotite occurs at depths of 100 km below intra-oceanic arcs (Plank et al., 2009) and a Benioff zone of ~125 km, then a maximum column length will be ~25 km. As shown in Figure 2.11, it would take the boron signal 100-5000 years to travel 25 km to the melt source, whereas the lithium signal will need between 10,000 years and 200,000 years to reach the melt source. If fluid velocities are 4-10 m/yr then, only the boron signal will reach the melt source while 238U and 230Th still maintain measurable isotopic disequilibrium. Other studies have made similar estimates of flux rates using 226Ra-230Th (Sigmarsson et al., 2002), the large excess of 226Ra over 230Th displayed in many young lavas requires fractionation, presumably due to fluid transport, to occur within 8 ka, or ~5half-lives (Sigmarsson et al., 2002). These time constraints give rise to fluid velocities of 10-100 m/yr (Sigmarsson et al., 2002). Only at the highest estimated fluid velocity, the lithium signal will reach the melt source within 10,000

38

years; this will satisfy the time constraints determined from measurements of U-series disequilibria and Ra-Th disequilibria (~10,000 yr; Figure 2.11). More recent studies have suggested that slab dehydration produces a zone of hydrated lithosphere, mainly consisting of chlorite, which is down dragged by corner flow to depths where melting may take place in the asthenosphere given the right subduction geometry (Grove et al., 2009). In this scenario mantle melting occurs 50-100 km below intra-oceanic arcs depending on the angle of slab-dip and the slab convergence rate. Assuming the Benioff zone is between 100 and 125 km, the minimum column length will be ~10 km. As shown in Figure 2.11, the boron signal needs only 100-2000 years to travel 10 km, whereas the lithium signal will need 4000100,000 years to reach the melt source. Even if the column length is very short, in order for the lithium signal to reach the melt source within the timescales constrained by U-series and Ra-Th disequilibria, fluid velocities between 10 m/yr and 100 m/yr are necessary. It should be noted that lithium partition coefficients for cpx/chlorite have been estimated from lithium concentrations in the eclogites from Syros, Greece, and are ~1, therefore it is expected that chlorite/fluid partitioning will be similar to clinopyroxene/fluid partitioning with respect to lithium (Marschall et al., 2006). This suggests that for the lithium signal to be correlated with the boron signal, as well as other fluid mobile elements, extremely high fluid volumes and velocities are required. The rather improbably high fluid volume and extremely rapid fluid velocity required to transport the lithium signal from the slab to the melt source is consistent with the lack of a slab-derived lithium signal in many arc volcanics. The lithium signal will not reach the melt source because it will preferentially partition into the mantle wedge, relative to other fluid mobile elements (such as boron). Given that subducting slabs are unlikely to generate the large fluid volumes required to transport a lithium signal to the melt source, the occurrence of a slab-like lithium signal in arc lavas implies a mode of fluid transport other than percolation. The rapid fluid velocities required by Ra-Th disequilibria have led to the suggestion of fluid transport by hydrofracture rather than percolation (Davies, 1997). Because fluid transport would be limited to fractures on the scale of 300-1500 m long and 10-200 mm wide (Davies, 1997), the fluid volume needed to generate very high fluid/rock ratios is more reasonable because the fluids only interact with a small volume of

39

rock. Since lithium partitions preferentially into the mantle relative to other fluid mobile elements, minimal rock interaction would result in more lithium being transported to the melt source.

