Determination of Oxygen Fugacity Using Olivine-Melt Equilibrium

Home , Fayalite, Melt inclusion

Determination of Oxygen Fugacity using Olivine-Melt Equilibrium: Implications for the Redox States of Mid-Ocean Ridge Basalt and Ocean Island Basalt Mantle Source Regions

THESIS

Presented in Partial Fulfillment of the Requirements for the Degree Master of Science in the Graduate School of The Ohio State University

Kenneth James Peterman, B.S.

Graduate Program in Earth Sciences

The Ohio State University

2017

Master's Examination Committee:

Dr. Michael Barton, Advisor

Dr. Berry Lyons

Dr. Tom Darrah

Kenneth James Peterman

2017

Abstract

In order to connect volcanic rocks to their mantle sources, it is essential to consider redox equilibria and their dependence on temperature, pressure, chemical composition, and oxygen fugacity. Oxygen fugacity (fO2) is an intensive variable that strongly affects the behavior of those elements in magmas that are sensitive to changes in redox state, such as Fe, and therefore Mg-Fe silicates, such as olivine. Since fO2 plays an important role in fractional crystallization, in principle, it is possible to estimate fO2 from analyses of olivine in equilibrium with the melt. This research describes a new method based on this principle called the Olivine-Melt Equilibrium Method. This method first calculates Fe3+ and Fe2+ from a relationship involving the partitioning of Mg and Fe2+ between olivine and melt. The ratio of Fe3+/Fe2+ expresses the change in the valence state of Fe, which is related to the redox state of the magma. The calculated Fe3+ and Fe2+ contents of the melt can then be used to determine the fO2 at which magma crystallized from a model described by Kress and Carmichael (1991). This model expresses a

3+ 2+ relationship between the Fe /Fe ratio of the melt, fO2, temperature, pressure, and melt composition.

The Olivine-Melt Equilibrium Method has the advantage that olivine and glass compositions are determined by an Electron Probe Micro Analyzer (EPMA), so analyzed

Fe3+ and Fe2+ of the melt is not required. This is useful because glass analyses in literature typically report all Fe as ∑FeO, rather than distinguishing between Fe2O3 and ii

FeO. Therefore, there is no need for scarce and specialized analytical methods, such as synchrotron-based techniques, to distinguish between the different oxidation states of Fe.

Additionally, this method takes advantage of the fact that olivine is ubiquitous in basaltic lavas, unlike Fe-Ti oxides used to estimate fO2 from geothermometer-oxybarometers.

We have calculated oxygen fugacities from published analyses of coexisting glass and olivine pairs in 982 samples from two different tectonic settings. The results

(expressed as ΔFMQ) for Mid-Ocean Ridge Basalts (MORB) from the Mid-Atlantic

Ridge (−1.55 ± 0.75), the East Pacific Rise (−0.65 ± 0.51), the Juan de Fuca Ridge (−0.77

± 0.42), and the Galápagos Spreading Center (+0.08 ± 0.48) agree with results obtained using other methods and average −1.09 ± 0.89. Ocean Island Basalts (OIB) from Iceland and the Galápagos Islands (ΔFMQ = −0.43 ± 0.71 and −0.33 ± 0.35 respectively) also yield values consistent with those obtained by other methods and fall in the same range as

MORB. However, lavas from the Canary Islands are more oxidized than typical MORB and OIB, with average ΔFMQ = +0.68 ± 0.52. The results for MORB and OIB potentially provide evidence for redox heterogeneity in the mantle, possibly as the result of crustal recycling. However it is necessary to evaluate the possibility that fO2 changes during magma ascent before concluding that the oxygen fugacities of erupted magmas directly reflect those of the mantle source regions.

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Acknowledgments

Foremost, I would like to thank Dr. Michael Barton for his guidance during my graduate studies, and for making insightful comments and revisions to this thesis. I would also like to thank Dr. Berry Lyons and Dr. Tom Darrah for their helpful advice and time spent reviewing this thesis. I would like to thank Friends of Orton Hall and its donors for supporting my research. Without their support, I would not have been able to collect samples from Iceland during the summer of 2016. I would like to thank my brother,

David Peterman, and Collin Oborn for their assistance with sample collection. These samples will tremendously benefit my research in the future. Additionally, FOH assisted with travel funding to attend the 2016 American Geophysical Union Fall Meeting.

Vita

December 24th, 1991 ...... Born, Columbus Ohio

June 2010 ...... Watkins Memorial High School

May 2014 – April 2015 ...... Undergraduate Teaching Associate, School

of Earth Sciences, The Ohio State

University

May 2015 ...... B.S. Earth Sciences, with Research

Distinction, The Ohio State University

August 2015 to present ...... Graduate Teaching Associate, School of

Earth Sciences, The Ohio State University

Fields of Study

Major Field: Earth Sciences

Table of Contents

Abstract ...... ii

Acknowledgments...... iv

Vita ...... v

List of Figures ...... ix

List of Tables ...... xi

Chapters:

1. Introduction ...... 1

1.1 Oxygen Fugacity ...... 2

1.2 Buffers ...... 3

2. Previous Methods for Estimation of Oxygen Fugacity...... 5

2.1 Determination of fO2 using Geothermometer-Oxybarometers ...... 5

3+ 2+ 2.2 Determination of fO2 using Fe /Fe Ratios ...... 7

3. Olivine-Melt Equilibrium – Previous Work ...... 9

3.1 Exchange of MgO and FeO between Olivine and Liquid ...... 9

4. Methodology ...... 13

4.1 Calculation of Oxygen Fugacity from Olivine-Melt Equilibrium ...... 13

5. Geological Settings ...... 18

5.1 Mantle Source Regions of MORB and OIB...... 18

5.2 Mid-Atlantic Ridge Geological Setting...... 20

5.3 East Pacific Rise Geological Setting ...... 21

5.4 Juan de Fuca Ridge Geological Setting...... 22

5.5 Galápagos Spreading Center and Galápagos Islands Geological Setting ...... 22

5.6 Iceland Geological Setting ...... 24

5.7 Canary Islands Geological Setting ...... 25

6. Results ...... 27

6.1 Application to Natural Samples ...... 27

6.2 Results for Each Geological Setting ...... 28

7. Discussion ...... 30

7.1 Comparison between Different Methods for Determining fO2 ...... 30

7.2 Redox Heterogeneity of MORB and OIB Mantle Source Regions ...... 31

7.3 Does the Oxidation State of Basalt Reflect that of its Source? ...... 33

7.4 Comparison of Olivine-Hosted Melt Inclusions and Pillow Glass ...... 39

7.5 Post Entrapment Modifications to Olivine-Hosted Melt Inclusions ...... 41

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8. Conclusions ...... 43

9. Future Work ...... 45

9.1 Proposed Research in Iceland ...... 45

9.2 Analysis of Different Tectonic Environments ...... 46

9.3 Determination of Temporal Variations in fO2 ...... 47

List of References ...... 48

Appendix A: Figures ...... 61

Appendix B: Tables ...... 86

Appendix C: Supplementary Tables ...... 97

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List of Figures

Figure 1. Oxygen Buffer Diagram ...... 62

Figure 2. KD as a Function of Temperature ...... 63

Figure 3. KD as a Function of Oxygen Fugacity ...... 64

Figure 4. Illustration of the Olivine-Melt Equilibrium Method ...... 65

Figure 5. Nd and Sr Isotope Compositions of Each Geological Setting ...... 66

Figure 6. Map of the Mid-Atlantic Ridge ...... 67

Figure 7. Map of the East Pacific Rise ...... 68

Figure 8. Map of the Juan de Fuca Ridge ...... 69

Figure 9. Map of the Galápagos Spreading Center ...... 70

Figure 10. Map of the Galápagos Islands ...... 71

Figure 11. Map of Iceland ...... 72

Figure 12. Map of the Canary Islands ...... 73

Figure 13. Histograms of Results for the Mid-Atlantic Ridge...... 74

Figure 14. Histograms of Results for the East Pacific Rise...... 75

Figure 15. Histograms of Results for the Juan de Fuca Ridge ...... 76

Figure 16. Histograms of Results for the Galápagos Spreading Center ...... 77

Figure 17. Histogram of Results for the Galápagos Islands ...... 78

Figure 18. Histograms of Results for Iceland ...... 79 ix

Figure 19. Histograms of Results for the Canary Islands ...... 80

Figure 20. Distribution of fO2 Calculated using Olivine-Melt Equilibrium ...... 81

