The Pennsylvania State University
The Graduate School
College of Earth and Mineral Sciences
THE EFFECT OF OXYGEN FUGACITY ON THE EQUILIBRIUM
PARTITIONING OF LITHIUM BETWEEN OLIVINE AND CLINOPYROXENE
A Thesis in
Geosciences
by
James A. Deane Jr.
© 2013 James A. Deane Jr.
Submitted in Partial Fulfillment
of the Requirements
for the Degree of
Master of Science
August 2013
The thesis of James A. Deane Jr. was reviewed and approved* by the following:
Maureen D. Feineman
Assistant Professor of Geosciences
Thesis Advisor
David H. Eggler
Emeritus Professor of Petrology
Peter J. Heaney
Professor of Geosciences
James D. Kubicki
Professor of Geoscience
Chris J. Marone
Professor of Geosciences
Associate Department Head of Graduate Programs
*Signatures are on file in the Graduate School
ii
ABSTRACT
Mantle olivines and clinopyroxenes from xenoliths sampled worldwide
commonly display disequilibrium distributions of lithium and its isotopes. While many
theories of open- and closed- system redistribution of lithium have been proposed, none
has been adequately demonstrated experimentally. Closed system redistribution of
lithium between olivine and clinopyroxene has been proposed to occur in response to
changing equilibrium partition coefficients driven by changes in temperature or oxygen
fugacity. Previous experiments (Yakob et al., EPSL. 330, 11-21, 2012) have shown that
there is no temperature dependence on Li partitioning between 700-1100°C. The goal of
this thesis is to determine the potential effect of oxygen fugacity on lithium partitioning
between olivine and clinopyroxene. To this end, a series of piston cylinder experiments
was designed to fix the oxygen fugacity at various mantle values while natural olivine
and diopside starting materials were equilibrated with a Li-bearing aqueous fluid.
Oxygen fugacity was buffered using the solid state fO2 buffers Fayalite-MagnetiteQuartz, Magnetite-Wüstite, and Mo-MoO2. A silver capsule was employed to minimize
the solubility of Fe in the capsule, thus keeping it in the experimental charge. Iron
retention is critical because the mechanism being tested is coupled substitution of Li+ and
Fe3+ for Mg2+ or Fe3+ in the olivine structure.
Results from these experiments show there is no apparent relationship between
oxygen fugacity and equilibrium partitioning of Li between olivine and clinopyroxene at
upper mantle conditions. Over the range of fugacities investigated, DLiol/di was 1.7±0.4
regardless of fO2. The lack of a relationship between Li partitioning and oxygen fugacity,
coupled with the previously determined lack of a relationship with temperature, suggests
that closed system redistribution of Li between mantle phases is unlikely to explain the
isotopic variability we see in xenoliths worldwide. Instead, open system interaction with
metasomatic fluids or melts could be the cause. This implies that Li isotopes in mantle
xenoliths are not necessarily representative of the ambient mantle from which they are
sourced, and therefore limits the use of Li isotopes as a passive tracer of recycled crustal
material in the mantle. However, Li isotopic fractionation during open-system processes
related to xenolith entrainment and eruption is likely to be useful as a geospeedometer to
determine the rates of these processes.
iii
TABLE OF CONTENTS
List of Tables.....................................................................................................................vii
List of Figures...................................................................................................................viii
Acknowledgements.............................................................................................................xi
1
INTRODUCTION...................................................................................................1
1.1
Li in the Earth..............................................................................................1
1.2
Influence of Oxygen Fugacity on Li Partitioning........................................5
2
METHODS............................................................................................................11
2.1
Starting Materials.......................................................................................11
2.2
Experimental Methods...............................................................................11
2.3
Analytical Methods....................................................................................13
2.3.1 Scanning Electron Microscopy..................................................................13
2.3.2 Electron Microscopy..................................................................................14
2.3.3 Laser Ablation Inductively Coupled Mass Spectrometry..........................14
2.3.4 X-ray Diffraction of fO2 buffers................................................................15
3
RESULTS..............................................................................................................20
3.1
Experimental Results.................................................................................20
3.2
Run Product Compositions........................................................................20
3.3
Oxygen Fugacity Control...........................................................................21
3.4
Partition Coefficients.................................................................................21
4
DISCUSSION........................................................................................................35
4.1
Attainment of Equilibrium in Experiments................................................35
4.2
Oxygen Fugacity........................................................................................35
4.2.1 Oxygen Fugacity in the Mantle..................................................................37
4.3
Li Partitioning as a Function of fO2...........................................................38
4.4
Li Variability in Mantle Minerals..............................................................40
4.4.1 Equilibrium Isotope Fractionation.............................................................40
4.4.2 Open vs. Closed System Redistribution of Li............................................41
5
CONCLUSIONS...................................................................................................60
6
6.1
FAILED EXPERIMENTS.....................................................................................62
Methodology..........................................................................................................62
iv
6.2
Experimental Designs............................................................................................62
References..........................................................................................................................65
Appendix: Sample Assembly and Piston Cylinder Procedures.........................................71
v
LIST OF TABLES
Table 2.1.1 Amounts of starting materials and run conditions for all experiments used
in this study........................................................................................................................16
Table 3.2.1 Major element concentrations of starting materials and experimental run
products..............................................................................................................................22
Table 3.2.2 Trace element concentrations of starting materials and experimental run
products..............................................................................................................................23
Table 3.4.1
Calculated Li partition coefficients (DLiol/cpx)............................................24
Table 4.4.1 Compilation of Li data from olivine and clinopyroxene mineral separates
from xenoliths worldwide. Data from Aulbach and Rudnick (2009)1, Aulbach et al.
(2008)2, Ionov and Seitz (2008)3, Jeffcoate et al. (2007)4, Kaliwoda et al. (2008)5, Magna
et al. (2006, 2008)6,7, Ottolini et al. (2004)8, Rudnick and Ionov (2007)9, Seitz and
Woodland (2000)10, Seitz et al. (2004)11, Tang et al. (2007b)12, and Woodland et al.
(2004)13..............................................................................................................................44
Table A.1
Pressure calibration for the piston cylinder apparatus used in this study..78
vi
LIST OF FIGURES
Figure 1.1.1 δ7Li variability in various terrestrial reservoirs. Figure 1 from Tang et al.,
2007a....................................................................................................................................7
Figure 1.1.2 The lithium geochemical cycle. References as follows: Seawater: Chan
and Edmond, 1988; Tomascak et al., 1999; James and Palmer, 2000; Rainwater: Millot et
al. 2010b; Continents: Teng et al., 2004; Mantle: Chan et al., 1992; Tomascak et al.,
2008; Altered Oceanic Crust: Chan et al., 1992, 2002a; Subducted sediment: Chan et al.,
2006; Arc lavas: Chan et al., 2002b; Tomascak et al., 2002; Moriguti et al., 2004;
Serpentinized mantle: DeCitre et al., 2002; Benton et al., 2004; Rivers: Huh et al., 1998;
Millot et al., 2010a; Hydrothermal fluids: Chan et al., 1993, 1994; Subduction reflux:
Zhang et al., 1998; James et al., 1999, Chan and Kastner, 2000; Alteration and sediment
uptake: Chan et al., 1992, 2002a, 2006; Plank and Langmuir, 1998; Tomascak et al.,
2008......................................................................................................................................8
Figure 1.1.3 A compilation of olivine/clinopyroxene lithium partition coefficients (D)
vs. isotopic fractionation factors (α) in mantle xenoliths worldwide. Figure 1 from Yakob
et al., 2012 with data taken from Aulbach and Rudnick (2009), Aulbach et al. (2008),
Ionov and Seitz (2008), Jeffcoate et al. (2007), Kaliwoda et al. (2008), Magna et al.
(2006, 2008), Ottolini et al. (2004), Rudnick and Ionov (2007), Seitz and Woodland
(2000), Seitz et al. (2004), Tang et al. (2007b), and Woodland et al. (2004)......................9
Figure 1.1.4 δ7Li variability between coexisting mantle phases. In most cases, olivine
has significantly higher δ7Li than coexisting clinopyroxene. This inter-mineral variability
remains unexplained. Modified from Figure 7 from Seitz et al. (2004) to include data
from Rudnick and Ionov (2007) and Tang et al. (2007b)..................................................10
Figure 2.2.1 a) Photograph of the piston cylinder apparatus used to conduct the
experiments presented in this thesis. b) Schematic of the piston cylinder showing the
arrangement of parts..........................................................................................................17
Figure 2.2.2 Sample assembly used in the piston cylinder experiments. The figure is to
scale...................................................................................................................................18
Figure 2.2.3 Schematic diagram of the double capsule experimental setup using two Ag
capsules separated by a 50/50 Ag-Pd foil. Schematic diagram shows the theoretical
progression of the experiment through time. All experiments were run at 900°C and 1
GPa.....................................................................................................................................19
Figure 3.1.1 Backscattered electron SEM images of run products from experiment TJ67. These crystals have not been mounted in epoxy and polished; rather they have been
mounted on electrically conductive tape to image via SEM and show morphology of
vii
unpolished crystals. Image A shows multiple olivine crystals surrounded by quench
material. Image B shows one large elongate clinopyroxene.............................................25
Figure 3.1.2 SEM backscattered electron image of run products from experiment TJ-52
buffered at Mo-MoO2. Sample was mounted in epoxy and polished with 600 grit siliconcarbide sandpaper and 1µm diamond paste. Clinopyroxene is typically more elongate
and brighter than olivine in backscattered electron images...............................................26
Figure 3.1.3 SEM backscattered electron image of run products from experiment TJ-63
buffered at Magnetite-Wüstite. Sample was mounted in epoxy and polished with 600 grit
silicon-carbide sandpaper and 1µm diamond paste. Clinopyroxene is typically more
elongate and brighter than olivine in backscattered electron images................................27
Figure 3.1.4 SEM backscattered electron image of run products from experiment TJ-67
buffered at FMQ. Sample was mounted in epoxy and polished with 600 grit siliconcarbide sandpaper and 1µm diamond paste. Clinopyroxene is typically more elongate
and brighter than olivine in backscattered electron images...............................................28
Figure 3.1.5 SEM backscattered electron image of run products from experiment TJ-68
which was unbuffered. Sample was mounted in epoxy and polished with 600 grit siliconcarbide sandpaper and 1µm diamond paste. Clinopyroxene is typically more elongate
and brighter than olivine in backscattered electron images...............................................29
Figure 3.2.1 Typical time resolved laser ablation spectrum for standard BCR-2g used in
sample-standard bracketing. Black vertical lines indicate the background differentiated
from the analysis signal. LA-ICP-MS data was reduced with LasyBoy V3.01, a data
reduction spreadsheet created by Joel Sparks at Boston University..................................30
Figure 3.3.1 XRD analysis of the fO2 buffer from TJ-52 is shown in black. The red
dotted line represents an XRD spectrum for Mo and the blue dashed line represents an
XRD spectrum for Mo-MoO2. TJ-52 shows clear peaks that overlap well with both
phases, suggesting that the experimental buffer was successful........................................31
Figure 3.3.2 XRD analysis of the fO2 buffer from TJ-63 is shown in black. The red
dashed line represents an XRD spectrum for magnetite. There is no confirmed wüstite in
this buffer, but the broad increase in intensity from ~30-45° represents data from an
amorphous phase with no defined crystal structure. This is being interpreted as the
remnants of the wüstite present at run conditions which decomposed upon quenching as
wüstite is not stable at earth surface conditions.................................................................32
Figure 3.3.3 XRD analysis of the fO2 buffer from TJ-67 is shown in black. The red
dashed line represents an XRD spectrum for fayalite. The blue dash-dot line represents
an XRD spectrum for magnetite. The green dotted line represents an XRD spectrum for
viii
quartz. While the strongest peaks in TJ-67 overlap with magnetite, there are overlapping
peaks that only correspond to quartz and fayalite as well, suggesting that all three phases
are present at the conclusion of the experiment and the buffer was successful.................33
Figure 3.4.1 Experimental results showing the measured Li partition coefficient
between olivine and clinopyroxene as a function of fO2. Open symbol represents
equilibrium partition coefficient from 3-day 900°C experiment presented by Yakob et al.,
2012....................................................................................................................................34
Figure 4.1.1 Time series of piston cylinder equilibrium partitioning experiments from
Yakob et al. (2012) showing the DLiol/cpx evolving over time. At 900°C the system
approaches equilibrium with respect to Li partitioning between olivine and clinopyroxene
within 24 hours. Figure 4 from Yakob et al., 2012..........................................................55
Figure 4.2.1 Solid state oxygen fugacity buffers rhenium-rhenium dioxide (Re-ReO2),
fayalite-magnetite-quartz (FMQ), magnetite-wüstite (MW), and molybdenummolybdenum dioxide (Mo-MoO2) shown as a function of temperature. Vertical black
line indicates experimental run conditions. Buffer curves calculated from: Mo-MoO2 =
O’Neill, 1986; Magnetite-Wüstite & Fayalite-Magnetite-Quartz = Myers and Eugster,
1983....................................................................................................................................56
Figure 4.4.1 Experimentally derived DLiol/cpx showing that the olivine/clinopyroxene
partition coefficient is ~2±0.5 regardless of temperature. Figure 6 from Yakob et al.,
2012. Data from Ottolini et al. (2009), Caciagli et al. (2011), Brenan et al. (1998), Yakob
et al. (2012) and Blundy and Dalton (2000)......................................................................57
Figure 4.4.2 Northrop and Clayton (1966) state that a linear relationship should exist
between ln(α)*103 vs. T-2*106 if isotopic fractionation exists as an equilibrium
phenomenon. Plotting ln(α)*103 vs. T-2*106 for all available xenolith data shows that in
most cases there is no observable relationship, and therefore the isotopic fractionation
seen in xenoliths is not the product of equilibrium isotope fractionation. A particular
suite from Rudnick and Ionov (2007), however, do seem to fit to a linear trend suggesting
a potential relationship between isotopic fractionation and temperature. Data from Table
4.4.1....................................................................................................................................58
Figure 4.4.3 Apparent Li partition coefficients measured from mantle xenolith mineral
separates worldwide show no relationship with temperature. Data from Table 4.4.1.......59
ix
ACKNOWLEDGEMENTS
The successes and failures of this research project proved to be huge test of
patience and dedication to succeed, and I couldn’t have completed this project without
the help from many wonderful people. First, I would like to thank my advisor, Dr.
Maureen Feineman. She has always been supportive and helpful with matters both
related and not to this project. Many experimental failures led to us talking about what to
do next, and while we got frustrated, she never allowed us to turn our backs on this
project. She has been a wonderful advisor and a great mentor for me to talk to about my
past, present and future. I would also like to thank the members of my thesis committee
who’ve helped me learn to think about scientific problems from many different angles.