2.5.5

Isotopic Evolution of Lithium-Bearing Fluids in the Mantle

2.5.5.1 Percolation and Rayleigh Distillation Interaction of lithium bearing fluids with mantle minerals will result in changes to the isotopic composition of both phases. If the shift in the isotopic composition of the fluid and the mantle are known, then the amount of fluid: rock interaction can be estimated. The degree of fractionation resulting from fluid-rock interaction can be modeled assuming a simple Rayleigh distillation model: (23)

 7 Li fluid  ( 7 Li slabfluid  10 3 ) f ( 1)  10 3

Where 7Lifluid is the altered fluid, 7Lislabfluid is the initial composition of the slab-derived fluid, and f is the fraction of the element in the fluid remaining after interaction with the mantle wedge. In this case  is calculated from the degree of cpx-fluid fractionation measured at 1100 oC in this study and is defined as: (24)

 = (7Limin + 1000)/(7Lifluid+1000)

The calculated here is 0.999, which is consistent with the  calculated using data from the study of Wunder et al. (2006). When  is < 1, continued interaction of the fluid with mantle minerals, i.e. distillation, will result in progressively heavier fluids. Figure 2.12 shows how the isotopic composition of a fluid with an initial δ7Li of +9.7 ‰ (the slab input estimated by Moriguti and Nakamura, 1998) would change during percolation through a mantle column. The isotopic composition of the altered fluid increases and is heavier than the initial slab-derived fluid and both the fore arc and back arc lavas of the Izu arc (δ7Li = +7.6 ‰ and +1.1 ‰, respectively) for any amount of fluid/rock interaction. It is important to note that the  used in the above calculations is determined for cpx-fluid fractionation at 1100oC. The temperature of the mantle will be lower near the subducted slab. Depending on the age of the crust, the rate of subduction, and the degree of frictional shear

40

heating the temperature in the mantle above the slab may be as low as 700-800oC (Peacock, 1993). Because fractionation of stable isotopes tends to increase with decreasing temperature (Urey, 1947) the fractionation occurring at the base of the mantle column, close to the top of the slab, may be even larger. Additionally, the degree of fractionation between fluids and the mantle may be greater than what has been calculated above since the value used in the above calculation was determined from cpx-fluid fractionation experiments and assumes that the fractionation of lithium between olivine and fluids is the same. For reference, curves for  values of 0.998 and 0.996 are plotted on Figure 2.12, showing the effects of greater fractionation factors on the fluid composition. As the fluids percolate through the mantle wedge the fraction of lithium in the fluid decreases, and the δ7Li of the fluid becomes progressively greater. Because lithium is readily taken up by mantle minerals, the fraction of lithium remaining in the fluid becomes very small, Xf → 0.2 (see above), and depending on , the isotopic composition becomes extremely fractionated with δ7Li ranging from +15 ‰ to +35 ‰ (depending on ; Figure 2.12). Because 7Li preferentially fractionates into fluids, interaction of slab fluids with mantle minerals, i.e. percolation, will generate heavier, more 7Li-rich fluids. A fluid with an initial δ7Li of +9.7 ‰ percolating through the mantle wedge could not generate the isotopic signature observed in the Izu arc lavas. Either the initial slab fluid is lighter, or the isotopic signature is the result of mixing the 7Li-rich altered fluid and a lighter mantle reservoir. Interestingly, to generate the isotopic composition of the Izu fore arc lavas using the Xf as calculated above, requires an initial slab input with δ7Li = +4 ‰; essentially a fluid with MORB-like δ7Li. For the Izu lavas of the backarc region, with δ7Li = +1.1 ‰, it is unlikely that any component of the slab derived fluid has made its way to the melt source by percolation, as even the least altered slab-derived fluid is isotopically heavier than the unaltered MORB source mantle. The implication here is that this isotopically light lithium is not due to percolation of the fluid through the mantle, but must be the signature of a component derived from an isotopically light reservoir, such as the residual slab or oceanic sediments (δ7Li ~ -2 ‰; Moriguti and Nakamura, 1998). Another possibility is that this signal is due to an entirely different mechanism of isotopic fractionation and transport, which is discussed in the following section.