Figure 21. Illustration of Crustal Recycling as a Source for Ocean Island Basalts ...... 82

Figure 22. Comparison of KD and ΔFMQ across an Olivine Zoning Profile ...... 83

Figure 23. Comparison of Average Olivine-Hosted Melt Inclusions and Glass ...... 84

Figure 24. Comparison of Corrected and Uncorrected Melt Inclusions ...... 85

List of Tables

Table 1. Olivine-Melt Equilibrium Data for the Mid-Atlantic Ridge ...... 87

Table 2. Fe3+/Fe2+ Ratio Data for the Mid-Atlantic Ridge ...... 87

Table 3. Fe-Ti Oxide Data for the Mid-Atlantic Ridge ...... 88

Table 4. Olivine-Melt Equilibrium Data for the East Pacific Rise ...... 88

Table 5. Fe3+/Fe2+ Ratio Data for the East Pacific Rise ...... 89

Table 6. Fe-Ti Oxide Data for the East Pacific Rise ...... 89

Table 7. Olivine-Melt Equilibrium Data for the Juan de Fuca Ridge ...... 90

Table 8. Fe3+/Fe2+ Ratio Data for the Juan de Fuca Ridge ...... 90

Table 9. Fe-Ti Oxide Data for the Juan de Fuca Ridge ...... 91

Table 10. Olivine-Melt Equilibrium Data for the Galápagos Spreading Center ...... 91

Table 11. Fe3+/Fe2+ Ratio Data for the Galápagos Spreading Center ...... 91

Table 12. Fe-Ti Oxide Data for the Galápagos Spreading Center ...... 92

Table 13. Olivine-Melt Equilibrium Data for the Galápagos Islands ...... 92

Table 14. Olivine-Melt Equilibrium Data for Iceland ...... 92

Table 15. Fe3+/Fe2+ Ratio Data for Iceland ...... 94

Table 16. Fe-Ti Oxide Data for Iceland ...... 94

Table 17. Olivine-Melt Equilibrium Data for the Canary Islands ...... 95

Table 18. Fe3+/Fe2+ Ratio Data for the Canary Islands ...... 95 xi

Table 19. Fe-Ti Oxide Data for the Canary Islands ...... 95

Table 20. S6+/ΣS Ratio Data for the Canary Islands ...... 96

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Chapter 1: Introduction

Oxygen fugacity (fO2) plays an essential role in many chemical processes within the Earth’s mantle and crust. These processes include the speciation of volatiles, evolution of the atmosphere and hydrosphere, metasomatism, magma genesis and magma evolution, and the crystallization of mineral assemblages in igneous and metamorphic rocks (e.g. Kasting et al., 1993). Oxygen fugacity is also an important factor controlling the stability of Fe-bearing silicates, and therefore strongly influences the compositions of melts during magma evolution (e.g. Osborn, 1959). Therefore, oxygen fugacity significantly affects which minerals can crystallize from magmas and which minerals can coexist in igneous rocks.

In principle, it is possible to determine the oxygen fugacity at which a particular magma crystallizes from analyses of the minerals in equilibrium with the melt. This is important because the oxygen fugacities of basaltic magmas reflect those of the mantle source regions (Carmichael 1991), and it is therefore essential to consider redox equilibria and their dependence on temperature, pressure, composition, and oxygen fugacity to understand both the origin and evolution of magmas. The purpose of this thesis is to compare the Olivine-Melt Equilibrium Method with other methods used to calculate oxygen fugacity. Additionally, the goal of this research is to apply this method to natural samples to discern if the mantle is heterogeneous with respect to redox state. 1

1.1 Oxygen Fugacity

In the early 20th century, G.N. Lewis, described a function termed fugacity, which was derived from the Latin root meaning “to escape” (Lewis, 1901). Fugacity is a measure of the escaping tendency of a real gas from a solution. Since oxygen is not necessarily present as a free gas, its reactivity in mineral assemblages is measured

(McCammon 2005).

Oxygen fugacity is a convenient parameter used to monitor the oxidation state of a system. This function measures the availability of oxygen to participate in chemical reactions, and gives information regarding whether Fe, for example, is likely to be found in its native state, as a divalent ion in a silicate mineral, or as a divalent or trivalent ion in an oxide mineral (Lindsley, 1991). Small changes in oxygen fugacity can drastically affect the pathway of magmatic differentiation and of fractional crystallization (Osborn

1959). These different pathways (Liquid Lines of Descent, or LLD’s) resulting from changes of oxygen fugacity reflect changes in mineral assemblages that are mostly the result of the change in the oxidation state of Fe (Philpotts, 2009).

Bowen and Schairer (1935) studied the relationship between oxide minerals and ferromagnesian silicates. They observed that minerals containing Fe2+ crystallize under relatively reduced conditions, whereas minerals containing Fe3+ crystallize under more oxidized conditions. High oxygen fugacities are responsible for early crystallization of magnetite yielding residual liquids depleted in ∑FeO (total iron as FeO) and (relatively) rich in SiO2. Magmas that show these characteristics are said to follow the silica- enrichment trend, or calc-alkaline trend (Frost and Lindsley, 1992). Low oxygen

2 fugacities suppress early crystallization of magnetite and yields residual liquids rich in

∑FeO at nearly constant SiO2, referred to as an iron-enrichment trend, or tholeiitic trend,

(Frost and Lindsley, 1992). The difference between these trends can be appreciated from

Le Châtelier's Principle and the following chemical reaction:

3Fe2SiO4 O2 2Fe3O4 3SiO2 Fayalite + Gas ↔ Magnetite + Quartz ……………………………(Eq. 1)

High fO2 pushes the equilibrium reaction to the right, in order to reestablish equilibrium conditions, thereby resulting in a silica-enrichment trend. Low fO2 pushes the reaction to the left, which favors crystallization of fayalitic olivine and yields the iron-enrichment trend.

1.2 Buffers

Oxygen fugacity varies as a function of temperature and pressure, and therefore comparisons of absolute fO2 need to refer to the same conditions if they are to be meaningful. As a result, oxygen fugacities are normalized to a geochemical buffer. Hans

Eugster first introduced the concept of oxygen controlling equilibria, or oxygen buffers

(Eugster, 1957). The term “buffer” is appropriate because the oxygen fugacities of univariant mineral assemblages are fixed at a single value at fixed temperature and total pressure. This is due to the fact that the equilibrium constants for these redox reactions involving end-member mineral phases can be written as a function only of oxygen fugacity. In other words, as long as the mineral phases are in equilibrium, the oxygen fugacity is fixed at a particular temperature and total pressure (Nordstrom and Munoz,

1994). As the temperature increases, the oxygen fugacity of all buffers also increases

(Figure 1). This reflects the fact that devolatilization reactions occur more readily at higher temperatures, which results in more oxygen escaping, increasing oxygen fugacity

(Lindsley, 1991). The particular buffers shown in Figure 1 involve mineral assemblages or redox equilibria used to control oxygen fugacity in experiments. However, these mineral assemblages rarely exist together in nature. For example, most magmas crystallize near the Fayalite-Magnetite-Quartz (FMQ) buffer (Lindsley, 1991), which is represented by Equation 1. Relatively few igneous rocks contain these minerals coexisting in equilibrium (Philpotts, 2009). Rare occurrences of very Fe-rich igneous rocks may contain this mineral assemblage, such as in the Skaergaard, Bushveld and

Bjerkreim-Sokndal igneous complexes (Barton, personal communication, 2017).