I’ve learned that focusing on one aspect of science is incredibly limiting, and true
productivity can only happen when thinking about problems in many different ways. I’d
like to specifically thank Dr. David Eggler, whose experimental expertise was
particularly helpful. Whenever there were experimental failures or mechanical
breakdowns, I would call Dave and he would come in and help me fix the machines or
redesign our experimental approach.
I’d like to thank Melanie Saffer, Nichole Wonderling and Mike Ackerson for their
help with the geochemical analyses performed in this study. I’d also like to thank Bruce
Watson, Dustin Trail and Jay Thomas for their help in designing the experimental setup
used in this study.
I would especially like to thank all of my friends and family, specifically my mom
and dad, my brothers Ryan and Brandon, and my lovely wife-to-be Camille, who have all
been incredibly supportive throughout the process of completing this research project.
Their guidance and encouragement has truly been an inspiration in keeping me focused
on finishing this project and completing graduate school. Thank you
x
1.
INTRODUCTION
1.1
Lithium in the Earth
Lithium is a trace element found both on the surface and within the interior of the
Earth. Because Li is very fluid mobile and found within oceanic basalts and sediments in
isotopic ratios that contrast with those inherent to the mantle (Chan and Edmond 1988;
Chan et al., 1994; Elderfield and Schultz 1996), there has been great hope over the past
two decades that lithium, and in particular the isotopes of Li, can serve as a geochemical
tracer of surface material recycled back into the mantle (Elliott et al., 2004; Seitz and
Woodland, 2000; Kent and Rossman, 2002). Research into lithium behavior in the
mantle has been an evolving field of geochemistry for the past 20 years, and there is still
much to learn about how lithium and its isotopes behave in Earth systems. Lithium
concentrations and isotopic ratios have been measured in whole rock samples, mineral
separates, and analyses of intra-mineral zoning from all around the world (e.g. Chan et
al., 1992; Chan et al., 1994; Coogan et al., 2005; Jeffcoate et al., 2007; Elliott et al., 2004;
Zack et al., 2003; etc.), but they have not yet led to a full understanding of how Li
behaves chemically in various environments, for example during recycling from the
Earth’s surface into the mantle at subduction zones.
Lithium is a light trace element with two stable isotopes, 6Li and 7Li (7.5% and
92.5% abundances, respectively). There is ~15% difference in mass between the two
isotopes, which is relatively large, and 6Li and 7Li are strongly fractionated by mass
dependent processes at low temperatures (Fritz and Whitworth, 1994; Pistiner and
Henderson, 2003; Huh et al., 1998; Huh et al., 2001). Lithium is a very fluid mobile
element and therefore can be mobilized and isotopically fractionated by weathering
processes at the Earth’s surface (Pistiner and Henderson, 2003; Rudnick et al., 2004;
Millot et al., 2010), circulation of hydrothermal fluids within spreading centers or
volcanic systems (Chan et al., 1993), or during dehydration/hydration reactions within
subduction zones (Chan et al., 2002; DeCitre et al., 2002; Zack et al., 2003; Leeman et
al., 2004). The fluid mobility of Li contributes to its potential as a geochemical tracer for
mass transfer within subduction zones. Lithium incorporated in altered oceanic crust and
subducted sediment is released and carried by slab-derived fluids back up through the
1
overlying mantle wedge and into volcanic arcs (Chan et al. 1992; Leeman et al. 2004)
and can be incorporated into mantle minerals such as olivine and pyroxene (Seitz and
Woodland, 2000; Kent and Rossman, 2002; Rudnick and Ionov, 2007). This process
should lead to a distinct chemical and isotopic signature in magmas and mantle
peridotites derived from volcanic arcs (Elliott et al., 2004).
Richter et al. (2003) showed that Li diffusion in silicate melt is 1-3 orders of
magnitude faster than diffusion of almost all other major and trace elements. The only
other species with comparable diffusion coefficients are light noble gases (He, Ne) and
water. Because diffusion rates are mass-dependent, and because the two isotopes of
lithium have such a large relative mass difference, significant isotope fractionation may
occur during diffusion. This can result in Li isotopic fractionation of up to 10-25‰ in
natural and experimental systems (Richter et al. 2003). Combining these two
observations, rapid diffusion could cause Li to be highly fractionated on short timescales
within minerals and rocks in natural samples. Thus Li zoning within phenocrysts or
between minerals in mantle xenoliths has the potential to tell us how long a particular
crystal took to cool below the closure temperature of Li diffusion (Coogan et al., 2005;
Jeffcoate et al., 2007; Parkinson et al., 2007).
Isotopically, Li is reported in terms of δ7Li [(7Li/ 6Lisample / 7Li/6LiL-SVEC − 1)103],
and expressed in permil (‰) variation relative to the L-SVEC lithium carbonate standard
reference material. The value of δ7Li varies widely in different Earth reservoirs (Fig.
1.1.1, from Tang et al., 2007a). Seawater, for example, has a very uniform δ7Li of ~
+33‰ (Chan and Edmond, 1988; Tomascak et al., 1999; James & Palmer, 2000) while
unaltered Mid-Ocean Ridge Basalt (MORB), representative of Earth’s upper mantle has
δ7Li values ranging from +2‰ to +7‰ (Chan et al., 1992; Seitz et al., 2004). Figure
1.1.2 shows the variability of Li isotopic signatures throughout the Earth’s hydrologic
cycle, the crust and the upper mantle, and how these various reservoirs interact with each
other.
Many mantle xenoliths worldwide (Table 1.1.1; Figure 1.1.3) display expected
“equilibrium” behavior with respect to Li distribution and isotope ratio between olivine
and diopside. Nonetheless, there are many mantle xenolith mineral separates (e.g., Seitz
2
et al., 2004; Rudnick and Ionov, 2007; Tang et al., 2007b; Fig. 1.1.4) that demonstrate
measurably different Li isotopic compositions. In the mantle xenoliths that display the
most extreme isotopic difference between olivine and clinopyroxene, the partition
coefficient (DLiol/di) decreases to values below 0.5 while the isotopic fractionation factor
(ol/di) increases to values greater than 1.020, signifying that clinopyroxene has a greater
Li concentration and lower δ7Li relative to olivine, representing an increase in 6Li in
clinopyroxene relative to olivine (e.g., Rudnick and Ionov, 2007; Tang et al., 2007b; Fig.
1.1.3).
Experiments have shown that the equilibrium partition coefficient for Li between
olivine and clinopyroxene (D = [Li]ol/[Li]cpx) is between 1.5 and 2 (Yakob et al., 2012;
Brenan et al., 1998; Blundy and Dalton 2000; Caciagli et al., 2011; Ottolini et al., 2009).
At the high temperatures and pressures characteristic of the mantle, isotope fractionation
between phases (e.g., olivine and diopside) at equilibrium is expected to be minimal.
Equilibrium isotopic fractionation between two phases at high temperature (>200°C) has
been shown to have a linear relationship between 1000*ln(α) (where α = isotopic
fractionation factor, i.e.: (7Li/6Li)ol/(7Li/6Li)cpx) and 1/T2 (e.g., Northrop and Clayton,
1966). The lack of any linear relationship with temperature in worldwide mantle
xenoliths suggests that isotopic variability does not represent equilibrium fractionation
(Discussed in more detail in section 4; Table 4.1.1, Figures 4.1.1, 4.1.2). Also, Tomascak
et al. (1999) showed that as the Kilauea Iki lava lake cooled, the fractional crystallization
of olivine, clinopyroxene, and plagioclase resulted in no measurable isotopic
fractionation of Li. Together, these observations suggest that at high temperature there is
little Li isotopic fractionation between olivine and clinopyroxene.
While it has been clearly shown that Li isotopic variability exists between olivine,
clinopyroxene and orthopyroxene within individual mantle peridotite samples (Jeffcoate
et al., 2007; Ionov and Seitz, 2008; Rudnick and Ionov 2007; Seitz et al., 2004; Aulbach
and Rudnick 2009; Tang et al., 2007; Magna et al., 2006), experimental studies have not
yet been able to confirm a mechanism that controls the variable Li distribution between
coexisting phases. In the most general sense, disequilibrium behavior of Li could be the
result of either closed-system or open-system processes.
3
1. Closed-system processes. In this category of hypotheses, the distribution
of Li is dependent on some condition that changes prior to or during
exhumation. For example, the Li partition coefficient between olivine and
clinopyroxene could be temperature dependent. A xenolith or peridotite
massif cools as it is exhumed. This change in temperature influences the
partition coefficient such that Li becomes more compatible in
clinopyroxene relative to olivine, resulting in Li moving from the olivine
to the clinopyroxene (thus decreasing the Dol/cpx) via diffusion. As has
been shown by Richter et al. (2003), 6Li diffuses faster than 7Li, so this
process of diffusion could cause dynamic isotopic fractionation on the
timescale of diffusion. The fractionation could then be locked into the
minerals as they continue to cool beyond the Li closure temperature.
Other variables, such as changes in oxygen fugacity, might likewise drive
closed-system redistribution of lithium and its isotopes.
2. Open-system processes. In this school of interpretation, lithium from an
external source infiltrates the mantle materials prior to or during
exhumation. Many mantle rocks, including all xenoliths, interact with a
metasomatic fluid or melt during their rise to the surface. This fluid or
melt may have a distinctly different Li isotopic signature from the ambient
mantle peridotites, and it may interact with each mantle mineral
differently. For instance, if Li diffuses from the fluid or melt into the
clinopyroxene faster than it diffuses into the olivine, different δ7Li values
may be generated in clinopyroxene relative to olivine.
The bottom line is that we do not yet have all the tools necessary to fully
understand the large variability in measured Li concentrations and δ7Li values from
mantle xenoliths worldwide. What is needed are experimental values of the partition
coefficients, diffusion coefficients, and isotopic fractionation factors for Li between
mantle phases, and an understanding of how these values are affected by environmental
and geochemical factors so we can put together a model to help explain what we see in
natural samples.
4
1.2 Influence of Oxygen Fugacity on Li Partitioning
It has been proposed that Li partitioning between olivine and diopside is
controlled by temperature, such that Li becomes more compatible in clinopyroxene
relative to olivine as cooling proceeds. However, experimental evidence shows that
DLiol/di remains close to 2 over a range of temperatures between 700˚C and 1400˚C
(Yakob et al., 2012; Caciagli et al., 2011). Below 700˚C, diffusion of Li in
clinopyroxene is too slow to accommodate dynamic redistribution (Coogan et al., 2005).
Another variable that may control the Li partitioning between mantle phases is oxygen
fugacity. The dominant species of iron in olivine is Fe2+, forming a solid solution with
Mg2+ on the octahedral M1 and M2 sites where there is a random distribution of Fe atoms
between the two sites (Spandler and O’Neill, 2010). It is possible, however, that under
oxidizing conditions Fe3+ could serve as a coupled substitution partner with Li+. Coupled
substitutions can preserve overall charge balance in the system, provided that the ionic
radii of the substituting species are near enough to that of the main structural ion to not
distort the crystal structure. Previous studies have shown that an increase in trivalent
cations within the olivine structure can have a significant effect on the partitioning of
monovalent cations (Grant and Wood, 2010). For example, an increase in Sc+3 within the
system allows for increased Li within experimental olivines (Grant and Wood, 2010). In
this case, Fe3+ and Li+ could substitute together into two M2 sites within the crystal
(replacing Mg2+ or Fe2+), preserving the overall charge of the coupled sites (+4). Should
more Fe3+ be present due to higher oxygen fugacity in the system, it is possible that more
Li+ could be preferentially incorporated into the olivine. Although this is not the only
mechanism by which Li can be incorporated into the mineral structures, it is hypothesized
that fO2 could therefore play a role in in the partitioning of Li between coexisting olivine
and clinopyroxene at mantle conditions. In a similar manner to the temperaturedependent partitioning hypothesis, if the partition coefficient is fO2 dependent, changing
environmental fO2 conditions prior to or during exhumation could drive diffusion of
lithium between olivine and diopside, resulting in kinetically induced isotopic
fractionation. It has been proposed that the fO2 effect would be more significant in olivine
relative to clinopyroxene, due to the greater Fe content in olivine (Caciagli, 2010).
Should Li incorporation into mantle minerals scale directly with Fe+3 content, and
5
assuming the Fe+3/Fe+2 is controlled within the minerals by the oxygen fugacity, then
olivine would be more strongly affected than clinopyroxene due to the respective iron
contents. Oxygen fugacity effects on equilibrium partitioning of Li have not been tested
experimentally for the olivine-clinopyroxene system. The goal of this research is to
measure equilibrium distribution of Li between olivine and diopside over a range of
oxygen fugacity values relevant to the upper mantle.
6
Figure 1.1.1. δ7Li variability in various terrestrial reservoirs. Figure 1 from Tang et al.,
2007a
7
Figure 1.1.2. The lithium geochemical cycle. References as follows: Seawater: Chan and
Edmond, 1988; Tomascak et al., 1999; James and Palmer, 2000; Rainwater: Millot et al.
2010b; Continents: Teng et al., 2004; Mantle: Chan et al., 1992; Tomascak et al., 2008;
Altered Oceanic Crust: Chan et al., 1992, 2002a; Subducted sediment: Chan et al., 2006;
Arc lavas: Chan et al., 2002b; Tomascak et al., 2002; Moriguti et al., 2004; Serpentinized
mantle: DeCitre et al., 2002; Benton et al., 2004; Rivers: Huh et al., 1998; Millot et al.,
2010a; Hydrothermal fluids: Chan et al., 1993, 1994; Subduction reflux: Zhang et al.,
1998; James et al., 1999, Chan and Kastner, 2000; Alteration and sediment uptake: Chan
et al., 1992, 2002a, 2006; Plank and Langmuir, 1998; Tomascak et al., 2008.
8
Figure 1.1.3. A compilation of olivine/clinopyroxene lithium partition coefficients (D)
vs. isotopic fractionation factors (α) in mantle xenoliths worldwide. Figure 1 from Yakob
et al., 2012 with data taken from Aulbach and Rudnick (2009), Aulbach et al. (2008),
Ionov and Seitz (2008), Jeffcoate et al. (2007), Kaliwoda et al. (2008), Magna et al.
(2006, 2008), Ottolini et al. (2004), Rudnick and Ionov (2007), Seitz and Woodland
(2000), Seitz et al. (2004), Tang et al. (2007b), and Woodland et al. (2004).
9
15.0
- Seitz et al. 2004
- Tang et al. 2007
- Rudnick and Ionov 2007
10.0
δ7Li
5.0
0.0
-5.0
-10.0
-15.0
olivine
clinopyroxene
Figure 1.1.4. δ7Li variability between coexisting mantle phases. In most cases, olivine
has significantly higher δ7Li than coexisting clinopyroxene. This inter-mineral variability
remains unexplained. Modified from Figure 7 from Seitz et al. (2004) to include data
from Rudnick and Ionov (2007) and Tang et al. (2007b).