41

2.5.5.2 Generation of 6Li-rich fluids Ab initio calculations have demonstrated that during mineral-solution reactions 6Li should be preferentially incorporated into octahedrally coordinated sites in solid phases (Yamaji et al., 2001). Recent work by Jahn and Wunder (2009) has examined how lithium speciation in hydrous fluids affects isotopic fractionation. From Ab initio molecular dynamic (AIMD) calculations, they have determined that during fluid-solid fractionation, 6Li will prefer sites with the higher coordination. Lithium in pyroxene and olivine, the most abundant mantle minerals, is in octahedral or six-fold coordination. When fluid densities are less than 1.0 g/cm3 coordination of lithium in the fluid is mainly three-fold (Jahn and Wunder, 2009), and therefore 7Li will preferentially partition into the fluid phase. As fluid density increases, the coordination of lithium in the fluid also increases. When fluid density is greater than 1.2 g/cm3, the proportion of 5-fold and 6-fold coordinated lithium increases and the proportion of 3-fold and 4-fold coordinated lithium decreases, and the overall average lithium coordination in the fluid is greater than 4.5 (Jahn and Wunder, 2009). When lithium coordination in the fluid becomes greater than lithium coordination in the mineral phase the sense of fractionation changes, and 6Li is predicted to preferentially partition into the fluid phase. This change in sense of fractionation has been observed during staurolite-fluid partitioning experiments at 3.5 GPa (Wunder at al., 2007). Lithium is in tetrahedral coordination in staurolite and preferentially incorporates 7Li at 3.5 GPa. Therefore, 6Li-rich fluids may be generated by mineral-fluid fractionation at high pressures. Figure 2.13 is a plot of fluid density vs. temperature, with the average calculated Li-coordination shown as degree of shading. Superimposed on this plot are the fluid densities calculated with the CORK-EOS (Holland and Powell, 1991) using the Perple_X computer program (Connolly, 2005) for Franciscan and Alpine subduction zones (Ernst, 1988) as well as the most direct path between the slab and a fore-arc volcano (Peacock, 1993). It is possible to generate 6Li-rich fluids when mineral-fluid interaction occurs at depths greater than ~125 km. These results suggest that as fluids percolate up through the mantle wedge, fluid density decreases and the average coordination of lithium in the fluids will decrease; 6Li will once again preferentially fractionate into the mineral phase and the fluids will become heavier. In order to preserve the 6Li-rich signal generated at depth, fluids need to reach the melt source having undergone minimal interaction with the mantle on their ascent path. Fluid transport by hydrofracture would satisfy this

42

requirement, as it results in high fluid-rock ratios, which would transport the 6Li-rich fluids to the melt source quickly, and with least amount of interaction with the mantle wedge.

2.5.5.3 Generation of 6Li-rich zones in the mantle Hyrdofracture of the mantle by slab-derived fluids is an appealing mechanism to transport lithium through the mantle, as channelized flow through hydrofractures would satisfy the high fluid: rock ratios and rapid fluid velocities required by mineral-fluid partitioning. This transport mechanism would also minimize isotopic fractionation by limiting mineral-fluid interaction, thereby effectively propagating a slab signal all the way to the melt source region. The isotopic composition of the mantle wall rock of the fractures would also shift, generating a local isotopically light region in the mantle wedge. The isotopic shift of the wall rock depends on the extent of reaction between the mantle and the slab-derived fluids. If mineral-fluid exchange is fast, then local mineral-fluid isotopic equilibrium will occur, and the isotopic shift will depend on the fluid-rock ratio. This process can be modeled after the approach of Abart (1995; after Taylor 1977) by calculating the progress of the reaction,  which is defined as the ratio between the observed isotopic shift in the rock and the maximum attainable isotopic shift: (25)

 Lif R   Lii R  i  Li F   R  F   Lii R

In this case, LifR is the final isotopic composition of the mantle, LiiR is the initial composition of the mantle, LifF is the composition of the metasomatizing fluid and R-F is the fractionation between the mantle and the fluid. The value of will be between 0 (no equilibration) and 1 (complete equilibration). Where mineral-fluid exchange is rapid, as is the case for lithium exchange, then the degree of equilibration depends on the lithium atom equivalent fluid-rock ratio, N (Taylor, 1977): (26)

N = - ln(1 - )