Oxygen fugacity is usually reported as a deviation from the value defined by the

FMQ buffer in log units at the same pressure and temperature. Many tholeiitic igneous rocks crystallize at oxygen fugacities about one log unit below the FMQ buffer (Lindsley,

1991). Few mafic rocks form at two log units below FMQ, and very rare mafic rocks with native Fe crystallize up to greater than four units below the FMQ buffer (Lindsley, 1991).

Most felsic rocks crystallize approximately one to two log units above FMQ, and a few silicic igneous rocks crystalize greater than three log units above FMQ (Lindsley, 1991).

Note that the range in oxygen fugacity is extremely large (~5 log units). Metamorphic rocks have an even wider range of oxygen fugacities compared to igneous rocks due to the fact that there is a much wider range of bulk compositions (Lindsley, 1991).

Chapter 2: Previous Methods for Estimation of Oxygen Fugacity

2.1 Determination of fO2 using Geothermometer-Oxybarometers

The equilibria involved in buffer reactions (Figure 1) illustrate the fact that oxygen fugacity places restrictions on what minerals can coexist and crystalize from certain melts. In other words, oxygen fugacity plays an important role in determining which minerals can crystallize and coexist in a rock. For example, there are no conditions on Earth in which hematite and fayalite can coexist and crystallize together (Philpotts,

2009). Since oxygen fugacity has such a strong effect on the pathways of fractional crystallization, it is possible to determine the oxygen fugacity at which a particular rock equilibrated from its mineral assemblage (Philpotts, 2009).

A number of methods have been used to determine the oxygen fugacities of lavas, including methods based on mineral equilibria involving two-oxides (Buddington and

Lindsley, 1964; Anderson and Lindsley, 1988; Anderson et al., 1991; Ghiorso and Sack,

1991) or oxide-silicate assemblages (Frost et al., 1988; Lindsley et al., 1990; Ballhaus et al., 1990, 1991; Ghiorso and Sack, 1991; Lindsley and Frost, 1992; Frost and Lindsley,

1992). The geothermometer-oxybarometer developed by Buddington and Lindsley

(1964) utilizes coexisting mineral phases which incorporate both Fe2+ and Fe3+ to estimate fO2, and is based on the exchange reactions: 5

Fe3O4 + FeTiO3 ↔ Fe2TiO4 + Fe2O3……………………………………………(Eq. 2)

1 3Fe2O3 ↔ 2Fe3O4 + 2O2…………………………………………………………(Eq. 3)

Further developments of this geothermometer-oxybarometer are described by

Anderson and Lindsley (1988), Anderson et al. (1991), and Ghiorso and Sack (1991).

Another method expands these Fe-Ti oxide geothermometer-oxybarometers to include silicates. This method described by Frost et al. (1988), Frost and Lindsley (1992), and

Lindsley and Frost (1992) can be used to calculate both the equilibrium temperature and oxygen fugacity at which the minerals crystallized. The Quartz-Ulvöspinel-Ilmenite-

Fayalite (QUILF) method is based on the equilibrium:

SiO2 2Fe2TiO4 2FeTiO3 Fe2SiO4 Quartz + Ulvöspinel ↔ Ilmenite + Fayalite ………………………(Eq. 4)

Igneous rocks may contain both the ilmenite-hematite solid solution, referred to as the rhombohedral phase; as well as the magnetite-ulvöspinel solid solution, known as the spinel phase (Philpotts, 2009). Therefore, it is possible to determine the oxygen fugacity and temperature at which these igneous rocks equilibrated, as long as the rock contains both of these phases (Philpotts, 2009). At low oxygen fugacities, the spinel phase is ulvöspinel-rich, and coexists with an ilmenite-rich rhombohedral phase (Philpotts, 2009).

As the oxygen fugacity increases, the spinel phase becomes more magnetite-rich

(Philpotts, 2009). At relatively high oxygen fugacities, the spinel phase is magnetite-rich, and the rhombohedral phase contains hematite (Philpotts, 2009).

The QUILF method allows estimates of oxygen fugacity because oxide and silicate compositions are interrelated. However, this mineral assemblage must be present in order to calculate the oxygen fugacity, although the presence of quartz is not required 6 as long as the activity of SiO2 can be calculated from equilibria involving fayalite and magnetite as illustrated in Equation 1.

These Fe-Ti oxide geothermometer-oxybarometers are capable of providing tight constraints on fO2, as well as equilibrium temperature. However, a major limitation of this method is that the appropriate mineral assemblage is rarely present in the most abundant magmas on earth - Mid-Ocean Ridge and Ocean Island Basalts (Frost and

Lindsley, 1992; Lindsley and Frost, 1992). Even if these assemblages are present, it is highly unlikely that all required components will be analyzed and published in research articles. Therefore it is desirable to develop another method to determine the oxygen fugacities of magmas in order to avoid these complications.

2.2 Determination of Oxygen Fugacity using Fe3+/Fe2+ Ratios

Other methods can be used to determine oxygen fugacity even if no coexisting

Fe-Ti oxides are present (such as in most MORB and OIB). One of methods involves the relationship between the molar fractions of Fe2O3 and FeO in the melt, oxygen fugacity, temperature, and melt composition. Sack et al. (1981) experimentally calibrated this relationship at 1 atm. They found that oxygen fugacity can be related to melt Fe3+ and

Fe2+ ratios by the equation:

Melt XFe2O3 b ln ( ) = a ln(푓O2) + + c + ∑diXi……………………………(Eq. 5) XFeO T where a, b, c, and di are constants determined by step-wise linear regression analysis of experimental data, and Xi is the mole fraction of component i in the melt. Values for the constants are given by Sack et al. (1981). Kilinc et al. (1983) and Kress and Carmichael

(1991) later refined the values of these constants and calibrated this method over a wider range of melt compositions.

This method can be used to estimate oxygen fugacity from glassy samples, yet it requires that the Fe3+ and Fe2+ contents of the melt are determined. Wet chemical techniques were originally used to determine Fe3+ and Fe2+, but this method is time consuming and requires a relatively large sample size. Recently, microbeam techniques such as Mössbauer Spectroscopy or Fe K-edge μ-X-ray Absorption Near Edge Structure

(μ-XANES) Spectroscopy have been used to determine Fe2+ and Fe3+. However, these specialized techniques can be expensive, difficult to use, and are not widely available

(Delaney et al., 1998).

The Fe3+/Fe2+ method cannot be applied to phyric lavas or intrusives, and is only appropriate for glasses and aphyric lavas. Additionally, the proportion of Fe2O3 and FeO of even glassy samples can be altered by chemical weathering, so that this method can only be used for fresh, unaltered glassy lavas (Frost et al., 1988). For these reasons, it

3+ 2+ seems unlikely that use of Fe /Fe to estimate fO2 will find widespread use by petrologists and geochemists in the near future.