10
2.
METHODS
2.1
Starting Materials
Natural crystals of forsteritic olivine (Mg1.8Fe0.2SiO4), from San Carlos, Arizona,
and diopside (CaMgSi2O6) from DeKalb, New York, were used in this study as starting
materials for the growth of olivine and clinopyroxene in the piston cylinder experiments
discussed below. Crystals were broken with a hardened steel crusher. Clean, inclusion
free, alteration free pieces were separated by hand under a stereoscopic microscope and
cleaned with ultrapure Milli-Q water. These pieces were then powdered with an agate
mortar and pestle while submerged in ethanol. Starting compositions for all materials
used in this study can be found in Tables 3.2.1 and 3.2.2. Natural quartz (SiO2) from
Brazil and synthetic anorthite (CaAl2Si2O8) were added to stabilize clinopyroxene in the
run products by saturating the fluid with Si, Ca and Al. The mass of powders added to
each experiment and oxygen fugacity buffer can be found in Table 2.1.1. The materials
used to buffer the oxygen fugacity within the experiments include synthetic fayalite
(Fe2SiO4), magnetite (Fe3O4), quartz (SiO2), wüstite (FeO), Mo (metal), and MoO2.
2.2
Experimental Methods
An end-loaded piston cylinder (PC) apparatus in the High-Pressure Laboratory at
the Pennsylvania State University was used to generate upper mantle pressure and
temperature conditions in which to run the equilibrium partitioning experiments (Figure
2.2.1). The piston cylinder is made up of two large hydraulic rams; the end load ram
controls the confining pressure, and the master ram controls the pressure on the sample
assembly. The master ram is 9 inches in diameter. By using the relationship P = F/A,
where P is pressure, F is force, and A is the area over which the force is applied, we can
generate high pressures by stepping the applied force of the 9” ram down to a smaller
area – in this case a ¾” tungsten carbide (WC) piston. The pressures generated in the
sample assembly can be as much as 4-5 GPa, depending on the piston diameter used.
The confining pressure is applied by a 10” diameter pressure plate, which is a series of
two concentric tapered steel rings around a tungsten carbide core. When the end load
11
ram applies pressure to the face of the plate, that pressure is transferred laterally to the
WC core in order to provide confining support for the sample assembly. The purpose of
the confining pressure provided by the end load ram and pressure plate is to prevent
lateral failure of the sample cell when high pressures are applied on the piston and sample
assembly by the master ram.
The experimental assembly (Figure 2.2.2) and method are adapted from Ayers et
al. (1992) and Trail et al. (2012). The capsule is designed so one Ag capsule is inverted
and set on top of a second Ag capsule, separated by a thin Ag-Pd foil. Silver capsules are
used in these experiments because Ag does not alloy with Fe, such that mantle-like Mg#s
(Mg/(Mg+Fe)) can be maintained in the run product olivines (~Fo90). Retention of Fe in
the system is imperative, as the purpose of these experiments is to determine the effect of
Fe3+ on Li partitioning in olivine. The lower capsule contains the experimental starting
materials. The experimental charge consists of powdered San Carlos Olivine (~7.5 mg),
DeKalb Diopside (~5 mg), synthetic anorthite (~1 mg) and natural Brazilian quartz (~1
mg) plus an aqueous solution containing 200 ppm Li and 100 ppm Ba in milli-Q water
(~50 mg). The lower capsule is then loosely capped with the Ag-Pd foil. In the upper
capsule the solid state fO2 buffer is mixed with water. The surface tension of the water in
the relatively small capsule prevents the powders and fluid from escaping the capsule
when it is inverted and placed on top of the Ag-Pd foil. Upon initial pressurization, the
foil seals to the rim of the upper and lower capsules, preventing any fluid escape when
the experiment is brought to run conditions.
The solid-state fO2 buffers used in these experiments, from highest to lowest fO2,
are fayalite-magnetite-quartz, magnetite-wüstite, and molybdenum-molybdenum dioxide.
At 900°C, these oxygen fugacity buffers correspond to values of logfO2 of -12.5, -15 and
-17 respectively. The range of fO2 values prescribed by these buffers spans the range of
oxygen fugacities expected in the Earth’s mantle. The Ag-Pd foil is used to physically
separate the solid state buffer assemblage from the experimental charge. Water is present
on both sides of the foil, which is permeable to H2. This allows H2 formed by the
dissociation of H2O during the run to move between chambers. If the buffer side of the
capsule contains more free oxygen than the olivine-diopside capsule at run conditions,
H2O on the experiment side will dissociate and H2 will diffuse through the foil to the
12
buffer side. If the buffer side contains less free oxygen, H2O in the buffer capsule will
dissociate and H2 will diffuse through the foil to the experiment side (Figure 2.2.3). As
long as H2O is present on both sides of the foil and all components of the solid-state
buffer remain at the end of the experiment, then both sides of the capsule were fixed at
the fO2 given by the buffer for the run conditions of interest.
The sample capsule is placed inside a NaCl-pyrex assembly with a straight-walled
graphite furnace and loaded into the piston cylinder apparatus. The capsule is initially
“cold-sealed” at ~0.8 GPa, then the experiment is brought up to run conditions. All
experiments were held at 1.0 GPa and 900˚C for ~72 hours. Experiments were quenched
by manually cutting the power to the furnace while maintaining the flow of water through
the cooling plates and coil; temperatures dropped below 300˚C in ~30 seconds. Detailed
experimental procedures are given in Appendix 1. Subsequent to quenching, the cleaned
sample capsule was weighed to ensure that no water was lost during the experiment.
Upon piercing the capsule, water should bubble out if all the initial water was retained
during the run. After opening, the capsule was left under a heat lamp overnight to
evaporate the remaining water, leaving behind only solid material. This solid material
was collected, weighed, cleaned in an ultrasonic bath, and labeled and stored in glass
vials. An aliquot was mounted in epoxy in a ¼” stainless steel annulus. The mounts
were then polished using 600 grit silicon carbide sand paper and 1 micron diamond paste
in preparation for analysis.
2.3
Analytical Methods
2.3.1
Scanning Electron Microscope
Experimental run products are initially analyzed on a FEI Quanta 200
Environmental Scanning Electron Microscope (ESEM) at the Materials Characterization
Laboratory at Penn State University. Backscattered electron (BSE) images are captured
to be used as maps for later analyses on other analytical instruments. Because the relative
brightness of grains in a BSE image is a function of the average atomic mass of the
material, olivine and clinopyroxene grains can usually be distinguished from each other
13
visibly. Individual energy dispersive x-ray spectroscopy (EDS) spot analyses are
performed on specific grains of interest to obtain a semi-quantitative understanding of the
composition of the material. Raw spectra were collected and converted to concentrations
using the Oxford Instruments INCA software package. This helped to ensure that both
Mg-rich olivine and diopside were successfully grown in each experiment.
2.3.2
Electron Microprobe Analysis
Fully quantitative major element compositions of the run products were
determined using a Cameca SX-100 Electron Probe Micro-analyzer (EPMA) at
Rensselaer Polytechnic Institute. Polished samples were carbon coated and analyzed for
Al, Ca, Cr, Fe, Mg, Mn, Ni, Si, and Ti using a focused electron beam with a current of 20
nA and an accelerating voltage of 15 kV. Natural diopside, tephroite, and kyanite, as
well as synthetic fayalite and forsterite were used as standards, and x-ray peak intensities
were converted to concentrations using the CIT-ZAF correction method.
2.3.3
Laser Ablation Inductively Coupled Plasma Mass Spectrometry
Trace element abundances were determined by laser ablation inductively coupled
plasma mass spectrometry (LA-ICP-MS) using a Thermo X-series II Quadrupole ICPMS connected to a New Wave UP-213 (213 nm) Neodymium-Yttrium-Aluminum-Garnet
(Nd-YAG) laser ablation system at the Laboratory for Isotopes and Metals in the
Environment (LIME) facility within the Penn State Institutes of Energy and the
Environment. A 30 µm diameter spot size was used to ensure that the laser analysis
sampled only the interior of the ~50 µm diameter crystals, and not any surface material.
Crystals were ablated with an output energy of 5-6 J/cm2 and a frequency of 20 Hz.
Ablated material was carried by a stream of helium (He) gas, subsequently mixed with
argon (Ar) gas before entering the Ar-plasma torch of the ICP-MS.
Batch analyses of 15 samples were bracketed by standard analyses to correct for
internal mass drift within the ICP-MS. Standards used in these analyses were USGS
basalt glass standards BCR-2g and GSD-1g. The isotopes measured were 7Li, 25Mg,
14
26
Mg, 43Ca, 47Ti, 60Ni, and 137Ba. Raw counts were converted to concentrations using
LasyBoy V3.01, a data reduction spreadsheet created by Joel Sparks at Boston
University. Calibration lines were constructed from selected bracketing standards for
each set of analyses, and elemental concentrations were calculated from backgroundcorrected counts-per-second of each isotope using 25Mg as a normalizing mass.
To ensure that no surface material or fluid inclusions were incorporated during the
analysis, barium (Ba) was added to the starting fluid. Barium is incompatible in the
silicate phases in the experiments, and should remain in the fluid. Therefore, any Ba
detected in the laser ablation analysis would have to originate from quenched material
adhered to the surface of the crystals or fluid inclusions within individual crystals. Any
analysis with Ba above detection limits was excluded from the results.
2.3.4
X-ray Diffraction of fO2 Buffers
X-ray diffraction (XRD) analyses were performed on the post-
experiment fO2 buffers to ensure that all solid phases were present at the conclusion of
the experiment. The XRD analyses were performed by Nichole Wonderling at the
Materials Characterization Lab at Penn State University using a PANalytical X’Pert Pro
MPD Theta-Theta Diffractometer system with Cu K-alpha radiation (1.54059 Å). Phase
identification was performed using Materials Data Incorporated (MDI) Jade 2010
software via comparison to the International Center for Diffraction Data (ICDDPDF) database.
15
Table 2.1.1
TJ-52
TJ-63
TJ-67
TJ-68
Amounts of starting materials and run conditions for all experiments used in this study.
Experimental starting materials (mg)
San Carlos DeKalb
fO2Buffer
Olivine
Diopside CaAl2Si2O8
Mo-MoO2
189.5
133.5
--MW
7.7
4.8
1.1
FMQ
8.5
6.9
1.2
Unbuffered
8.0
5.6
1.0
SiO2
--0.8
1.0
0.9
Li-Ba
fluid
12.0
41.1
67.7
60.7
fO2 Buffers (mg)
MoFeOMoO2 Fe3O4
56.6
----11.4
---------
SiO2
----4.1
---
Fe2SiO4
----4.8
---
Fe3O4
----4.5
---
Li-Ba
fluid
18.4
38.9
39.5
---
--- = not used for this experiment
Mo-MoO2 mixed together in 4:1 ratio by mass
FeO-Fe3O4 mixed together in 3:1 ratio by mass
All experiments run at 900°C and 1 GPa
16
a
b
Figure 2.2.1. a) Photograph of the piston cylinder apparatus used to conduct the
experiments presented in this thesis. b) Schematic of the piston cylinder showing the
arrangement of parts
17
Ag-Pd
foil
Ag Capsules
Ceramic
Annulus
and Lid
1
mm
3.9 mm
6.4 mm
6.4 mm
9.8 mm
44
mm
9.9 mm
2 mm
11.5 mm
Solid
Ceramic
15
mm
10
mm
10 mm
Graphite
Sleeve
and Plug
42
mm
Center
Drilled
Ceramic
1.7 mm
9.8 mm
44
mm
11.6 mm
13.9 mm
Pyrex Sleeve
Base
Plug
12.7
mm
16
mm
1.7 mm
15.8 mm
9.8 mm
47
mm
Pyrophyllite
Sleeve
12.7
mm
16 mm
19 mm
44
mm
14 mm
1.6 mm
Thermocouple
18.8 mm
Insulator
NaCl Sleeve
58 mm
Lead Foil
Figure 2.2.2. Sample assembly used in the piston cylinder experiments. The figure is to scale.
18
1
Mineral powders and
fluid sealed in capsule
2
Minerals grow incorporating Li;
fO2 buffer reaction progresses
2Fe3O4 + 3SiO2
M
F
Q
Buffer powders
Ag-Pd
foil
Experiment powders
O2
H2
O2
O2
Li
3Fe2SiO4 + O2
Magnetite and quartz react to form
fayalite and free oxygen (or vice
versa). H 2 diffuses through the
Ag-Pd foil to create water from
excess free O 2, balancing the
activity of O 2 in both capsules.
H2
Li
Li
Silver capsule
Figure 2.2.3. Schematic diagram of the double capsule experimental setup using two Ag
capsules separated by a 50/50 Ag-Pd foil. Schematic diagram shows the theoretical progression
of the experiment through time. All experiments run at 900°C and 1 GPa.
19
3.
RESULTS
3.1
Experimental Results
Successful experiments were characterized by the growth of both olivine and
clinopyroxene 50 m in diameter, as well as retention of the fluid added at the
beginning of the experiment. Within the oxygen fugacity buffer chamber, all
components, including water, must be present at the conclusion of the experiment to be
considered a successfully buffered experiment. The two chambers needed to remain
physically isolated (that is, the Ag-Pd foil must have remained intact) for the experiment
to be considered a success.
Crystals were mounted in epoxy and analyzed using scanning electron
microscopy (SEM) to identify the phases that were grown in the experiment. X-ray
energy dispersive spectroscopy (EDS) analysis confirmed that both olivine and
clinopyroxene were present in each experiment. Backscattered electron (BSE) images
were taken and used as maps to help identify phases and show relative crustal sizes
during subsequent analysis on the LA-ICP-MS (Figures 3.1.1- 3.1.5). Typical
experimental crystal sizes ranged from 30-100 µm, with olivine typically growing larger
than clinopyroxene.
3.2
Run Product Compositions
Major element compositions of the run products and starting materials are
presented in Table 3.2.1. A representation of a typical LA-ICP-MS spectrum showing
the collection of a background and signal is shown in Figure 3.2.1. The olivine run
products contain ~10 wt% FeO* (all iron calculated as FeO), confirming that the Ag
capsule successfully preserved Fe within the system. Trace element (Li and Ba)
compositions of the run products are presented in Table 3.2.2. Barium content was used
to monitor for unintentional sampling of fluid inclusions or grain boundaries during
analysis, and data were filtered to include only those samples with Ba content below
detection limits. The Ba content of clinopyroxenes in experiment TJ-52 was above
detection limits in all analyses. In this case, clinopyroxenes with < 10 ppm Ba were
20
included. Lithium contents ranged from ~10 to ~27 ppm in olivine, and from ~6 to ~13
ppm in diopside.