The final isotopic composition of the mantle wall rock is estimated here given an initial mantle composition of δ7Li = +4 ‰ (Moriguti and Nakamura, 1998; Tomascak et al., 2008; Nishio et al., 2002). This model assumes complete equilibration between the mantle wall rock and fluid

43

will result in the mantle wall rock having a δ7Li‰ lower than the initial fluid composition (cpx-fluid = -1 ‰ at 1100oC; this study, Wunder et al., 2005). Because the isotopic composition of the initial slab fluid is not very well constrained, three different initial fluid compositions are used in this illustration; δ7Li = +10 ‰ (estimate for the Izu arc fluids by Moriguti and Nakamura; 1998), 0 ‰ and -10 ‰, the latter being arbitrary values reflecting generation of 6Lirich fluids at depth (Jahn and Wunder; 2009). Given a DLibulk = 0.42, N becomes (1/ DLibulk)*N in weight units. Assuming the isotopic composition of the fluid does not change, which is the case if the fluid/rock ratio is high, the isotopic composition of the vein/hydrofracture wall rock will approach complete equilibrium with the fluid; equal to R-F of -1 ‰. Because the timescales of lithium diffusion in olivine and clinopyroxene are rapid compared to the fluid transport times (2 m/1hr; see Ch.3, vs. 100 m/yr; Sigmarsson et al., 2002) the isotopic shift in the wall rock is a function of the amount of fluid available. Figure 2.14 shows the evolution of the final isotopic composition of the metasomatized mantle wall rock with increasing fluid/rock ratio. The isotopic shift of the wall rock depends on the extent of reaction between the mantle and the slab-derived fluids; if the fluid/rock ratio is high, the isotopic composition of the vein/hydrofracture wall rock will approach complete equilibrium with the fluid equal to R-F of -1 ‰. Also plotted are the δ7Li values of the Izu fore arc and back arc lavas (which are typical of the range of δ7Li values found in many arc lavas; Tomascak et al., 2002). The entire range of δ7Li values found in arc lavas can be achieved by metasomatizing the mantle with fluids that have δ7Li between 0 ‰ and +10 ‰ and fluid/rock ratios >1.2. Both these values are reasonable given the isotopic composition of subducted material and typical fluid/rock ratios for hydrofractured zones. Values of δ7Li in seafloor sediments range from -5 ‰ to +20 ‰ (Marschall et al., 2007 and references therein) and altered oceanic crust has δ7Li of ~ +14 ‰ (Moriguti and Nakamura, 1998; Chan et al., 1992). Slab derived fluids with δ7Li between 0 ‰ and +10 ‰ could be achieved during dehydration of the slab, recalling that isotopic fractionation in a cool slab at depths greater than ~125 km are likely to produce fluids that are isotopically lighter than the solid (Jahn and Wunder, 2009). The fluid/rock ratios in vein systems can be very high with typical values from calcite or quartz vein systems ranging from 70-100 cm3 of fluid per 100 cm3 of rock to as much as 1400 cm3 of fluid per 100 cm3 of rock (Spear, 1993).