Given the difficulties encountered applying the methods described above to natural basalts, it is desirable to develop another method for estimating the oxygen fugacities of lavas. This research utilizes a novel method to determine fO2 based on olivine-melt equilibrium (Barton, 2003; Miller et al., 2005). This new method, the

Olivine-Melt Equilibrium Method, is advantageous because it avoids many of the complications outlined above.

Chapter 3: Olivine-Melt Equilibrium – Previous Work

3.1 Exchange of MgO and FeO between Olivine and Liquid

Roeder and Emslie (1970) experimentally determined the distribution of Mg and

Fe2+ between olivine and melt:

2+ 2+ MgOl + FeLiq ↔ MgLiq + FeOl ……………………………………………(Eq. 6)

They determined the distribution coefficient (KD) that relates the partitioning of Mg and

Fe between olivine and liquid:

Liq XOl X K = FeO MgO = 0.30 ± 0.03……………………………………………(Eq. 7) D Liq XOl XFeO MgO

2+ where Xi refers to the mole fractions of Fe and Mg in olivine and liquid (Roeder and

Emslie, 1970). If the distribution coefficient determined for a particular olivine-melt pair does not fall in this range (0.30 ± 0.03), then it is almost certain that the olivine did not crystallize from the host liquid, and is therefore most likely a xenocryst (Roeder and

Emslie, 1970).

The variation in the value of the distribution coefficient as a function of temperature is given by Roeder and Emslie (1970) as:

130 log(K ) = − 0.7044……………………………………………………(Eq. 8) D T

In their study, variation in KD is relatively small (0.26 to 0.36, or potentially 0.29 to 0.34 when analytical uncertainty is taken into account) at high temperatures (1150 to 1300°C).

This is the consequence of the small value for the numerator (130), which reflects the similar enthalpies of fusion of forsterite and fayalite (Roeder and Emslie, 1970).

Therefore, the distribution coefficient is essentially constant regardless of temperature variations over the range expected for basaltic magmas. Figure 2 also affirms that KD is not dependent on temperature, as it is plotted against the molar fractions of SiO2 –

((Na2O+K2O)/SiO2). The experimental values plot on the line of predicted KD values for two different temperature conditions, demonstrating a lack of temperature dependence on

KD. The experiments performed by Roeder and Emslie (1970) also demonstrated that although the composition of olivine is dependent on oxygen fugacity, the value of KD is essentially independent of oxygen fugacity. This is shown in Figure 3, as the variation in

KD does not fluctuate over a large range of fO2 (>10 log units).

There is a relatively large uncertainty in the value of the distribution coefficient,

(±0.03) (Roeder and Emslie, 1970). This is typical of older analyses of olivine and glass obtained with electron microprobes. Older analyses are less accurate and precise compared to those obtained with modern instruments (Barton, personal communication,

2014). Another source of the uncertainty in the distribution coefficient may stem from uncertainties in the influence of melt composition. In other words, this relatively large variation in KD may reflect the wide variation in chemical compositions used in order to determine it.

Longhi et al. (1978) confirmed the work of Roeder and Emslie, showing that the distribution coefficient was equal to 0.30 ± 0.03, and showed that this value is nearly independent of temperature. They did this by analyzing the distribution of Fe and Mg between coexisting olivine and different lunar basaltic liquids. However, they suggested that the distribution coefficient varies with changes in liquid composition (Longhi et al.,

1978). The value of KD varies with changes in silica concentration, even though it is strictly independent of silica activity by definition (Longhi et al., 1978). Longhi et al.

(1978) suggested that the change in KD with composition is the result of the influence of silica concentration on the mixing properties of Fe and Mg in liquids with different chemistry. They suggested that the Mg activity increases with increasing degrees of polymerization of the melts. These variations in KD can be accounted for by empirical corrections to the expression for KD to take into account variations in silica content

(Longhi et al., 1978).

Gee and Sack (1988) also concluded that the value of the distribution coefficient varies with composition. They combined the results from experiments on multiply- saturated liquids coexisting with olivine, nepheline, leucite, and spinel with or without

Ca-rich pyroxene and melilite at 1 atm pressures with published data from other experiments at 1 atm (Gee and Sack, 1988). They concluded that the value of the distribution coefficient for Mg and Fe2+ between olivine and melt decreases as the compositions of the melts change from tholeiite to alkali basalt, to ugandite, to melilite nephelinite. They defined this compositional dependence of KD using the parameters for pseudoternary liquidus projections in the CMAS system:

S S 2 ln(K ) = −6.06 + 8.104 [ ] − 4.742 [ ] ………………(Eq. 9) D (S+CA+M) (S+CA+M)

Where S is SiO2 – 2(Na2O) – 4(K2O), CA is CaO + Al2O3 + Fe2O3 + 3(Na2O) + K2O –

(P2O5 / 3), and M is FeO + MnO + MgO. Toplis (2005) agreed with Gee and Sack (1988) that it was an oversimplification to state that KD = 0.30 is constant over a range of liquid and olivine compositions, so that melt chemistry does in fact influence the value of KD.

Various KD equations have been tested by Barton (personal communication,

2015) using olivine-melt pairs produced in experiments on natural basalts. This work demonstrates that the expression reported by Gee and Sack (1988) best describes the compositional dependence of KD and allows olivine compositions in equilibrium with melt, and vice-versa, to be calculated accurately (i.e. the error associated with KD is

±0.008 compared to the value of ±0.03 determined by Roeder and Emslie (1970)).

Another equation for KD has been developed by Michael Barton, which is based on the principle that alkalis (Na2O + K2O) influence the polymerization of the silicate melt (Barton, personal communication, 2014). Gee and Sack (1988) represented the compositional dependence on KD using the CMAS system, whereas this KD directly relates to the alkalis in the melt. Barton compared this KD equation to that of Gee and

Sack (1988) and found that calculated values of KD are virtually identical to those calculated using Equation 9, which confirms the effect of alkalis on the melt structure.

Chapter 4: Methodology

This research proposes an alternative, innovative method used to calculate the oxygen fugacity at which basaltic magmas crystallized, and avoids the complications with other methods outlined above. The purpose of this research is to use the Olivine-

Melt Equilibrium Method to determine the oxygen fugacities of magmas from various tectonic settings, including Mid-Ocean Ridge Basalts and Ocean Island Basalts. The objective is to test whether the Olivine-Melt Equilibrium Method yields results that are consistent with those obtained in previous studies using different methods, and to ascertain whether there is significant variation in the redox states of different mantle source regions. If the mantle is heterogeneous with respect to redox state, oxygen fugacity is the first variable to investigate.

4.1 Calculation of Oxygen Fugacity from Olivine-Melt Equilibrium

This method used to calculate oxygen fugacity, originally described by Barton

(2003), has been refined in subsequent work (Barton, personal communication, 2014).

Oxygen fugacity strongly affects the behavior of elements in magmas that are sensitive to changes in redox state. Fe is one of these elements; and therefore the stabilities and compositions of ferromagnesian silicates, such as olivine, are sensitive to changes in redox state. In principle, therefore, it is possible to estimate fO2 from analyses of olivine 13 in equilibrium with melt as long as the compositions of both olivine and melt are known.