3.3
Oxygen Fugacity Control
A solid-state oxygen fugacity buffer consumes “reactants” to generate “products”,
including O2. These are reversible reactions, and can run in either direction to maintain
fO2 in the capsule, such that the designation of which components are the reactants and
which are the products is arbitrary. In order to confirm that the fO2 in the capsule was
maintained at the buffer value for the given run conditions (shown in Fig. 2.2.4),
however, it is critical that all reactants and products, including H2O, be present at the end
of the experiment. The solid-state buffer materials were analyzed via x-ray diffraction
(XRD) at the Materials Characterization Laboratory at Penn State University to confirm
that all solid phases were indeed present at the end of the experiment. The results from
the x-ray analyses showed that all necessary phases were present. TJ-52 showed peaks
for Mo-metal and MoO2. TJ-63 showed major peaks for magnetite and an amorphous
phase. Throughout the quenching process, wüstite becomes unstable and likely reverted
to magnetite and a poorly-crystalline Fe phase. TJ-67 showed peaks for fayalite,
magnetite and quartz. X-ray spectra from these buffers are presented in Figures 3.3.13.3.3.
3.4
Partition Coefficients
The equilibrium partition coefficient (DLiol/cpx) is calculated by dividing the
average Li concentration in the olivine by the average Li concentration in the diopside in
each experiment. This calculation yields DLiol/cpx = 1.7±0.4 for all experiments (Table
3.4.1 and Figure 3.4.1). Partition coefficients are plotted against log fO2 in Figure 3.4.1.
21
Table 3.2.1. Major element concentrations of starting materials and experimental run products.
fO2 buffer
Mo-MoO2
Olivine
MW
FMQ
Unbuffered
Mo-MoO2
Clinopyroxene
MW
FMQ
Unbuffered
TJ-52
n = 10
TJ-63
n=7
TJ-67
n = 10
TJ-68
n=7
San Carlos
Olivine*
n = 17
TJ-52
n=9
TJ-63
n=7
TJ-67
n = 10
TJ-68
n=9
DeKalb
Diopside*
n = 14
49.67 (1.30)
41.47 (0.29)
40.32 (0.98)
40.76 (1.02)
40.91 (0.59)
54.63 (1.02)
54.99 (0.88)
54.63 (1.10)
54.88 (0.77)
54.54 (0.90)
TiO2
0.00 (0.00)
0.00 (0.01)
0.00 (0.02)
0.00 (0.02)
0.00 (0.02)
0.01 (0.02)
0.02 (0.02)
0.02 (0.01)
0.02 (0.02)
0.02 (0.02)
Al2O3
0.01 (0.02)
0.01 (0.02)
0.02 (0.04)
0.02 (0.03)
0.02 (0.03)
0.54 (0.22)
0.57 (0.19)
0.58 (0.06)
0.61 (0.10)
0.81 (1.18)
Cr2O3
0.02 (0.04)
0.02 (0.02)
0.02 (0.03)
0.02 (0.03)
0.02 (0.02)
0.01 (0.02)
0.00 (0.03)
0.02 (0.02)
0.00 (0.02)
0.00 (0.02)
FeO
9.11 (1.74)
8.67 (1.25)
9.63 (2.36)
9.07 (1.62)
8.75 (0.94)
0.89 (0.12)
0.92 (0.16)
0.91 (0.08)
0.89 (0.09)
1.01 (0.29)
MnO
0.13 (0.06)
0.13 (0.04)
0.12 (0.03)
0.12 (0.04)
0.12 (0.04)
0.06 (0.04)
0.05 (0.04)
0.06 (0.04)
0.06 (0.04)
0.06 (0.03)
MgO
49.67 (1.30)
49.05 (1.94)
49.13 (2.27)
50.04 (1.11)
49.70 (1.35)
18.06 (0.54)
17.71 (0.64)
17.42 (0.68)
17.49 (0.40)
17.59 (0.47)
NiO
0.37 (0.04)
0.33 (0.12)
0.37 (0.10)
0.36 (0.11)
0.37 (0.16)
0.01 (0.02)
0.02 (0.03)
0.01 (0.04)
0.00 (0.03)
0.01 (0.03)
CaO
0.07 (0.04)
0.11 (0.05)
0.10 (0.05)
0.11 (0.02)
0.13 (0.06)
25.62 (0.54)
25.52 (0.30)
25.33 (0.22)
25.50 (0.20)
25.31 (0.63)
Na2O
0.00 (0.02)
0.02 (0.02)
0.02 (0.05)
0.01 (0.03)
0.04 (0.06)
0.37 (0.16)
0.33 (0.31)
0.44 (0.10)
0.44 (0.05)
0.47 (0.17)
K2O
0.00 (0.00)
0.03 (0.01)
0.03 (0.02)
0.02 (0.01)
0.03 (0.02)
0.01 (0.00)
0.02 (0.02)
0.02 (0.01)
0.02 (0.01)
0.03 (0.03)
TOTAL
100.5
99.8
99.8
100.5
100.1
100.2
100.2
99.4
99.9
99.9
SiO2
n = number of analyses
2σ uncertainties are shown in parentheses
Data collected using the Cameca SX-100 Electron Probe Micro Analyzer at Rensselaer Polytechnic Institute
* Data collected using the JEOL JXA-8900 Electron Probe Micro Analyzer at the University of Maryland
All experiments run at 900°C and 1 GPa
22
Table 3.2.2. Trace element concentrations of starting materials and experimental
run products
fO2 Buffer [Li] (ppm) 1σ
[Ba] (ppm) 1σ
n
Olivine
San Carlos Olivine*
TJ-52
TJ-63
TJ-67
TJ-68
Mo-MoO2
MW
FMQ
Unbuffered
1.80
20.6
10.6
9.59
26.8
0.10
2.0
3.7
2.03
6.6
BDL
BDL
BDL
BDL
BDL
BDL
BDL
BDL
BDL
BDL
3
9
5
3
3
Clinopyroxene
DeKalb Diopside*
5.10
0.40
BDL
BDL 8
TJ-52
Mo-MoO2
11.1
2.8
9.40
1.30 3
TJ-63
MW
6.90
1.04
BDL
BDL 3
TJ-67
FMQ
5.79
1.36
BDL
BDL 3
TJ-68
Unbuffered
13.2
3.8
BDL
BDL 4
All experiments run at 900°C and 1 GPa
Trace element data collected using the Thermo X-Series II Quadrupole ICP-MS
with New Wave UP-213 laser ablation system at PSU
* Data collected using the Thermo-Finnigan Element II ICP-MS with New Wave
UP-213 laser ablation system at the University of Maryland
n = number of analyses
BDL = below detection limit
23
Table 3.4.1. Calculated Li partition coefficients (DLiol/cpx)
Sample
fO2 Buffer
D(ol/cpx)
1σ
TJ-52
Mo-MoO2
1.9
0.5
TJ-63
MW
1.5
0.6
TJ-67
FMQ
1.7
0.5
TJ-68
Unbuffered
2.0
0.7
AVERAGE
1.7
0.2
All experiments run at 900°C and 1 GPa
24
B
A
Clinopyroxene
Olivine
Figure 3.1.1. Backscattered electron SEM images of run products from experiment TJ-67.
These crystals have not been mounted in epoxy and polished; rather they have been
mounted on electrically conductive tape to image via SEM and show morphology of
unpolished crystals. Image A shows multiple olivine crystals surrounded by quench
material. Image B shows one large elongate clinopyroxene.
25
Clinopyroxene
Olivine
Figure 3.1.2. SEM backscattered electron image of run products from experiment TJ-52
buffered at Mo-MoO2. Sample was mounted in epoxy and polished with 600 grit siliconcarbide sandpaper and 1µm diamond paste. Clinopyroxene is typically more elongate
and brighter than olivine in backscattered electron images.
26
Olivine
Clinopyroxene
Figure 3.1.3. SEM backscattered electron image of run products from experiment TJ-63
buffered at Magnetite-Wüstite. Sample was mounted in epoxy and polished with 600 grit
silicon-carbide sandpaper and 1µm diamond paste. Clinopyroxene is typically more
elongate and brighter than olivine in backscattered electron images.
27
Clinopyroxene
Olivine
Figure 3.1.4. SEM backscattered electron image of run products from experiment TJ-67
buffered at FMQ. Sample was mounted in epoxy and polished with 600 grit siliconcarbide sandpaper and 1µm diamond paste. Clinopyroxene is typically more elongate
and brighter than olivine in backscattered electron images.
28
Olivine
Clinopyroxene
Figure 3.1.5. SEM backscattered electron image of run products from experiment TJ-68
which was unbuffered. Sample was mounted in epoxy and polished with 600 grit
silicon-carbide sandpaper and 1µm diamond paste. Clinopyroxene is typically more
elongate and brighter than olivine in backscattered electron images.
29
BCR-2g
1,000,000.0
Signal
Counts/Second
Background
Li
100,000.0
Mg
Mg
10,000.0
Ca
Ti
Ni
1,000.0
100.0
0.000
Ba
10.000 20.000 30.000 40.000 50.000 60.000
Time (sec)
Figure 3.2.1. Typical time resolved laser ablation spectrum for standard BCR-2g used in
sample-standard bracketing. Black vertical lines indicate the background differentiated
from the analysis signal. LA-ICP-MS data was reduced with LasyBoy V3.01, a data
reduction spreadsheet created by Joel Sparks at Boston University.
30
20000
18000
16000
TJ-52
14000
Mo
Intensity (counts)
12000
MoO2
10000
8000
6000
4000
2000
0
15
25
35
45
55
65
2-Θ (deg)
Figure 3.3.1. XRD analysis of the fO2 buffer from TJ-52 is shown in black. The red
dotted line represents an XRD spectrum for Mo and the blue dashed line represents an
XRD spectrum for Mo-MoO2. TJ-52 shows clear peaks that overlap well with both
phases, suggesting that the experimental buffer was successful.
31
14000
12000
TJ-63
10000
Magnetite
Intensity (counts)
8000
6000
4000
2000
0
15
25
35
45
55
65
2-Θ (deg)
Figure 3.3.2. XRD analysis of the fO2 buffer from TJ-63 is shown in black. The red
dashed line represents an XRD spectrum for magnetite. There is no confirmed wüstite in
this buffer, but the broad increase in intensity from ~30-45° represents data from an
amorphous phase with no defined crystal structure. This is being interpreted as the
remnants of the wüstite present at run conditions which decomposed upon quenching as
wüstite is not stable below 550°C at 1 atm.
32
10000
9000
8000
TJ-67
Fayalite
Magnetite
Quartz
7000
Intensity (counts)
6000
5000
4000
3000
2000
1000
0
15
25
35
45
55
65
2-Θ (deg)
Figure 3.3.3. XRD analysis of the fO2 buffer from TJ-67 is shown in black. The red
dashed line represents an XRD spectrum for fayalite. The blue dash-dot line represents an
XRD spectrum for magnetite. The green dotted line represents an XRD spectrum for
quartz. While the strongest peaks in TJ-67 overlap with magnetite, there are overlapping
peaks that only correspond to quartz and fayalite as well, suggesting that all three phases
are present at the conclusion of the experiment and the buffer was successful.
33
3.0
DLi(ol/cpx)
2.5
2.0
Buffered experiments
from this study
1.5
Unuffered experiment
from this study
1.0
Average 900°C
experiment from
Yakob et al., 2012
0.5
0.0
-18.0
-16.0
-14.0
-12.0
-10.0
log fO2
Figure 3.4.1. Experimental results showing the measured Li partition coefficient between
olivine and clinopyroxene as a function of fO2. Open symbol represents equilibrium
partition coefficient from 3-day 900°C experiment presented by Yakob et al., 2012. It is
assumed that the oxygen fugacity of the unbuffered experiments is near Ni-NiO, which has
a logfO2 of ~-12 at 900°C.
34
4.
DISCUSSION
4.1
Attainment of Equilibrium in Experiments
Previous piston-cylinder experiments conducted in the High-Pressure Laboratory
at Penn State (Yakob et al. 2012) demonstrated that Li partitioning between olivine and
clinopyroxene achieves equilibrium on the order of hours to days. A time series
conducted at 900°C and 1.5 GPa shows that the measured Li partition coefficient
between olivine and clinopyroxene remains constant (within error) for experiments
between 1 and 7 days in duration (Figure 4.1.1). Modeling of the diffusion of Li between
olivine and clinopyroxene supports this claim. Using the published diffusion coefficients
for Li in olivine (“fast path”; Dohmen et al., 2010) and clinopyroxene (Coogan et al.,
2005), Li partitioning between olivine and clinopyroxene approaches equilibrium within
crystals 50µm and smaller in 24 hours or less at 900°C. The experiments in this study
were held at 900˚C for ~3 days, allowing ample time for crystals to achieve equilibrium
with respect to Li partitioning.
4.2
Oxygen Fugacity
In gaseous systems, fugacity is defined as the difference between the chemical
potential of a real gas at a given pressure and temperature and the chemical potential of
an ideal gas at temperature and a standard pressure of 1 bar.
fi = exp{[µ(T,P) - µ(T,1 bar, pure)]/RT}
Where: fi = fugacity of gas i
μ = chemical potential
R = gas constant
T = temperature
Fugacity is the pressure an ideal gas would have at the same chemical potential as a real
gas. At high temperatures and low pressures we can think of fugacity as the partial
pressure of a gas in a system. Oxygen fugacity can be then defined as the chemical
35
potential or partial pressure of oxygen within a system. The concept of oxygen fugacity
(fO2) can be applied to solid, liquid and gaseous systems when we want to understand the
chemical potential of oxygen, which is important in understanding the oxidation
environment of a system, oxidation-reduction processes, speciation of ions with multiple
valence states, liquidus temperatures in melts and crystallization sequences.
Oxygen fugacity within solid systems is most easily thought of as the chemical
potential or activity of oxygen within a system. Typically in solid systems, oxygen
fugacity is reported as it relates to various solid state fO2 buffers such as fayalitemagnetite-quartz (FMQ) or magnetite-wüstite (MW) (Fig. 4.2.1). These buffer systems
work because O2 is a component on one side of a chemical reaction. The three oxygen
fugacity buffers used in this study were:
Fayalite + Oxygen ↔ Quartz + Magnetite
3 Fe2SiO4 + O2 ↔ 3 SiO2 + 2 Fe3O4
Wüstite + Oxygen ↔ Magnetite
3 FeO + ½ O2 ↔ Fe3O4
Molybdenum + Oxygen ↔ Molybdenum dioxide
Mo + O2 ↔ MoO2
Presuming all solid phases are present in the system (e.g. fayalite, magnetite and quartz),
the chemical activity of O2 can be calculated at a given temperature and pressure.
Generally speaking, any metal-oxide pair can be a solid-state fO2 buffer so long as it is
calibrated with respect to pressure and temperature.