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2.6 Conclusions Lithium is moderately incompatible in the mantle during mineral – fluid exchange reactions. The DLi measured in this study ranges from 1.34 – 0.14 in olivine, to 0.32 – 0.09 in plagioclase and, 0.32 – 0.07 in clinopyroxene. Lithium partitioning between clinopyroxene and hydrous fluids is a function of temperature, decreasing with increasing temperature from 800oC to 1100oC at 1 GPa and appears to increase with increasing Al2O3 content of the pyroxene. Olivine-fluid partitioning of lithium is not a function of temperature, but appears to be sensitive to Mg/Fe content, although this needs to be investigated more systematically. Lithium partitioning in anorthite is a function of feldspar composition, similar to the partitioning of other cations in the feldspar-fluid system. Lithium partitioning between olivine and clinopyroxene appears to be independent of temperature; however, preliminary experiments examining the effect of REE content and fO2 suggest that DLiol/cpx may be a function of crystal chemistry. Isotopic fractionation between clinopyroxene and fluid has been measured as well as between olivine and clinopyroxene. The isotopic fractionation between clinopyroxene and fluid at 900oC is ~ +1 ‰ (±2 ‰) and the measured isotopic exchange between olivine and clinopyroxene is ~ +5 ‰ (±4 ‰). Isotopic fractionation between clinopyroxene and fluids is a function of temperature and consistent with what has been observed in the spodumene – fluid system. The fractionation between spodumene and hydrous fluids results in an enrichment of 7Li in the fluid from +3.5 ‰ at 500oC to ~ +1.0 ‰ at 900oC and 2.0 GPa (Wunder et al., 2006). Application of these data to models of fluid-rock interaction in the mantle wedge reveals that lithium is a moderately incompatible element in the mantle during mineral-fluid exchange reactions. Because lithium is not a conservative element, it cannot be used to deconvolve the proportions of slab-derived fluid and altered and unaltered MORB-source involved in generating arc lavas. However, constraining how lithium behaves in the mantle provides some insight into the lithium and lithium isotopic trends, or lack thereof, observed in arc lavas. The absence of high Li/Y ratios in arc lavas with high B/Be, or MORB-like δ7Li in lavas with high B/Be contents (such as the lavas from the Sunda arc, Indonesia; Tomascak et al., 2002), can be explained by partitioning lithium into mantle minerals as fluids percolate through the mantle wedge. In these cases, transport through the mantle wedge completely removed the lithium signal from the slab-derived fluid. Convergent margin lavas, such as the Izu forearc lavas, with

45

δ7Li values greater than the mantle values (δ7Li ~ +4 ‰) are likely the result of some component of slab fluid-mantle interaction during percolation. It is important to note that very high fluid fluxes are implied if a 7Li signal from slab-derived fluids is to reach the melt source by percolation. Low δ7Li values (< MORB; δ7Li ~ +4 ‰) that correspond with high Li/Y ratios are likely generated near the slab and transported to the melt source with a minimum amount of interaction with the mantle wedge; here transport through hyrdofractures is a likely mechanism. The trend of increasing δ7Li with decreasing Li/Y, which is observed in most arc lavas (Tomascak et al., 2002), could be viewed as a spectrum between the two scenarios. Where low Li/Y values correspond with high δ7Li, large fluid fluxes were most likely percolating through the mantle wedge. Where high Li/Y values correspond with low δ7Li, the fluids were likely generated at depth and transported through the mantle through hydrofractures, having minimal interaction with the wedge. Intermediate values could be a result of some component of both these mechanisms. Transport of slab-derived fluids through hydrofractures in the mantle can also explain the lack of clear and consistent correlations between lithium and other fluid mobile elements. Fluids transported to the melt source through hydrofractures would be subject to differing degrees of mantle interaction (variable fluid/rock ratios and transport velocities). Lithium is moderately compatible in the mantle and diffuses rapidly; therefore, lithium contents and isotopic compositions will be very sensitive to variations in the types of mineral-fluid interaction. The lithium isotopic evolution of the mantle will also be affected by these processes, as it is such an efficient sink for lithium. Dehydration reactions in the subducting slab at depths less than ~125 km, where fluid density is relatively low, and the predicted predominance of three-fold and four-fold coordinated lithium in the fluid will generate 7Li-rich fluids and result in localized 6Li enrichment of the mantle. Hydrous fluids generated deeper than ~125 km are predicted to contain lithium in coordination states greater than four-fold, and therefore likely to be enriched in 6Li, at least initially, giving rise to a zone of 7Li-rich mantle at depth. Xenoliths with δ7Li values greater than MORB are uncommon, but have been found in blueschists from Syros (Greece), eclogites from Dabishan (China), Cima di Gagnone and Trescolmen (Alps) and lherzolites from Northern Japan and SE Austrailia (Nishio et al., 2004; Marschall et al., 2007).

46

Table 2.1 Composition of Starting Material Dekalb Diopside SiO2

54.77

(0.80)1

Al2O3

0.66

(0.10)

San Carlos Olivine

40.95

(0.02)