This new method is based on this fact: that variation in the melt redox state and oxygen fugacity affect the stability and composition of the coexisting olivine solid solutions

(Barton, written communication, 2014). Lindsley and Frost (1992) and Kress and

Carmichael (1991) have shown that there is little driving force for chemical diffusion of oxygen into or out of the melt phase during crystallization – that is, magmas behave essentially as closed systems with respect to oxygen during crystallization. As magma adiabatically ascends and decompresses in a closed system, the oxidation state of the melt will not change and will reflect the redox state of the mantle source region (Kress and

Carmichael, 1991). Fresh, unaltered, crystal-poor, glassy lavas are the best indicators of the relative oxidation state of mantle source regions because they represent a snapshot of equilibrium conditions of the melt (Kress and Carmichael, 1991). Therefore, this new method can be used to infer the redox states of mantle source regions from calculated oxygen fugacities for crystallization at or near the surface for magmas that behave as closed systems with respect to oxygen during ascent.

Olivine is among the first phases to crystallize from basalts of different chemical compositions. It has a complete solid solution between Fe and Mg end-members: fayalite

2+ (Fe2SiO4) and forsterite (Mg2SiO4), where the olivine has higher Mg/(Mg+Fe ) than the coexisting melt (Roeder and Emslie, 1970). This method takes advantage of the fact that olivine is ubiquitous in Earth’s upper mantle and basalts, contrary to the limited availability of Fe-Ti oxides used in geothermometer-oxybarometers for estimation of fO2.

The Olivine-Melt Method utilizes the olivine-melt Fe2+ – Mg distribution coefficient to calculate the Fe3+/Fe2+ ratio of the melt. This ratio expresses the dominant valence state of Fe, and hence the redox state of the magma. Previous work demonstrates that the KD described by Gee and Sack (1988) (Equation 9) best reproduces experimental data for coexisting olivine-melt pairs. The relationship between Fe and oxygen in the melt can be defined as:

2FeO 1/2 O2 Fe2O3 melt + gas ↔ melt …………………………………………(Eq. 10)

This expresses the effect of oxygen on the speciation of Fe in the melt (Kress and

Carmichael, 1991).

It is necessary to calculate Fe3+ and Fe2+ contents of melts because FeO and

Fe2O3 contents of glassy samples are rarely analyzed. This is because most of the routine analytical instruments, such as the Electron Probe Micro-Analyzer (EPMA), cannot distinguish between Fe2O3 and FeO. As a result, probe analyses report total Fe only

(usually expressed as FeO). However, the actual value of FeO in the melt can be calculated from the composition of the olivine and knowledge of KD. The distribution coefficient is fixed, and the molar fractions of forsterite, fayalite, and MgO in the melt are known from analysis. The FeO content in the melt is the only unknown quantity, and can be determined from the expression of KD. The difference between the total (analyzed) Fe, reported as FeO, and calculated FeO yields the amount of Fe2O3, which is represented by

Equation 11:

FeO + (0.8998)Fe2O3 = FeOTotal

FeO (푎푛푎푙푦푧푒푑)−FeO(푐푎푙푐푢푙푎푡푒푑) Total = Fe O …………………………(Eq. 11) 0.8998 2 3

Hence, the Fe3+/Fe2+ ratio of the melt is known, and can be used to calculate oxygen fugacity using the equation described by Kress and Carmichael (1991) (Equation 5).

Therefore, the Olivine-Melt Equilibrium Method is advantageous because there is no need for specialized analytical equipment (such as Mössbauer Spectroscopy and K-edge

μ-XANES Spectroscopy) to determine the FeO and Fe2O3 contents of the melt.

This method of calculating Fe2O3 using olivine-melt equilibrium is illustrated in

Figure 4. The olivine-melt pair used in this Figure is from an experimental run (#8) on basalt (RE 46-8) from Iceland (Yang et al., 1996). The olivine and glass in this experimental run were analyzed via EPMA, with total Fe reported as FeO. The

° temperature (1198 C) and oxygen fugacity (logfO2 = −8.80) of the experimental run were controlled and reported, with fO2 near FMQ. The value of KD and the mole percent forsterite are used to calculate the amount of FeO in the melt, and this value together with the measured amount of total Fe are used to calculate Fe2O3 in the melt, which is required for use of Equation 5 to determine the oxygen fugacity. Figure 4A displays the relationship between the Mg# (the molar fractions of Mg/(Mg + Fe2+)) and the Fe3+/ΣFe ratio of the melt. The FeO and Fe2O3 in the melt are directly correlated because FeO +

3+ 2+ 0.8998(Fe2O3) = FeOTotal. As Fe increases in the melt (as Fe is oxidized), Fe decreases and the Mg# increases.

Figure 4B shows the relationship between the predicted forsterite mole percent and the Mg# of the melt. The predicted olivine composition in equilibrium with the melt is determined by Equation 9 from Gee and Sack (1988). The composition of olivine 16 varies as a function of the Mg# and therefore as the molar fraction of Fe2+ varies. Since the Mg# and Fe3+/ΣFe are related (Figure 4A), the equilibrium olivine is a function of the

Fe3+/ΣFe ratio in the melt. The true, analyzed, olivine composition can then be used to determine the Fe3+/ΣFe ratio in the melt, as shown in Figure 4C, where the equilibrium olivine is plotted against Fe3+/ΣFe. This value is used to calculate oxygen fugacity using

Equation 5, as shown in Figure 4D. The calculated log(fO2) for this particular experimental run (−8.77) agrees very well with the value controlled during the experiment (−8.80).

Chapter 5: Geological Settings

5.1 Mantle Source Regions of MORB and OIB

Basalt is the primary constituent of spreading centers, as well as hot spot volcanoes. However, these lavas are chemically and isotopically heterogeneous depending on the mantle source region of the erupted basalt. Mid-Ocean Ridge Basalts are erupted from spreading centers sourced from the upper mantle. Ocean Island Basalts are produced via intraplate hotspot volcanism, and are thought to originate from a deeper mantle source, possibly at or near the core-mantle boundary, as mantle plumes. The magmas erupted in these two different tectonic environments have been intensively studied to establish the chemical and isotopic differences and to explore the existence and origin of different mantle sources.

The chemical and isotopic heterogeneity of the Earth’s mantle has been attributed to isotopically depleted and isotopically enriched mantle domains (Zindler and Hart,

1986). The upper mantle is thought to have been depleted by partial melting and crust extraction, while some have suggested that the enriched component of plumes is due to recycling of subducted oceanic lithosphere entrained in convecting mantle (e.g. Hofmann and White, 1982). The mantle source regions for OIB are also heterogeneous at different scales. This compositional heterogeneity is thought to be due to crustal recycling. Seismic 18 tomography data can indicate a larger proportion of high velocity crustal material beneath certain locations (e.g. Hofmann 1997).

The isotope ratios of both Sr and Nd are a useful combination to evaluate the compositional differences between MORB and OIB mantle source regions because the

Rb-Sr isotope system behaves differently during melting than the Sm-Nd system (Faure and Mensing, 2005). During partial melting, Rb preferentially enters the melt phase.

Therefore, volcanic rocks have higher Rb/Sr ratios compared to their source in the mantle. Conversely, during partial melting in the Sm-Nd system, Nd preferentially enters the melt phase. As a result, volcanic rocks have lower Sm/Nd ratios than their source in the mantle. These differences in both isotope systems result in an upper mantle with low

87Sr/86Sr and high 143Nd/144Nd because the upper mantle was depleted in Rb and Nd during partial melting throughout the geological past (Faure and Mensing, 2005). Figure

5 shows the Nd and Sr isotope compositions of magmas erupted in each of the tectonic settings investigated in this research. On average, OIB have higher 87Sr/86Sr and lower

143Nd/144Nd compared to MORB. Although values from Iceland and some of the

Galápagos data overlap with MORB, they are still towards the enriched end of the data array.