Relatively speaking, “high fO2” environments are oxidizing and “low fO2”
environments are reducing. This is important in any environment where multivalent
cations, such as iron, are present because the oxidation environment plays a major control
on the speciation of the cation. As described in the Experimental Methods section, the
piston cylinder experiments presented here used San Carlos olivine and DeKalb diopside
as starting materials; these have ~10% and <1% FeO by weight, respectively. The
oxidation environment in which olivine and clinopyroxene crystals are grown from these
starting materials dictates the Fe3+/Fe within the system, and in turn the amount of Fe3+
incorporated into the minerals.
36
4.2.1
Oxygen Fugacity in the Mantle
The oxygen fugacity of the upper mantle has been calibrated as a function of Fe
activity in spinel/orthopyroxene/olivine xenolith assemblages. Results have shown that
the ambient fO2 of the upper mantle ranges from 2 log units above to 3 log units below
the FMQ solid state buffer (Mattioli and Wood, 1986; O’Neill and Wall, 1987). Lee et
al. (2005) developed a trace element method of determining the oxygen fugacity of
various mantle environments by calibrating the V/Sc system with fO2; V and Sc are both
non-fluid-mobile and behave very similarly in volcanic processes, which mean they are
generally unfractionated during the genesis of mid-ocean ridge basalts (MORBs) and arc
lavas. However, the speciation of V is redox sensitive while that of Sc is not. Therefore,
V/Sc ratios give an opportunity to calculate the oxygen fugacity of the mantle while
removing outside effects due to melt formation or metasomatism. Application of this
method to natural basalts suggests that the oxygen fugacity of arc lavas and MORBs
range from 1.25 log units below FMQ to 0.5 log units above FMQ (Lee et al. 2005).
More recently, Cottrell and Kelley (2011) used µ-XANES (micro x-ray absorption near
edge spectroscopy) to calculate the Fe+3/ΣFe of MORB glasses worldwide, showing that
the oxygen fugacity of primary MORB melts is very near the FMQ buffer, in close
agreement with the results from Lee et al. (2005). Beneath Archean cratons, the fO2
determined by Mössbauer spectroscopy of garnet and spinel from peridotite xenoliths
shows the ambient fO2 decreases with depth to nearly 5 log units below FMQ (Woodland
and Koch 2003). More recent experimental constraints on garnet oxythermobarometry in
xenoliths shows that even the most reducing deep lithosphere environments have fO2
~3.5 log units below the FMQ buffer, at least an order of magnitude more oxidizing than
previously thought (Stagno et al. 2013). As depth increases through the mantle towards
the core-mantle boundaries, oxygen fugacity continues to decrease to values below the
iron-wüstite buffer, where Fe and Ni metal can be stable (Frost and McCammon 2008).
Many different analytical and experimental methods have been applied to assess
the oxygen fugacity of the interior of the Earth. In summary, in the upper mantle near
mid-ocean ridges and within volcanic arcs, the fO2 is near the FMQ buffer (Mattioli and
Wood, 1986; O’Neill and Wall, 1987; Lee et al., 2005; Cottrell and Kelley, 2011).
37
Deeper within the mantle, the ambient fO2 decreases to values as low as 5 log units below
FMQ (Woodland and Koch, 2003; Stagno et al. 2013). The range of solid-state buffers
(plus the unbuffered experiment) used in this study was chosen to cover the full range of
fO2 values expected in the upper mantle.
4.3
Li Partitioning as a Function of fO2
The partitioning of Li between mantle phases could be influenced by the oxygen
fugacity of the surrounding environment. Being a monovalent cation, Li does not
partition into the crystal structure of olivine by direct substitution; rather, Li needs a
coupled substitution partner to preserve overall charge balance. One potential
mechanism allowing Li incorporation includes a coupled substitution with a trivalent
cation, where Li and the trivalent cation occupy the M1 and M2 sites typically reserved
for divalent cations within the structure. Fe and Mg typically make up the cations that
occupy these M1 and M2 sites within the olivine crystal structure. Fe has two valence
states, and oxygen fugacity effectively controls the Fe3+/Fe2+ within a system. A more
oxidizing fO2 environment would result in an increase in the relative abundance of Fe3+
within the system, effectively allowing for more coupled substitution with monovalent
cations, such as Li, into the olivine structure. Considering olivine within the mantle
contains ~10 wt. % FeO, and clinopyroxene contains ~3.5 wt. % FeO, more Fe3+ can be
incorporated in olivine relative to clinopyroxene. This is to say that an increasingly
oxidizing environment could allow for more Li to be incorporated into olivine rather than
clinopyroxene. As such, oxygen fugacity has been proposed as a possible controlling
factor in the partitioning of Li between mantle phases.
Li has been shown to be significantly affected by coupled substitution in olivine.
Experiments have shown that an increase in trivalent cations, specifically Al, Sc and Ga,
can lead to an increase in Li content of the olivine that scales with the increase in
trivalent cations (Grant and Wood, 2010). Caciagli et al. (2011) also show that an
increase in trivalent cations can show an increase in partitioning between mantle
minerals. An experiment with significantly increased Yb resulted in an ol/cpx partition
38
coefficient of 10, attributed to coupled substitution of Li+ with Yb3+. In a more complex
manner, Li has been linked to P concentrations in olivine (Mallmann et al. 2009). The
modal abundance of P is significantly higher than Li in this study, so a direct charge
balancing mechanism does not seem plausible. Mallmann et al. (2009) suggest that Li is
being incorporated into vacancies within the olivine structure created by the distortions
caused by P incorporation.
Caciagli (2010) showed that varying oxygen fugacity may have an effect on Li
partitioning between mantle phases. Two experiments were run with the intention of
varying the oxygen fugacity of the system, one at Re-ReO2 (highly oxidizing, 5 log units
> FMQ) and one at Mo-MoO2 (highly reducing, 5 log units < FMQ). The results from
these experiments are not entirely clear, as no olivine was grown in the oxidizing
experiment and the olivine/clinopyroxene partition coefficient determined in the reducing
experiment was 0.7±0.4, slightly lower than in unbuffered equilibrium experiments
(Caciagli, 2010) (Figure 4.4.1). This suggests that oxygen fugacity could possibly play a
role in equilibrium partitioning of Li between mantle minerals, but more data at a range
of oxygen fugacities is required.
The results from our experiments, however, show no apparent relationship
between oxygen fugacity and Li partitioning between olivine and clinopyroxene. We
have studied a wide range of oxygen fugacities spanning nearly all fugacities relevant to
the Earth’s mantle. The measured partition coefficient between olivine and
clinopyroxene in our run products show that under all fO2 conditions tested, the
equilibrium partition coefficient is 1.7±0.4 (Table 3.2.3). This, along with the results
from Yakob et al. (2012), suggests that closed-system redistribution of Li between mantle
minerals cannot explain the compositional and isotopic variability seen in mantle
xenoliths from around the world.
39
4.4
Li Variability in Mantle Minerals
The apparent disequilibrium distribution of Li and its isotopes is observed within
some mantle peridotites collected from localities worldwide. As shown in Figure 1.1.3,
the observed Li partition coefficients between olivine and clinopyroxene at some
localities deviate from experimentally derived equilibrium partition coefficients
determined from a variety of studies (Figure 4.4.1; Yakob et al., 2012; Ottolini et al.,
2009; Caciagli et al., 2011; Brenan et al., 1998; Blundy and Dalton, 2000). In addition,
the isotopic distribution of Li in many of the “non-equilibrium” samples shows
significant variation. One trend, most evident in xenoliths from Tok, Siberia (Rudnick
and Ionov, 2007) and Hannuoba, China (Tang et al., 2007b), shows increasing isotopic
fractionation with a decreasing partition coefficient. In other words, the clinopyroxenes in
these xenoliths are unusually enriched in isotopically light Li. Elevated concentrations of
isotopically light Li have also been observed in clinopyroxenes hosted in mantle
xenoliths from Kyushu, Japan, although the olivines in these samples were not analyzed
(Nishio et al., 2002).
4.4.1
Equilibrium Isotope Fractionation
The isotopic variability shown in some mantle xenoliths could be explained if the
equilibrium isotopic fractionation factor between olivine and clinopyroxene (ol/di) at
high temperature is greater than 1, such that 7Li is preferentially taken up by olivine
relative to clinopyroxene, resulting in a higher δ7Li for olivine. Seitz et al. (2004)
observed an apparent temperature dependence on isotopic fractionation as mantle
xenolith samples from Siberia, Germany and Austria, with the highest equilibration
temperatures (~1150°C) correlate with the lowest Δ7Liol-cpx (Δ7Liol-cpx = 1) (Δ7Liol-cpx =
δ7Liol-δ7Licpx), while samples with lower equilibration temperatures (~850-900°C) show
increased isotopic fractionation (Δ7Liol-cpx = 3.5-4).
The idea that equilibrium isotopic fractionation takes place at high temperatures
goes against thermodynamic convention, which suggests isotopic effects are minimized
with increasing temperature. If the isotopic fractionation observed between mantle
40
minerals is an equilibrium effect related to temperature, at high temperatures (>200 °C)
there should be a linear relationship between 1000*ln(α) vs. 106/T2 [where α =
7
Li/6Liol/7Li/6Licpx] (Northrop and Clayton 1966). A compilation of Li concentrations and
isotopic ratios measured in bulk mineral separates or individual crystals from mantle
xenoliths worldwide is presented in Table 4.4.1. The dataset reveals no apparent
relationship between temperature and isotopic fractionation when 1000*ln(α) is plotted
against 106/T2 (Figure 4.4.2). This suggests that the Li isotopic variability observed in
mantle xenoliths is not representative of equilibrium fractionation.
Furthermore, Tomascak et al. (1999) demonstrated a lack of measurable Li
isotopic fractionation as olivine, clinopyroxene, and plagioclase sequentially crystallized
out of a basaltic melt at temperatures between 1050 and >1200 ˚C at Kilauea Iki lava
lake. Throughout the process of fractional crystallization, the isotopic composition of the
residual melt remains unchanged. This suggests that there is no inherent isotopic
fractionation of Li in silicate systems including olivine and diopside and therefore the
isotopic fractionation factor is indeed close to 1 at temperatures above 1000°C.
4.4.2
Open vs. closed system redistribution
Redistribution of Li between coexisting phases could result in the pattern
displayed by mantle xenoliths that deviate from equilibrium behavior (Dol/cpx <1,
αol/cpx>1). Diffusive re-distribution of Li between mantle minerals has interesting
implications for isotopic fractionation. As has been shown by Richter et al. (2003), 6Li
diffuses faster than 7Li, such that diffusion could cause dynamic isotopic fractionation on
geologically short timescales by as much as 25‰. The fractionation could then be locked
into the minerals as they continue to cool beyond the Li closure temperature. A driving
force is required, but if Li is mobilized and allowed to move from olivine into
clinopyroxene, the partition coefficient would decrease and the isotopic fractionation
factor would increase. One potential driver for closed system Li redistribution is a
temperature dependence of the Li partition coefficient between olivine and
clinopyroxene. A xenolith or peridotite massif cools as it is exhumed. This change in
41
temperature might influence the partition coefficient such that Li becomes more
compatible in clinopyroxene relative to olivine as temperature decreases, resulting in Li
moving from the olivine to the clinopyroxene (thus decreasing the Dol/cpx) via diffusion.
By analogy, Coogan et al. (2005) showed that the Li partition coefficient between
clinopyroxene and plagioclase is dependent on temperature, and used this temperature
dependence to develop a “geo-speedometer” to determine cooling rates for mid-ocean
ridge basalts.
Ionov and Seitz (2008) invoked this hypothesis to propose that eruption style, and
therefore cooling rate, could be a mechanism for controlling Li distribution between
mantle minerals. They observed greater Li isotopic fractionation in presumably rapidly
cooled xenoliths sampled from explosive eruptions (tuff and scoria deposits) relative to
xenoliths sampled from lava flows, where cooling rates would be substantially slower.
The enhanced cooling rates in the xenoliths from explosive eruptions would allow the Li
isotope ratios to be locked into the cooling crystals near the peak of fractionation, while
the slower cooling rates would allow the minerals to approach equilibrium with one
another prior to dropping below the closure temperature. Kaliwoda et al. (2008) invoked
the same hypothesis, however in the opposite sense, to explain Li zoning in individual
crystals in mantle xenoliths from Saudi Arabia. They observed a rim-ward increase in Li
content of the olivines, while the rims of clinopyroxenes decreased their Li
concentrations relative to the cores. They concluded that this was the result of
temperature dependent partition coefficients changing such that Li moves from
clinopyroxene to the olivine upon cooling.
Yakob et al. (2012) experimentally tested whether the Li partition coefficient
between olivine and clinopyroxene is temperature dependent. Piston cylinder
experiments at 1.5 GPa and temperatures ranging from 700-1100°C all yielded DLiol/di =
2.0  0.2, suggesting that closed-system re-distribution of Li in response to changing
temperature is not responsible for the observed isotopic fractionation between olivine and
clinopyroxene in natural mantle xenoliths. Apparent DLiol/di measured from mineral
separates of mantle xenoliths worldwide also shows no correlation with temperature
(Figure 4.4.3). The experiments of Yakob et al. (2012), combined with the results from
42
this study, have failed to identify a suitable driving force for closed system redistribution
of Li capable of generating the non-equilibrium Li signature seen in mantle xenoliths.
This implies that the mechanism controlling the non-equilibrium Li signature seen in
mantle xenoliths is an open system mechanism.
Xenoliths can gain isotopically distinct Li by open-system interaction with a
metasomatic fluid or melt prior to or during their rise to the surface. This fluid or melt
may interact with each mantle mineral differently, depending on the relative diffusion
rates of lithium in each phase. Rudnick and Ionov (2007) propose a few different
mechanisms for interaction with an infiltrating fluid or melt. For example, if Li diffuses
from the fluid or melt into the clinopyroxene faster than it diffuses into the olivine, on a
short timescale such that equilibrium is not attained, the concentration of Li in
clinopyroxene would increase relative to that in olivine. Likewise, the isotopic
composition of the clinopyroxene would be relatively light due to faster diffusion of the
lighter 6Li. Alternatively, if Li diffuses into olivine faster than clinopyroxene, the olivine
could come to equilibrium with the fluid or melt, while the isotopically lighter
clinopyroxene still captures kinetically induced isotopic fractionation. The final, and
preferred, proposed mechanism of Rudnick and Ionov (2007), is that Li diffuses into
clinopyroxene faster than olivine, with the added condition that the fluid:rock ratio is
quite small. This would result in early uptake of isotopically light Li by clinopyroxene,
which would draw down the 6Li in the fluid such that the remaining Li would become
isotopically heavy. As time progresses, isotopically heavy Li would migrate into the
olivines, allowing for greater diversity in the Li isotope composition of minerals coexisting within the same xenolith sample. This hypothesis has yet to be experimentally
tested, although independent Li diffusion studies by Coogan et al. (2005) and Dohmen et
al. (2010) suggest that Li diffusion could be up to three orders of magnitude faster in
clinopyroxene relative to natural mantle olivine.