Likewise, differences in helium isotope ratios have also provided information about MORB and mantle plume sources. Helium was entrained in the crust and mantle during formation of the Earth by accretion of planetessimals (Faure and Mensing, 2005).

3He is primordial, as it is formed by the β-decay of unstable tritium (3H). 4He is radiogenic, as it is produced through α-decay of U, Th, and their daughters. The ratio of

3He/4He provides information about the ratio of primordial (mantle derived) He to radiogenic (crustal) He. OIB typically display higher 3He/4He ratios compared to MORB, suggesting a more primordial, less degassed source (Faure and Mensing, 2005).

The differences between these isotope ratios in volcanic rocks can be used to infer differences in the mantle source regions. These differences indicate that the Earth’s mantle is mineralogically, chemically, and isotopically heterogeneous, and that OIB and

MORB are derived from different mantle sources. The purpose of this research is to determine whether the mantle is also heterogeneous with respect to redox state, and to establish whether there are systematic differences in redox state between the different mantle sources. This research focuses on seven different locations from two different types of tectonic environment. One environment includes four different mid-ocean ridge localities: the Mid-Atlantic Ridge, the East Pacific Rise, the Juan de Fuca Ridge, and the

Galápagos Spreading Center. The other tectonic environment includes three different ocean island localities of hotspot volcanism. These are the Galápagos Islands, Iceland, and the Canary Islands.

5.2 Mid-Atlantic Ridge Geological Setting

The Mid-Atlantic Ridge (MAR) marks the divergent boundary between the North

American and Eurasian Plates in the northern Atlantic Ocean, and the boundary between the South American Plate and African Plate in the southern part of the Atlantic Ocean

(Figure 6). Many samples have been dredged or collected from various points along the ridge, but most datasets used in this research were produced during the FAMOUS

(French American Mid-Ocean Undersea Study) project between 1973 and 1974 (Bryan 20 and Moore, 1977). Samples were collected by the submersible Alvin, or were retrieved by dredging and core sampling from surface ships during this time. This location was chosen because it is relatively well sampled. Most samples are related to distinctive topographical or geological submarine features, which were located by acoustic navigation systems (Bryan and Moore, 1977). These features were selected in order to collect samples of the most recent fresh lava flows, because these samples are likely to contain fresh, unaltered volcanic glass.

The FAMOUS region is located on the Mid-Atlantic Ridge at 37°47’N latitude.

The ridge segment is approximately 45 km long, and is bounded on the north and south by offsets of 25 km (Laubier et al., 2012). The northern segment is known as the North

FAMOUS segment, and the southern segment is the AMAR segment.

The FAMOUS region has relatively slow, asymmetric spreading rates for mid- ocean ridges, at 7 mm/yr to the west and 15 mm/yr to the east (Bryan and Moore, 1977).

This slow spreading rate may cause intermittent blocking of the volcanic vents when lava accumulates or when the walls of the vents collapse. These phenomena could halt eruptions, which explain the gaps in time recorded by the ages of the lava flows on the central valley of the FAMOUS region (Bryan and Moore, 1977).

5.3 East Pacific Rise Geological Setting

The East Pacific Rise (EPR) is a fast spreading ridge (Herron, 1972) located in the southeastern Pacific Ocean (Figure 7). The EPR is bounded to the north by the Gulf of

California, where it joins the Pacific-North American Plate transform. The southern boundary is a triple junction with the Pacific Antarctic Ridge and the Chile Rise (Searle 21

2013). The EPR represents the divergent boundary between the Pacific Plate and the

North American, Cocos, Nazca, and Antarctic Plates. The full-spreading rate increases from about 60 mm/yr in the Gulf of California to about 160 mm/yr towards the south

(Searle 2013). Because the EPR has a faster spreading rate than many ridges, eruptions are frequent, but not voluminous. Additionally, the faster spreading rate is responsible for the much broader profile of the EPR compared to slower spreading ridges, such as the

MAR (Searle 2013).

5.4 Juan de Fuca Ridge Geological Setting

Figure 8 shows the Juan de Fuca Ridge (JdFR), an intermediate-rate spreading center with a full rate of 58 mm/yr (Van Wagoner and Leybourne, 1991). It is oriented roughly north-northwest in the northwestern Pacific Ocean, approximately 300 km from the coast of North America (Van Wagoner and Leybourne, 1991). It separates the Pacific

Plate and Juan de Fuca Plate, which is a remnant of the former Farralon Plate, currently subducting beneath the North American Plate. The JdFR segment is approximately 500 km long, and is bounded to the north and south by the Sovanco and Blanco Fracture

Zones respectively (Dixon et al., 1986). The Sovanco fracture zone separates the JdFR from the Explorer Ridge to the north, and the Blanco Fracture Zone separates the JdFR from the Gorda Ridge to the south (Dixon et al., 1986).

5.5 Galápagos Spreading Center and Galápagos Islands Geological Setting

The Galápagos Islands are adjacent to, but off the axis of the nearby Galápagos

Spreading Center (GSC), which is located approximately 200 km north of the islands

(Cushman et al., 2004) (Figure 9). It is located in the equatorial Pacific Ocean, and separates the Cocos and Nazca Plates. The GSC trends east-west, and has an intermediate spreading rate, at a full rate of about 45 to 57 mm/yr (Cushman et al., 2004). The GSC is influenced by the Galápagos mantle plume, which is responsible for rising bathymetry due to thermal effects of the plume. The highest point along the GSC, closest to the mantle plume near 91°W, is about 1 km higher than the distal ends of the ridge over a

1,300 km segment length (Cushman et al., 2004). Furthermore, isotope compositions are indicative of plume-ridge interaction. For example, samples dredged from the GSC closest to the western portion of the Galápagos Archipelago exhibit high 87Sr/86Sr and low 144Nd/143Nd ratios, as well as high incompatible element concentrations, relative to

MORB (Cushman et al., 2004).

The Galápagos Islands are located in the equatorial Pacific Ocean, between the

EPR and South America, nearly 1,000 km away from the continent (Figure 10). The

Archipelago was formed via hotspot volcanism, and the current location of the Galápagos mantle plume is beneath the most volcanically active island: Fernandina (Allan and

Simkin, 2000). This research focuses on samples from two different volcanoes

(Fernandina and Santiago) from the Galápagos Archipelago. Lavas from other volcanic centers on the Galápagos have compositions that can be described as mixtures of lavas erupted from these two volcanoes, which represent two distinctive endmembers. Lavas erupted on Fernandina display relatively high 3He/4He ratios suggesting a more primordial mantle source (Koleszar et al., 2009). Additionally, these lavas are enriched in incompatible trace elements relative to MORB. In contrast, lavas from Santiago are

23 chemically heterogeneous but include samples depleted in incompatible trace elements, which closely resemble n-MORB (Koleszar et al., 2009). The compositions of lavas from the Galápagos Islands therefore appear to represent a mixture of magmas derived from ridge and plume mantle sources. This is one of the few places that exhibits magmatism that reflects a mixture of these two mantle endmembers, without one dominating the other (Harpp et al., 2003).

5.6 Iceland Geological Setting

Iceland represents a unique geological setting, located at the intersection of the

Mid-Atlantic Ridge and the Greenland-Iceland-Faeroe Ridge (Figure 11). It has an anomalous elevation up to ~3,000 m above the surrounding sea floor (Thordarson and

Larsen, 2007). Iceland’s crustal thickness varies between 10 km and 40 km (Leftwich et al., 2005). This is extremely unusual compared to normal oceanic crust, which is 7 ± 1 km thick on average (Bown and White, 1994). Only 30% of the landmass is exposed above sea level; however the island and submarine plateau covers about 350,000 km2

(Thordarson and Larsen, 2007).

The oldest rocks on Iceland are only 14 to 16 Ma, however its construction is believed to have initiated at ~24 Ma through multiple eruptive phases (Thordarson and

Larsen, 2007). Seismic and geochemical data suggests the presence of a mantle plume located beneath Iceland (Thordarson and Larsen, 2007). The Iceland mantle plume is hypothesized to have been active for the last 65 Ma, and is responsible for constructing the nearly 2,000 km long North Atlantic Igneous Province. Iceland is the only part of this

24 province that is presently active (Thordarson and Larsen, 2007). The currently active fissure swarms and central volcanoes are displayed in Figure 11.

Hart et al. (1973) showed that 87Sr/86Sr ratios differ as a function of distance along the Reykjanes Ridge towards Iceland. The Reykjanes Ridge is composed of basalt erupted from the upper mantle until about 200 km from the shore of Iceland. Here, the

87Sr/86Sr ratios increase due to a more enriched mantle source related to the Iceland mantle plume (Hart et al., 1973). Not only do Sr isotopes vary systematically along the

Reykjanes Ridge in proximity to Iceland, but 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb ratios do as well (Sun et al., 1975). These ratios systematically increase towards Iceland about 200 km from land, as do 87Sr/86Sr ratios, confirming a change in mantle sources

(Faure and Mensing, 2005). Iceland also has very high 3He/4He ratios, up to 23 times the atmospheric ratio (Ra) (Poreda et al., 1992). These values are greater than that of MORB, and are also characteristic of other OIB, such as Hawaii. These high 3He/4He ratios have been interpreted as indicating that basalts on Iceland are derived from a less degassed, more primordial source than MORB along the southern part of the Reykjanes ridge

(Poreda et al., 1992). This is consistent with evidence from the other isotopic systems mentioned above.

5.7 Canary Islands Geological Setting

The Canary Archipelago (Figure 12) is a chain of seven volcanic islands oriented roughly west-east, stretching 600 km, and is located about 90 km off the coast of northwest Africa (Stroncik et al., 2009). These ocean islands are produced via hotspot volcanism through some of the oldest known oceanic crust, which is about 160 Ma 25

(Nikogosian et al., 2002). Islands are systematically older to the east, from Fuerteventura

(>20 Ma) to El Hierro (<1.7 Ma) in the west (Stroncik et al., 2009). The islands of La

Palma and El Hierro are the youngest islands associated with this hotspot. This is the current locus of the mantle plume, located on the westernmost end of the hotspot track.

Although the island of La Palma is currently in its shield building stage, the lavas are slightly alkalic instead of tholeiitic, which is typical of other ocean islands, such as

Hawaii. Eruptions of mafic magma on top of thicker lithosphere, such as in the Canary

Islands, characteristically produce more silica undersaturated (basanitic) lavas.

Conversely, eruptions of mafic magma on top of thinner, faster moving plates, such as the Hawaiian Islands, typically produce tholeiitic lavas (Nikogosian et al., 2002). The low-flux, low-buoyancy plume beneath the Canaries must rise through thick and old oceanic lithosphere. As a result, primary magmas are low degree, high pressure melts

(Nikogosian et al., 2002).

Chapter 6: Results

6.1 Application to Natural Samples

The datasets used in this research were retrieved from published research articles for each of the geological settings described above. These natural samples used with the

Olivine-Melt Equilibrium Method mostly consist of melt inclusions in olivine, which are small portions of melt that have quenched into an amorphous glassy phase, which olivine trapped during its crystallization. Additionally, all locations contain data from analyzed glassy rims of pillow basalts. This dataset is also composed of a few analyses of interstitial matrix glass, glass crusts in or around xenoliths, tephra glass, hyaloclastite glass, and whole rock analyses. The error associated with the following results using the

Olivine-Melt Equilibrium Method is ~±0.3 log units (Barton, personal communication,

2015).

The Fe3+/Fe2+ ratio model (Equation 5) of Kress and Carmichael (1991) was also used to calculate oxygen fugacities for samples with analyzed values of FeO and Fe2O3

(or Fe3+/Fe2+) for comparison with the Olivine-Melt Method. The samples mostly consist of glasses but some whole-rock analyses were also used. Additionally, values of both temperature and log(fO2) obtained from Fe-Ti oxide geothermometer-oxybarometers

27 were compiled for comparison with results obtained from the olivine-melt method.

These values were retrieved from published papers.

6.2 Results for Each Geological Setting

The Olivine-Melt Equilibrium Method was used to calculate oxygen fugacities from 409 olivine-melt pairs from the Mid-Atlantic Ridge. The average ΔFMQ for these samples is −1.55 ± 0.75. A summary of the specific sample locations, reference, and type of glass present (inclusion or matrix) is displayed in Table 1. Lavas with analyzed values of FeO and Fe2O3 yield an average ΔFMQ (for 69 samples) of −0.76 ± 0.63. Table 2 gives a summary of these results, as well as the method used to analyze the Fe3+ of the melt. A single sample retrieved for the Mid-Atlantic Ridge contains coexisting Fe-Ti oxides and yields ΔFMQ = −0.02 (Table 3). The oxygen fugacities obtained using different methods for the MAR are displayed as histograms in Figure 13.

The Olivine-Melt Equilibrium Method was used to calculate oxygen fugacities from 126 total samples from the East Pacific Rise. The average ΔFMQ for these samples is −0.65 ± 0.51 (Table 4). Lavas with analyzed values of FeO and Fe2O3 yield an average

ΔFMQ of −0.30 ± 0.60 for 87 total samples (Table 5). The average ΔFMQ from coexisting Fe-Ti oxide pairs is −0.35 ± 0.21 (Table 6). These data are shown in Figure 14.

The Olivine-Melt Equilibrium Method yields an average ΔFMQ of −0.77 ± 0.42 for 59 samples from the Juan de Fuca Ridge (Table 7). Analyzed values of Fe3+/Fe2+ in

25 samples yield an average ΔFMQ of −0.20 ± 0.58 (Table 8), whereas Fe-Ti oxide oxybarometry for seven samples resulted in an average ΔFMQ of −0.11 ± 0.35 (Table 9).

The results using these three methods are displayed as a histogram in Figure 15. 28

The Olivine-Melt Equilibrium dataset for the Galápagos Spreading Center consisted of 96 total samples, with an average ΔFMQ of 0.08 ± 0.48 (Table 10).

Analyzed values of Fe3+/Fe2+ in 105 lavas yielded an average ΔFMQ of −0.15 ± 0.50

(Table 11). Fe-Ti oxide data yielded an average ΔFMQ of 0.22 ± 0.45 for 7 samples

(Table 12). These data are represented in Figure 16.

The Galápagos Islands dataset only consists of olivine-melt pairs (Table 13). The average ΔFMQ using the Olivine-Melt Equilibrium Method is −0.33 ± 0.35 for 145 total samples, and is displayed as a histogram in Figure 17.

The Iceland Olivine-Melt Equilibrium dataset consists of 78 total samples. The average ΔFMQ for these samples is −0.43 ± 0.71 (Table 14). Ten lavas with analyzed

Fe3+/Fe2+ data for this location yield an average ΔFMQ of −0.17 ± 0.38 (Table 15), whereas Fe-Ti oxide data for 90 samples average −0.46 ± 0.35 (Table 16). The results obtained using the three methods are shown in Figure 18.