43
Table 4.4.1 Compilation of Li data from olivine and clinopyroxene mineral separates from xenoliths worldwide. Data
from Aulbach and Rudnick (2009)1, Aulbach et al. (2008)2, Ionov and Seitz (2008)3, Jeffcoate et al. (2007)4, Kaliwoda et
al. (2008)5, Magna et al. (2006, 2008)6,7, Ottolini et al. (2004)8, Rudnick and Ionov (2007)9, Seitz and Woodland (2000)10,
Seitz et al. (2004)11, Tang et al. (2007b)12, and Woodland et al. (2004)13.
--- = values not reported
Sample
Location
Li (ppm)
δ7Li
Olivine
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
89-6611
Lashaine, Tanzania
0.7
---
---
---
---
---
---
1090
89-6641
Lashaine, Tanzania
---
---
1.4
-2.5
---
---
---
1230
89-669
1
Lashaine, Tanzania
2.7
---
3.6
-3.2
0.75
---
---
1040
89-6711
Lashaine, Tanzania
1.7
13.9
---
---
---
---
---
---
89-6721
Lashaine, Tanzania
1.1
4.2
1.6
0.4
0.69
3.8
1.004
1110
89-6741
Lashaine, Tanzania
1.5
5
2.2
---
---
---
---
1250
89-6751
Lashaine, Tanzania
0.6
---
---
---
---
---
---
1240
89-6801
Lashaine, Tanzania
1.5
2.3
0.9
-1.5
1.67
3.8
1.004
1150
89-719
1
Lashaine, Tanzania
1.1
2.8
1.2
-3.9
0.92
6.7
1.007
1150
89-7721
Olmani, Tanzania
4.1
4.1
1.5
-2
2.73
6.1
1.006
950
89-7731
Olmani, Tanzania
0.7
3
---
---
---
---
---
1080
89-7741
Olmani, Tanzania
3
3.3
---
---
---
---
---
---
89-7761
Olmani, Tanzania
3.7
6.2
1.6
-0.4
2.31
6.6
1.007
1120
89-777
1
Olmani, Tanzania
3.5
5.7
1.9
-1.4
1.84
7.1
1.007
1120
89-7781
Olmani, Tanzania
2.6
9.3
---
---
---
---
---
---
44
Sample
Location
Li (ppm)
δ7Li
Olivine
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
LB-29
2
Labait, Tanzania
1.9
3.4
---
---
---
---
---
---
LB-312
Labait, Tanzania
1.8
3.3
0.9
1.5
2.00
1.8
1.002
---
KAT-172
Labait, Tanzania
3.4
4.7
2.2
-0.3
1.55
5
1.005
---
LB-42
Labait, Tanzania
2.3
4.7
1.2
-2.6
1.92
7.3
1.007
---
LB-452
Labait, Tanzania
2.4
4.7
2
0.1
1.20
4.6
1.005
---
LB-6
Labait, Tanzania
1.9
3.7
---
---
---
---
---
---
LB-172
Labait, Tanzania
2.7
4
---
---
---
---
---
---
LB-212
Labait, Tanzania
2.2
2.5
3.8
-4.9
0.58
7.4
1.007
---
LB-462
Labait, Tanzania
4.4
5.2
3
-6.7
1.47
11.9
1.012
---
LB-512
Labait, Tanzania
4.8
6.6
---
---
---
---
---
---
LB-592
2
Labait, Tanzania
3.2
6.6
---
---
---
---
---
---
LB-2
2
Labait, Tanzania
5
10.1
---
---
---
---
---
---
Av-13
Avacha, Siberia
1.35
---
---
---
---
---
---
919
Av-23
Avacha, Siberia
1.59
3.25
---
---
---
---
---
906
Av-33
Avacha, Siberia
1.42
-2.1
---
---
---
---
---
946
Av-43
Avacha, Siberia
0.91
2.95
---
---
---
---
---
920
Av-53
Avacha, Siberia
1.17
---
---
---
---
---
---
---
Av-6
3
Avacha, Siberia
1.75
4.71
---
---
---
---
---
993
Av-73
Avacha, Siberia
1.26
---
---
---
---
---
---
---
45
Sample
Location
Li (ppm)
δ7Li
Olivine
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
Av-8
3
Avacha, Siberia
1.54
2.24
---
---
---
---
---
978
Av-93
Avacha, Siberia
1.32
---
---
---
---
---
---
---
Av-103
Avacha, Siberia
1.21
---
---
---
---
---
---
---
Av-113
Avacha, Siberia
1.81
0.3
---
---
---
---
---
930
Av-123
Avacha, Siberia
1.24
---
---
---
---
---
---
---
Av-13
3
Avacha, Siberia
1.4
---
---
---
---
---
---
---
Av-143
Avacha, Siberia
1.28
---
---
---
---
---
---
---
Av-153
Avacha, Siberia
1.37
1.53
---
---
---
---
---
960
Av-163
Avacha, Siberia
1.24
0.86
---
---
---
---
---
989
Av-173
Avacha, Siberia
1.33
---
---
---
---
---
---
---
621-163
Vitim, Siberia
1.15
6.3
5.08
-17.9
0.23
24.2
1.025
1176
313-105
4
Vitim, Siberia
2.32
4.7
1.12
1.1
2.07
3.6
1.004
1034
313-1054
Vitim, Siberia
2.4
4.9
1.18
1
2.03
3.9
1.004
1034
313-1024
Vitim, Siberia
1.97
3.7
1.15
3.5
1.71
0.2
1.000
1053
314-564
Vitim, Siberia
1.87
3
1.3
6.6
1.44
-3.6
0.996
889
Mo-Z-14
Tariat, Mongolia
1.93
3
1.11
2.6
1.74
0.4
1.000
890
4230-164
Tariat, Mongolia
1.75
3.1
1.15
4.7
1.52
-1.6
0.998
867
8505-2
Dariganga, Mongolia
1.4
3
1.25
1.6
1.12
1.4
1.001
914
SC4
San Carlos, AZ
2.38
5.2
1.48
-7.8
1.61
13
1.013
1002
4
46
Sample
Location
Li (ppm)
δ7Li
Olivine
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
4
San Carlos, AZ
2.31
4.8
1.5
-8.7
1.54
13.5
1.014
1002
SC4
San Carlos, AZ
2.29
5
1.48
-8.1
1.55
13.1
1.013
1002
Th294
Icelandic basalt
1.25
3.3
---
---
---
---
---
---
AH4
Hawaiian basalt
1.42
3.5
---
---
---
---
---
---
97KC06 (1960) 4
Hawaiian basalt
1.18
6.6
0.85
1.6
1.39
5
1.005
---
4
Hawaiian basalt
1.28
6.5
0.78
1.8
1.64
4.7
1.005
---
SA84-32 cores5
Harrat Uwayrid, Saudi Arabia
1.382
---
0.917
---
1.51
---
---
971
SA84-32 rims5
Harrat Uwayrid, Saudi Arabia
1.47
---
0.666
---
2.21
---
---
988
SA84-123 cores5
Harrat Uwayrid, Saudi Arabia
1.296
---
1.179
---
1.10
---
---
805
SA84-123 rims5
Harrat Uwayrid, Saudi Arabia
1.179
---
0.62
---
1.90
---
---
844
SA84-124b cores5
Harrat Uwayrid, Saudi Arabia
1.481
---
0.581
---
2.55
---
---
884
5
Harrat Uwayrid, Saudi Arabia
1.6
---
0.522
---
3.07
---
---
884
SA84-128/1 cores5
Harrat Uwayrid, Saudi Arabia
1.563
---
0.98
---
1.59
---
---
998
SA84-128/1 rims5
Harrat Uwayrid, Saudi Arabia
1.563
---
0.853
---
1.83
---
---
1019
SA84-128/2 cores5
Harrat Uwayrid, Saudi Arabia
1.625
---
1.003
---
1.62
---
---
1051
SA84-128/2 rims5
Harrat Uwayrid, Saudi Arabia
2.081
---
0.833
---
2.50
---
---
1061
SA84-38 cores5
Harrat Uwayrid, Saudi Arabia
1.577
---
1.393
---
1.13
---
---
874
5
Harrat Uwayrid, Saudi Arabia
1.569
---
0.544
---
2.88
---
---
852
SA84-97 cores5
Harrat Uwayrid, Saudi Arabia
2.163
---
1.19
---
1.82
---
---
799
SC
97KC05 (1955)
SA84-124b rims
SA84-38 rims
47
Sample
Location
Li (ppm)
δ7Li
Olivine
SA84-97 rims
5
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
Harrat Uwayrid, Saudi Arabia
2.281
---
0.718
---
3.18
---
---
809
SA84-128/3 cores5
Harrat Uwayrid, Saudi Arabia
2.968
---
1.936
---
1.53
---
---
987
SA84-128/3 rims5
Harrat Uwayrid, Saudi Arabia
3.373
---
1.843
---
1.83
---
---
1028
SA84-91 cores5
Harrat Uwayrid, Saudi Arabia
3.883
---
1.766
---
2.20
---
---
1039
SA84-91 rims5
Harrat Uwayrid, Saudi Arabia
3.823
---
1.619
---
2.36
---
---
1060
5
Harrat Uwayrid, Saudi Arabia
1.923
---
1.134
---
1.70
---
---
1122
SA84-166 rims5
Harrat Uwayrid, Saudi Arabia
2.029
---
1.089
---
1.86
---
---
1118
SA84-50 cores5
Harrat Uwayrid, Saudi Arabia
1.934
---
1.016
---
1.90
---
---
---
SA84-50 rims5
Harrat Uwayrid, Saudi Arabia
1.577
---
0.873
---
1.81
---
---
---
SA84-63 cores5
Harrat Uwayrid, Saudi Arabia
3.865
---
2.085
---
1.85
---
---
900
SA84-63 rims5
Harrat Uwayrid, Saudi Arabia
4.067
---
2.451
---
1.66
---
---
981
San Carlos, AZ
1.6
3.64
0.52
3.32
3.08
0.32
1.000
---
314-586
Vitim, Siberia
1.9
3.76
1.4
5.21
1.36
-1.45
0.999
---
8520-096
Atsagin-Dush, Mongolia
1.5
3.55
0.82
5.96
1.83
-2.41
0.998
---
MPH 79/16
Tariat, Mongolia
2
3.75
1.4
2.57
1.43
1.18
1.001
---
96-26
Kilbourne Hole, NM
1.7
3.56
---
---
---
---
---
---
Atsagin-Dush, Mongolia
1.74
5.21
1.11
4.34
1.57
0.87
1.001
---
BM90-6
Balmuccia, Italy
1.16
---
1.18
---
0.98
---
---
---
BM90-12L8
Balmuccia, Italy
1.1
---
2.36
---
0.47
---
---
---
SA84-166 cores
SC
6
8250-157
8
48
Sample
Location
Li (ppm)
δ7Li
Olivine
8
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
LH12
Lherz, France
1.75
---
1.77
---
0.99
---
---
---
Z20688
Zabargad Island
2.09
---
3.9
---
0.54
---
---
---
Z20708
Zabargad Island
2.73
---
2.1
---
1.30
---
---
---
Z20998
Zabargad Island
2
---
1.75
---
1.14
---
---
---
BR118
Massif Central, France
1.92
---
1.13
---
1.70
---
---
---
8
Massif Central, France
1.67
---
0.99
---
1.69
---
---
---
RP688
Massif Central, France
1.87
---
1.13
---
1.65
---
---
---
RP708
Massif Central, France
1.88
---
0.95
---
1.98
---
---
---
Z78
Massif Central, France
2.26
---
0.62
---
3.65
---
---
---
Bt398
Massif Central, France
2.91
---
1.45
---
2.01
---
---
---
Szt10638
Carpatho-Pannonian Basin
1.2
---
0.54
---
2.22
---
---
---
8
Carpatho-Pannonian Basin
1.55
---
0.84
---
1.85
---
---
---
Ib/kl8
Dreiser Weiher, Germany
1.34
---
0.75
---
1.79
---
---
---
D588
Dreiser Weiher, Germany
1.42
---
0.82
---
1.73
---
---
---
D508
Dreiser Weiher, Germany
1.36
---
0.67
---
2.03
---
---
---
6--19
Tok, Siberia
1.3
8.1
3.3
-6.9
0.39
15
1.015
1010
6--29
Tok, Siberia
2
---
2.3
---
0.87
---
---
1001
6--3
9
Tok, Siberia
1.5
11.8
6.7
-9.3
0.22
21.1
1.021
976
7--59
Tok, Siberia
1.6
11.9
4.8
-9.8
0.33
21.7
1.022
985
BR12
Bol082
49
Sample
Location
Li (ppm)
δ7Li
Olivine
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
8--5
9
Tok, Siberia
2.1
0.8
4.1
-12.9
0.51
13.7
1.014
1004
8--69
Tok, Siberia
1.5
8.1
4.3
-10.7
0.35
18.8
1.019
985
8--399
Tok, Siberia
1.8
8.3
3.2
-14.6
0.56
22.9
1.023
964
1--29
Tok, Siberia
4.2
0.6
4.4
-3
0.95
3.6
1.004
910
1--39
Tok, Siberia
4.6
---
9.3
---
0.49
---
---
907
1--13
Tok, Siberia
3.1
---
8.2
---
0.38
---
---
933
2--69
Tok, Siberia
3.6
---
5.7
---
0.63
---
---
980
2--99
Tok, Siberia
3.1
---
2.4
---
1.29
---
---
874
3--49
Tok, Siberia
4.3
---
5.7
---
0.75
---
---
910
3--199
Tok, Siberia
2.6
---
0.9
---
2.89
---
---
931
5--39
Tok, Siberia
2.9
---
2.2
---
1.32
---
---
907
6--0
9
Tok, Siberia
2.3
---
---
---
---
---
---
890
7--19
Tok, Siberia
3.7
---
4.2
---
0.88
---
---
985
8--19
Tok, Siberia
2.3
---
10
---
0.23
---
---
1005
8--29
Tok, Siberia
3
---
9
---
0.33
---
---
976
8--39
Tok, Siberia
3.8
---
1.7
---
2.24
---
---
956
8--79
Tok, Siberia
4.3
---
6
---
0.72
---
---
968
8--8
9
Tok, Siberia
3.2
---
8.6
---
0.37
---
---
955
8--119
Tok, Siberia
4.4
---
6.7
---
0.66
---
---
957
9
50
Sample
Location
Li (ppm)
δ7Li
Olivine
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
8--31
9
Tok, Siberia
3
2.3
8.3
-7.3
0.36
9.6
1.010
887
8--509
Tok, Siberia
2.3
-1.7
2.3
-9.1
1.00
7.4
1.007
992
10--29
Tok, Siberia
4.3
---
1.3
---
3.31
---
---
914
10--49
Tok, Siberia
2.8
---
1.6
---
1.75
---
---
926
10--89
Tok, Siberia
3.2
---
2.7
---
1.19
---
---
950
10--16
9
Tok, Siberia
1.1
3.3
1.6
-1.7
0.69
5
1.005
957
10--179
Tok, Siberia
1.5
2.1
1.4
-1.9
1.07
4
1.004
1011
10--199
Tok, Siberia
1.3
2.8
1.2
-2.1
1.08
4.9
1.005
951
2--19
Tok, Siberia
3.6
---
11.5
---
0.31
---
---
---
2--29
Tok, Siberia
2.4
---
3
---
0.80
---
---
909
2--39
Tok, Siberia
4
---
3.7
---
1.08
---
---
908
2--4
9
Tok, Siberia
3.7
---
8.4
---
0.44
---
---
---
2--109
Tok, Siberia
3.6
---
8.7
---
0.41
---
---
1024
3--29
Tok, Siberia
3.2
2
1.6
-0.8
2.00
2.8
1.003
949
3--229
Tok, Siberia
3.4
---
2
---
1.70
---
---
---
8--109
Tok, Siberia
3
---
7.3
---
0.41
---
---
984
10--19
Tok, Siberia
2.7
---
1.7
---
1.59
---
---
964
10--3
9
Tok, Siberia
3.6
---
1.6
---
2.25
---
---
982
10--119
Tok, Siberia
5.5
---
7.4
---
0.74
---
---
920
51
Sample
Location
Li (ppm)
δ7Li
Olivine
9
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
13--8
Sikhote-Alin
4.8
4.7
6.6
-6.3
0.73
11
1.011
905
8802-19
Sikhote-Alin
1.5
4
1.2
3
1.25
1
1.001
987
Ke 1906/110
Chyulu Hills, Kenya
---
---
18.67
---
---
---
---
778
Ke 771/310
Chyulu Hills, Kenya
---
---
1.18
---
---
---
---
878
MC 21/210
Cantal, France
1.18
---
3.62
---
0.33
---
---
852
Chyulu Hills, Kenya
2.13
---
0.96
---
2.22
---
---
912
MC 34/110
Deves, France
2.44
---
1.36
---
1.79
---
---
940
Ke 50310
Chyulu Hills, Kenya
---
---
2.41
---
---
---
---
974
Ke 1921/510
Chyulu Hills, Kenya
---
---
1.56
---
---
---
---
977
Ke 1916/310
Chyulu Hills, Kenya
---
---
1.59
---
---
---
---
1014
Vi 313-1010
Vitim, Siberia
1.28
---
0.75
---
1.71
---
---
1050
San Carlos, AZ
1.58
---
1
---
1.58
---
---
1052
MC 45/310
Vivarais, France
1.44
---
0.92
---
1.57
---
---
1066
Pa110
Pali Aike, Chile
1.27
---
0.74
---
1.72
---
---
1120
Mo 2210
Tariat, Mongolia
1.42
---
0.79
---
1.80
---
---
1137
D 4510
Dreiser Weiher, Germany
1.23
---
0.71
---
1.73
---
---
1147
Causses, France
1.26
---
0.74
---
1.70
---
---
1180
Chyulu Hills, Kenya
1.1
---
0.73
---
1.51
---
---
1346
Marsabit, Kenya
1.34
---
0.71
---
1.89
---
---
1367
Ke 1928/3
SC 1
10
10
MC 49/110
Ke 514/1
10
Ke 604/610
52
Sample
Location
Li (ppm)
δ7Li
Olivine
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
10
Chyulu Hills, Kenya
1.61
---
0.86
---
1.87
---
---
1363
Ke 1927/210
Chyulu Hills, Kenya
1.22
---
0.61
---
2.00
---
---
1364
SC11
San Carlos, AZ
1.7
3.1
---
---
---
---
---
---
SC11
San Carlos, AZ
1.8
2.2
---
---
---
---
---
---
San Carlos, AZ
1.7
3.4
1.9
-0.8
0.89
4.2
1.004
1052
Vitim, Siberia
1.6
4.2
1
2.5
1.60
1.7
1.002
1050
Ka 16711
Kapfenstein, Austria
1.8
3.2
1.1
1.5
1.64
1.7
1.002
1040
Ia/17111
Dreiser Weiher, Germany
1.3
3.1
0.8
0.4
1.63
2.7
1.003
842
Ia/21111
Dreiser Weiher, Germany
1.8
1.4
1
-2.4
1.80
3.8
1.004
969
Ib/311
Dreiser Weiher, Germany
1.4
3
0.8
1.3
1.75
1.7
1.002
1185
Ib/811
Ke 1927/8
SC-111
11
Vi 313-10
Dreiser Weiher, Germany
1.6
4.5
1
2.8
1.60
1.7
1.002
1190
11
Dreiser Weiher, Germany
1.2
3.9
0.8
3
1.50
0.9
1.001
1152
Ib/K111
Dreiser Weiher, Germany
1
2.4
0.9
2.3
1.11
0.1
1.000
1121
Ib/5811
Dreiser Weiher, Germany
1.3
2.1
---
---
---
---
---
---
D112
Damaping, China
1.8
3.34
3.01
-8.25
0.60
11.59
1.012
---
D212
Damaping, China
1.66
6.39
3.49
-5.59
0.48
11.98
1.012
---
D312
Damaping, China
1.19
5.37
3.59
-5.32
0.33
10.69
1.011
---
D4
12
Damaping, China
1.51
5.35
2.41
-3.25
0.63
8.6
1.009
---
D512
Damaping, China
1.37
5.2
2.72
-7.54
0.50
12.74
1.013
---
Ib/24
53
Sample
Location
Li (ppm)
δ7Li
Olivine
Li (ppm)
δ7Li
DLiol/cpx
Δ7Liol-cpx
αol/cpx
T (°C)
Clinopyroxene
70972
13
Victoria, Australia
4.75
---
1.55
---
3.06
---
---
---
7098713
Victoria, Australia
2.87
---
1.4
---
2.05
---
---
---
7100013
Victoria, Australia
3.13
---
1.13
---
2.77
---
---
---
7100413
Victoria, Australia
2.67
---
2.42
---
1.10
---
---
---
7100613
Victoria, Australia
4.05
---
0.63
---
6.43
---
---
---
13
SH 45
Victoria, Australia
2.63
---
1.28
---
2.05
---
---
---
7698913
Victoria, Australia
1.75
---
0.89
---
1.97
---
---
---
7096513
Victoria, Australia
1.98
---
0.59
---
3.36
---
---
---
7100313
Victoria, Australia
2.29
---
0.98
---
2.34
---
---
---
7100813
Victoria, Australia
2.92
---
0.71
---
4.11
---
---
---
7699313
Victoria, Australia
3.42
---
1.42
---
2.41
---
---
---
76995
13
Victoria, Australia
1.34
---
1.88
---
0.71
---
---
---
7699713
Victoria, Australia
1.73
---
0.52
---
3.33
---
---
---
SH 3513
Victoria, Australia
2.11
---
1.27
---
1.66
---
---
---
7100113
Victoria, Australia
5.32
---
1.13
---
4.71
---
---
---
7102313
Victoria, Australia
2.22
---
1.35
---
1.64
---
---
---
7699113
Victoria, Australia
5.14
---
1.18
---
4.36
---
---
---
13
Victoria, Australia
5.71
---
0.6
---
9.52
---
---
---
76994
54
3.5
3
Dol.cpx
2.5
2
1.5
1
0.5
0
0
2
4
6
8
Run Duration (days)
Figure 4.1.1. Time series of piston cylinder equilibrium partitioning experiments from
Yakob et al. (2012) showing the DLiol/cpx evolving over time. At 900°C the system
approaches equilibrium with respect to Li partitioning between olivine and clinopyroxene
within 24 hours. Figure 4 from Yakob et al., 2012.
55
-5.00
-7.00
log fO2
-9.00
-11.00
-13.00
-15.00
FMQ
-17.00
MW
-19.00
Mo-MoO2
-21.00
700
900
1100
1300
1500
1700
Temperature (°C)
Figure 4.2.1. Solid state oxygen fugacity buffers rhenium-rhenium dioxide
(Re-ReO2), fayalite-magnetite-quartz (FMQ), magnetite-wüstite (MW), and
molybdenum-molybdenum dioxide (Mo-MoO2) shown as a function of
temperature. Vertical black line indicates experimental run conditions. Buffer
curves calculated from: Mo-MoO2 = O’Neill, 1986; Magnetite-Wüstite &
Fayalite-Magnetite-Quartz = Myers and Eugster, 1983.
56
5
Yakob et al. 2012
Ottolini et al. 2009
4
Brenan et al. 1998
Dol/cpx
Blundy and Dalton 2000
3
Caciagli et al. 2011
2
1
0
600
700
800
900
1000
1100
1200
1300
1400
Temperature (C)
Figure 4.4.1. Experimentally derived DLiol/cpx showing that the olivine/clinopyroxene
partition coefficient is ~2±0.5 regardless of temperature. Figure 6 from Yakob et al.,
2012. Data from Ottolini et al. (2009), Caciagli et al. (2011), Brenan et al. (1998), Yakob
et al. (2012) and Blundy and Dalton (2000)
57
Isotopic Fractionation
30
25
ln(α)*103
20
15
10
5
0
-5
0.6
0.8
1
1.2
1.4
1.6
1.8
(1/T^2)*106
Figure 4.4.2. Northrop and Clayton (1966) state that a linear relationship should exist
between ln(α)*103 vs. T-2*106 if isotopic fractionation exists as an equilibrium
phenomenon. Plotting ln(α)*103 vs. T-2*106 for all available xenolith data shows that in
most cases there is no observable relationship, and therefore the isotopic fractionation
seen in xenoliths is not the product of equilibrium isotope fractionation. A particular
suite from Rudnick and Ionov (2007), however, do seem to fit to a linear trend
suggesting a potential relationship between isotopic fractionation and temperature. Data
from Table 4.4.1
58
4
3.5
3
DLiol/cpx
2.5
2
1.5
1
0.5
0
750
950
1150
1350
Temperature (°C)
Figure 4.4.3. Apparent Li partition coefficients measured from mantle xenolith mineral
separates worldwide show no relationship with temperature. Data from Table 4.4.1.
59
5.
CONCLUSIONS
If equilibrium with respect to lithium distribution in mantle minerals is defined by
DLiol/di = 2.0  0.5 and 7Liol-di = 0  5‰, then many, but not all, mantle xenoliths display
non-equilibrium behavior with respect to Li partitioning and/or Li isotopic fractionation
between coexisting mineral phases. The lack of any clear relationship between isotopic
fractionation and temperature suggests that the isotopic variability seen in mantle
xenoliths is not an equilibrium phenomenon, as isotopic effects should be minimized at
high temperature, but rather a short lived kinetic effect resulting from the differing
diffusivities of 6Li and 7Li. The driving force behind Li redistribution could be a result of
closed or open system processes.
One theory that has been invoked in the literature without experimental backing is
that equilibrium elemental partitioning of Li between olivine and clinopyroxene changes
as a function of temperature such that Li becomes more compatible in clinopyroxene as
cooling progresses. The resulting closed-system diffusion of lithium from olivine to
clinopyroxene would be accompanied by isotopic fractionation, as the lighter 6Li diffuses
more rapidly than the heavier 7Li. While diffusion is in progress, then, clinopyroxene
would become isotopically lighter and olivine would become isotopically heavier. This
dynamic fractionation could be preserved in the minerals if the xenolith temperature
drops below the closure temperature for Li diffusion before equilibration is complete.
However, Yakob et al. (2012) demonstrated that the equilibrium partition coefficient for
lithium between olivine and diopside is independent of temperature between 700˚C and
1100˚C, suggesting a temperature dependent partition coefficient is not the driver for Li
redistribution.
Another quasi-closed-system process that could drive Li diffusion between
olivine and clinopyroxene is a change in the ambient oxygen fugacity. Substitution of Li+
for Mg2+ or Fe2+ in the olivine structure requires coupled substitution with a trivalent ion
in order to maintain charge balance. One candidate for a substitution partner is Fe3+.
Because Fe3+/Fe is a function of oxygen fugacity, a shift toward higher fO2 in the
mantle would result in greater Fe3+ incorporation in olivine, and potentially greater
opportunity for Li substitution in turn. By this logic, systems at higher fO2 should shift
60
toward higher DLiol/di and systems at lower fO2 should shift toward lower DLiol/di. I have
shown through piston cylinder experiments using solid-state fO2 buffers that the partition
coefficient for Li between olivine and clinopyroxene remains between 1.5-2.0 at 1.0 GPa
and 900˚C, over a range of oxygen fugacity values representative of the upper mantle
(~FMQ to ~5 log units below FMQ).
This result suggests that closed-system redistribution of Li between phases is not
a feasible method of generating the Li signatures observed in mantle xenoliths. Another
possibility is that the Li signature is instead affected by open-system interaction with
metasomatic fluids or melts prior to or during exhumation. This implies that the Li
signature of many mantle xenoliths may not be representative of the ambient mantle, but
is instead a mixture of the mantle and the metasomatic agent. The resetting of Li isotopic
distribution in mantle xenoliths on short timescales prior to or during exhumation limits
the applicability of stable Li isotopes as a passive tracer of recycled material in the
mantle, but opens the possibility of using Li as a geo-speedometer to measure the
timescales of short-lived processes such as magma migration and volcanic eruptions.
61
6.
FAILED EXPERIMENTS
6.1
Methodology
We wanted to develop an experimental setup to effectively control the oxygen
fugacity within the sample chamber of our piston cylinder experiments while at high
pressure and temperature for the duration of the experiment. Considering a piston
cylinder experiment is effectively a closed system, the only experimental design that
could control the oxygen fugacity was one that included a solid-state oxygen fugacity
buffer which fixes the activity of O2 via a chemical reaction that occurs at run conditions.
Assuming all necessary phases are present at the conclusion of the experiment, one can
assume the activity of O2, and thus the oxygen fugacity was fixed within the buffer.
Many experimental designs were constructed and tested until a successful design was
chosen that controlled all important variables.
6.2
Experimental Designs
The first few experimental designs attempted to use as similar a design as possible
to that described by Yakob et al. (2012), where a Pt-lined Ni capsule was used to grow
olivine and clinopyroxene at varying temperatures. The first design was an attempt to
use a double capsuled experimental setup in the original Pt-lined Ni capsule. A fO2
capsule was created out of 3-mm gold tubing and put inside the experimental chamber.
The tubing was crimped and welded on one end, filled with powder for an oxygen buffer
and ultrapure Milli-Q water, crimped and welded shut. The capsule was weighed before
and after adding every component to ensure material was retained in the capsule after
each step. In these experiments, water is present on both sides of the small inner capsule,
so H+ can become available via the breakdown of water and diffuse through the
semipermeable gold walls to allow for chemical communication between the two
chambers while physically isolating the buffer from the experiment.
These first experiments had a few issues. For one, the amount of material that
could be retained in the small inner buffer capsule was very small. The buffer being used
in these first few experiments was Ni-NiO, and these buffer capsules were completely
62
full of NiO at the end of the experiment. Additionally, Ni escaped the inner capsules and
creating bright green Ni-enriched olivine and clinopyroxene. We tried running
experiments with the oxygen fugacity buffer in the bigger outside chamber and the actual
experiment in the small Au capsule, however we still found large amounts of NiO in the
run products.
Before developing this experimental design further, we came to a revelation that
the method had a major flaw. The experiments were being run in a Pt-lined Ni capsule.
At run conditions, Fe alloys very strongly with Pt, thus we were producing Fe-free
olivine and clinopyroxene. The premise behind our research was that the relative
abundance of Fe3+ as controlled by the oxygen buffer was the controlling factor in Li
partitioning between the mantle phases. If we remove all the Fe from the system, there
are no major redox sensitive elements in the experiment, meaning that controlling the fO2
should not have any effect on Li partitioning.
To address the Fe loss issue, we decided that the best method would be to run the
experiments in a silver capsule with no noble-metal lining. Fe and Ag do not alloy at run
conditions, so the Fe present in the starting materials should remain there throughout the
experiment, and the oxygen fugacity buffer can effectively control the speciation of Fe.
The first Ag capsule design was very similar to its predecessor. An inner Ag-Pd capsule
was constructed in a similar manner as before to serve as the oxygen fugacity buffer
capsule within the system. The Ag-Pd capsule was filled with fayalite, magnetite, quartz
and ultrapure Milli-Q water, placed inside the larger Ag capsule with the experimental
powders. This resulted in growing beautifully large brown fayalite crystals in the
experiment. Fe from the buffer must have escaped the inner capsule, saturating the
experimental charge with Fe and growing Fe-rich olivine.
A new Ag capsule design was constructed where multiple holes were drilled as
close together as possible into the Ag capsule, creating a multi-chambered Ag charge.
The experiment and the oxygen fugacity buffer would go in separate chambers, with the
hope that H+ could still diffuse through the much larger barrier separating the two
chambers. It is unclear if the buffer worked in these experiments, because the
experimental charge as a whole collapsed each time this design was used. The Ag walls
63
were too thin and weak to support the run conditions, and they would fail resulting in a
loss of water to the outside assembly and a failed experiment.
Finally, the inverted capsule technique described in the Methods section above
was developed. This involved having two Ag capsules, one filled with the oxygen
fugacity buffer and one filled with the experiment. The buffer capsule would be inverted
and set on top of the experiment capsule, separated by a thin Ag-Pd foil. The foil
allowed for the chemical communication between the two chambers while preventing
physical contamination.
64
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70
Appendix: Sample Assembly and Piston Cylinder Procedures
The sample assembly consists of (Figure 2.2.2):

2 Ag capsules

Ag-Pd foil

Solid, center drilled, and annulus MgO ceramic pieces

MgO ceramic disk

NaCl tube

Graphite tube

Graphite disk

Pyrex tube

Lead foil

Steel base plug

Pyrophyllite ring

Ceramic thermocouple insulator (“spaghetti”)
Three pieces of “crushable” MgO ceramic is used to surround the Ag capsule.
The bottom piece is solid ceramic, the middle piece resembles an annulus inside which
the crucible sits, and the upper piece has a hole drilled through the center where the
thermocouple can enter and sit directly on top of the sample capsule. To prevent the
metal wires of the thermocouple from coming in contact with the Ag capsule, a thin disk
of MgO ceramic is laid on top of the Ag capsule and MgO ceramic annulus to provide
insulation.
The four pieces of MgO ceramic are all made by hand on a lathe in the lab. The
lathe allows us to shave the outside of the ceramic pieces to the correct diameter, if
necessary, so they can fit snugly inside the rest of the sample assembly. The annulus
ceramic piece is drilled with a stationary drill bit of the correct diameter while the
ceramic piece spins in the lathe, and each individual piece is cut from a longer ceramic
rod using a handsaw.
71
The sample capsule and ceramic insulators are then placed inside a graphite
furnace. The height of the three ceramic pieces should be no longer than the height of the
graphite because the furnace must make complete contact with the baseplug in order for
current to be conducted uniformly to the furnace. A current is passed through the
graphite, which has high electrical resistivity. Current is converted into heat according to
the relationship P = RI2, where P is power (heat transfer in watts), R is electrical
resistance (in ohms) and I is the current (in amperes). A graphite disk of the same
diameter is carefully glued to the bottom of the graphite sleeve in order to allow nearly
complete contact between the two graphite pieces. The graphite disk is in contact with
the piston, which serves to ground the furnace. Surrounding the graphite is a Pyrex
(glass) tube of the same height as the graphite sleeve and disk. The Pyrex tube serves to
physically and electrically insulate the graphite sleeve from the rest of the system, so the
electrical current can only pass through the graphite. The graphite disk is cemented to
both the bottom of the graphite sleeve and the inside of the Pyrex tube, and is left to
harden for a few moments before being inserted into the piston cylinder apparatus.
A salt cell is used to transfer pressure to the sample assembly hydrostatically and
with a minimum of frictional loss. The cell is made of NaCl using a custom die-press.
NaCl powder is poured into the die, which has the shape of a ring, and pressurized so the
powder recrystallizes to form a coherent tube of NaCl. These are prepared ahead of time
and kept in an oven at 100°C to prevent atmospheric water from compromising the
structure of the NaCl tube. The salt cell should be exactly the same length as the sample
assembly. If the cell is too long, it can be ground to the appropriate length using coarse
grit sandpaper. The salt cell is then wrapped in lead foil, which serves as a lubricant as
well as a barrier to prevent salt from oozing into gaps or cracks in the tungsten carbide
core of the pressure plate during an experiment. The lead foil wraps around the salt cell,
but does not cover the entire circumference of the tube; a 2-3mm gap remains. The lead
foil also wraps around part of the bottom end of the salt cell. Should the lead foil be
wrapped completely around the bottom the salt cell, it could potentially come into contact
with the graphite furnace, causing a short in the electrical circuit controlling the
temperature of the system and resulting in an experimental failure.
72
Before loading an experiment into the piston cylinder apparatus, a thermocouple
must be made in order to measure and control the temperature of the sample during the
experiment. Pt wire is threaded through one of the holes of a two-holed piece of ceramic
thermocouple insulator (“spaghetti”) 4.7 cm in length, while Pt/Rh(10%) wire is thread
through the other. Upon exiting the ceramic, the two wires are lightly crimped and
welded using an arc welder to form a bead of metal that rests on top of the thermocouple
ceramic. The other ends of the wire are covered in insulating Teflon tubes to prevent
them from contacting each other or any other metallic surface of the apparatus, which
could cause an electrical short during the experiment. Before being inserted into the
experiment, the freshly welded thermocouple must be annealed over a Bunsen-burner
flame to heal the metal.
When both the sample assembly and the thermocouple are prepared, the
experiment is ready to be put together. First, a (insert correct number here) inch diameter
tungsten carbide “pusher” is inserted into a guide ring. The guide ring serves to center
the pusher on the master ram. The lower cooling plate is centered on the pusher by a
snug-fitting brass sleeve. The cooling plate has input and output quick-release couplings
where hoses can be attached to circulate cooling fluid (water) to cool the pressure vessel
during the experiment. An upper cooling plate will later be aligned and these two are
attached to one another by hoses so water can flow through and cool both the top and the
bottom of the pressure vessel. A ¾-inch tungsten carbide piston is cleaned and placed on
top of the pusher, inside the aligning brass sleeve, and centered using two steel guide
rings. A sheet of Mylar with a ¾-inch center-punched hole is fit snugly around the piston,
extending beyond the edges of the pressure vessel. The Mylar sheet serves as an
insulator to prevent current from flowing through the pressure vessel instead of the
graphite furnace, and also to prevent water from escaping out of the cooling plate. The
core of the pressure vessel is then lubricated with a thin layer of Molykote and placed on
top of the bottom cooling plate, centered on the piston. The inner bore has the same
diameter as the piston, so it is important to lubricate the hole to prevent frictional forces
from building up and resisting pressurization or causing fracturing or chipping of the
brittle tungsten carbide. Surrounding the pressure plate is a cooling ring which is
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connected by hoses to the upper and lower cooling plates. This ring serves to cool the
pressure plate from the outside during the experiment.
In order to load the sample assembly into the piston cylinder apparatus, first the
lead-wrapped salt cell is inserted into the bore of the lubricated pressure plate core. The
sample assembly is inserted carefully inside the salt cell. If using an unwelded capsule
such as the one described above, it is particularly critical to avoid jarring or tiling the
same assembly in order to avoid water loss from the capsule prior to cold-sealing. On top
of the salt cell and sample assembly, but still inside the pressure plate bore, a steel base
plug is inserted and surrounded by a pyrophyllite ring. The base plug serves as a
conductor for the electrical current which flows through the graphite furnace, and as a
soft metal that upon pressurization vertically, expands horizontally and seals the end of
the pressure plate. The pyrophyllite ring insulates the electrical current from the pressure
plate.
The base plug and pyrophyllite ring are then aligned flush with the surface of the
pressure plate by adjusting the master ram, and another Mylar sheet is centered on the
surface. This Mylar sheet serves to prevent water leaks, but also electrical shorts as it
electrically insulates the pressure plate from the upper cooling plate. A ½” hole is
punched in the middle of the Mylar, which allows electricity to flow to the base plug. It
is important to be sure that the hole is centered over the base plug such that no part of the
pressure plate is exposed, as this could result in a short circuit.
The upper cooling plate is then aligned using a series of thin metal rods with
slightly varying lengths and stiffnesses. Holes in the center of the cooling plate, base
plug, and upper MgO crushable ceramic piece have the same diameter as the
thermocouple ceramic, such that they must all be perfectly aligned for the thermocouple
to pass through and rest on top of the sample capsule. When the upper cooling plate is
properly aligned, the thermocouple is inserted into the hole and pressed down until the
entire thermocouple ceramic disappears from view. The insulated wires that emerge
from the top are taped down into a groove cut into the upper cooling plate and led to the
positive and negative leads that connect to the temperature control system. When the
thermocouple is connected, the control system can read the temperature of the system
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based on the ambient EMF generated by the temperature at the nose of the thermocouple.
The EMF changes with temperature, resulting in changes in the voltage read at the
control center. The voltage changes are calibrated with temperature based on the
electrical properties of the individual thermocouple wires, and a specific temperature is
displayed.
A stack of steel spacers are placed on top of the upper cooling plate to fill void
space between the cooling plate and the upper block of the PC frame. When all the plates
are stacked, a 1-2mm gap between the stack and the upper block should remain. At this
point, the sample assembly is complete and pressurization can begin. Initially the end
load ram is raised to close the gap between the plate and the PC frame, after which point
continued pumping pressurizes the pressure vessel, raising the confining pressure on the
sample. The end load pressure is raised to the pressure necessary for the given
experimental conditions (Table A1). Then, the master ram is raised, which pushes the
piston up, causing the sample pressure to increase. The initial pressure increase takes
significant time because the crushable components of the sample assembly, the ceramic,
salt cell, and Pyrex tube, all have to break and compress before the pressure can increase.
Initial pressurization is to about 0.8 GPa, where we assume the Pt lid and the Pt lined
capsule have sealed together. This is important because when we heat the sample, an
incomplete seal between the lid and capsule would allow fluid to escape, and the
experiment would fail (although this may not be apparent until after the run was taken
down). The water hoses are connected to the cooling plates and ring, draining to the sink,
and the water is turned on in a fairly strong, steady stream. A failsafe switch will prevent
power from being supplied to the furnace if the water flow is too low or interrupted.
At this point, we begin bring up the temperature by turning on the power to the
PC. High voltage from the wall-mounted electrical box is converted to high current using
a step down transformer. The current is brought up slowly using a rheostat to avoid
shocking the system. Current is applied from the bottom of the PC, through the master
ram, pusher, piston, graphite furnace, base plug, upper cooling plate, metal stack and into
the PC frame where the electrical circuit is completed. A Eurotherm temperature
controller receives the temperature signal from the thermocouple and supplies current to
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the furnace as necessary to maintain the temperature and the desired setpoint for the
duration of the experiment. We maintain a stable temperature within 1°C throughout the
duration of the experiment.
Throughout the process of increasing the temperature to the set value, we
continue to increase the pressure of the master ram until we reach the desired
experimental pressure (Table A1). After a few minutes of fluctuations in pressure and
temperature where minor manual adjustments to both may be necessary, the system
reaches a steady state and the pressure and temperature remain constant for the duration
of the experiment.
Upon completion of the desired experiment duration, a kill switch is used to
instantly cut off electrical power to the heating system, quenching the experiment. The
temperature typically drops to < 300°C (well past the closure temperature for most high
temperature geochemical exchange) in less than 30 seconds, and <100˚C in 60 seconds.
This rapid cooling ensures that the chemical signature of the crystals grown during the
experiment is representative of what was stable at the experimental run conditions. Once
the sample has cooled to 30˚C, the water system is turned off and a hose is attached to
one of the nozzles and air is forced through the cooling lines to clear out any remaining
water.
The pressure on the master ram and end load ram is released just enough to open a
gap between the spacer stack and the PC frame without allowing any weight to be taken
up by the piston. First, the stack of spacers is removed from the top cooling plate. The
top cooling plate is jolted to the side, snapping the thermocouple ceramic at the cooling
plate – pressure plate interface. The cooling plate is removed and cleaned, dried, and put
away for storage. The Mylar sheet and cooling ring are both removed, and the pressure
plate and lower cooling plate are lifted together to a bench where they can be separated.
The pressure plate is then put below a hydraulic ram where a steel rod pushes the sample
assembly, which is now stuck inside the pressure plate due to radial expansion, out of the
bottom of the plate. Both the pressure plate and lower cooling plate are cleaned, dried,
and put away. The remaining sample assembly is a condensed version of what was put
into the system initially. The salt cell, Pyrex sleeve, graphite tube and ceramic pieces are
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all delicately chipped away from the sample capsule using a small hammer, pliers, and/or
wire cutters so that the sample capsule is all that remains.
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Table A.1 Pressure derivation for the piston cylinder apparatus used in this study
1 inch
Master Ram –
Gauge pressure (bars) /
1 kbar sample pressure
10
End Load Ram –
Gauge pressure (psi) /
1 kbar sample pressure
592
¾ inch
5.625
333
½ inch
2.5
148
Piston diameter
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