The Olivine-Melt Equilibrium Method was used for 69 samples from the Canary

Islands (Table 17), giving an average ΔFMQ 0.68 ± 0.52. Analyses of Fe3+/Fe2+ for 58 samples yielded an average ΔFMQ of 0.47 ± 0.56 (Table 18); and coexisting Fe-Ti oxides in 88 samples were used to calculate an average ΔFMQ of 0.83 ± 0.80 (Table 19).

Additionally, S6+/ΣS ratios were used to estimate oxygen fugacities for 49 samples, averaging at 0.86 ± 0.68 above FMQ (Table 20). Results for all four of these methods are summarized in Figure 19.

Chapter 7: Discussion

7.1 Comparison between Different Methods for Determining fO2

As seen in Figures 13 through 19, the results obtained using the Olivine-Melt

Equilibrium Method are virtually identical with results obtained using other methods used to calculate fO2. These other methods include various Fe-Ti oxide geothermometer- oxybarometers and analyzed Fe3+/Fe2+ ratios of glassy samples. In addition, the results agree with the typical values of ΔFMQ for MORB and OIB reported in literature. A large majority of the results from each of these locations range from just above to about a log unit below the FMQ buffer, which is typical for MORB and OIB magmas (e.g. Ballhaus,

1993; Lindsley, 1991; Frost et al., 2004).

Some of the Fe3+/Fe2+ ratio results for certain locations studied in this research, such as the Juan de Fuca Ridge and EPR, yield slightly higher fO2 compared to the

Olivine-Melt Equilibrium Method. Many of these datasets were retrieved from Cottrell and Kelley (2011), who determined Fe3+/ΣFe, via Fe K-edge μ-XANES spectroscopy.

They suggest that MORB may be slightly more oxidized than previously thought.

Recently, Berry et al. (2015) suggests that there may be a calibration error responsible for these high Fe3+/ΣFe, and if this is true, then this may explain the apparent difference between the Olivine-Melt Equilibrium Method and some of the slightly higher Fe3+/Fe2+ 30 estimates. Although, the values are similar when the error in the Fe3+/Fe2+ method is taken into account, which is about ±0.5 log units. However, these results are not significantly higher if these uncertainties are taken into account, so they will be accepted until future studies prove that these values were incorrect.

7.2 Redox Heterogeneity of MORB and OIB Mantle Source Regions

It has been shown that lavas from hotspot volcanoes originate from a chemically and isotopically different mantle source compared to mid-ocean ridges (Zindler and Hart,

1986). Therefore, they should also reflect any differences in thermodynamic properties, such as oxygen fugacity, between mantle source regions. All of the results presented in this research yield oxygen fugacities generally near the FMQ buffer. However, there are slight differences between each location. The average values calculated using olivine- melt equilibrium vary by about 2.23 log units at the most, between the Mid-Atlantic

Ridge (most reduced) and the Canary Islands (most oxidized). Figure 20 shows the ranges and averages of fO2 calculated using the Olivine-Melt Equilibrium Method for each location. This displays that on average, the three OIB localities are slightly more oxidizing than the all MORB localities, except for the Galápagos Spreading Center. This plume-influenced spreading center is more oxidized on average compared to normal

MORB. When considering the error in the Olivine-Melt Equilibrium Method (~0.3 log units), the Iceland and Galápagos averages do, however, fall within the range of MORB averages. The Canary Islands data are clearly different, over a log unit more oxidizing than any of the other normal MORB localities on average. Others have also noted this difference in calculated fO2, suggesting that OIB sources can be slightly more oxidized 31 than MORB (Ballhaus, 1993; Haggerty and Tompkins, 1983; Williams et al., 2004;

Oppenheimer et al., 2011; Green and Falloon, 2015). However, there is still uncertainty regarding this evaluation (Gerlach, 2004a; Roeder et al., 2004). It is therefore evident that it is necessary to better constrain the redox state of magmas from all tectonic environments.

It has been proposed that the redox state of basalts (especially MORB) directly reflect the redox states of their mantle source regions (Carmichael and Ghiorso, 1986;

Carmichael, 1991; Wallace and Carmichael, 1992; Bézos and Humler, 2005; Lee et al.,

2005; Cottrell and Kelley, 2011). If the oxygen fugacities calculated from these basalts truly represent their mantle source regions, then this would suggest that the deeper mantle source regions of certain locations, such as the Canaries, are more oxidized than the upper mantle source of spreading centers. This is unexpected, as the plume hypothesis suggests that mantle plumes may originate from the lower mantle, at or near the core- mantle boundary. Indeed, the lower mantle likely contains some native Fe, requiring oxygen fugacities several log units below FMQ (Frost et al., 2004). It has been proposed that relatively oxidized, subducted oceanic lithosphere may accumulate in the lower mantle, being recycled into these mantle plumes (e.g. Arculus, 1985; Haggerty and

Tompkins, 1983; Hofmann, 1997; Kellogg et al., 1999; Green and Falloon, 2015; illustrated in Figure 21). Therefore, these higher oxygen fugacities could be the consequence of crustal recycling in the source for OIB.

Crustal recycling may also explain the chemical and isotopic differences between

MORB and OIB source regions. However it is unclear how this process could reproduce

32 the relatively oxidized character of the Canaries because oceanic lithosphere and asthenosphere is relatively reduced. The thin veneer of marine sediments is more oxidized than the oceanic crust, but does not contribute nearly as much by volume.

Moreover, the subcontinental lithosphere is more oxidized than MORB, although similar in composition (Wood et al., 1990). The process of continental lithospheric delamination during subduction may possibly contribute a larger amount of oxidized crustal material, although it is still unclear how this process could create such oxidized magmas as the

Canaries.

7.3 Does the Oxidation State of Basalt Reflect that of its Source?

The oxygen fugacity, and hence redox state, of a melt will reflect that of its source region as long as it is a closed system with respect to oxygen (Kress and Carmichael,

1991). However, several processes may alter fO2 of ascending magma relative to that of its mantle source region. These include partial melting, decompression, fractionation of the melt during ascent, the exchange of oxygen with the ambient mantle, and degassing

(Ballhaus, 1993). The effects of post-eruptive oxidation and volatile loss are minimized in submarine lavas erupted at mid-ocean ridges because these lavas are immediately quenched, unlike subaerially erupted lavas (Wallace and Carmichael, 1992). Therefore, subaerially erupted OIB are more likely to be affected by post-eruptive oxidation and volatile loss, which may explain their higher oxygen fugacities relative to MORB.

Moreover, the effects of degassing may be lessened at MOR because they often reside under 2.5 to 3.5 km of water, which exerts a significant pressure and inhibits degassing

(Carmichael and Ghiorso, 1990). Additionally, for OIB samples, the location and method 33 at which samples are collected is crucial for accurately determining their oxidation state.

The fO2 of subaerial lavas can differ if they are quenched close to volcanic vents compared to those that have quenched after traveling some distance away from the vent

(Rhodes and Vollinger, 2005).

The process of contamination may oxidize magmas, depending on the contaminant. A melt could be oxidized if magmas rise through older and thicker lithosphere that has been metasomatically altered, and therefore has been oxidized.

Additionally, hotspots located closer to continental margins must pass through thicker packages of relatively oxidized sediment eroded from nearby continents (Thirlwall et al.,

1997). This could explain the Canary Islands results, which are the most oxidized OIB analyzed in this research. The anomalous lithospheric thickness below the Canaries is a consequence of the relatively slow movement of the African Plate across the hotspot, and they are located on old (~160 Ma) crust (Nikogosian et al., 2002). The crust is about 16 km thick, and consists of about 6 km of sediments derived from northwest Africa.

Additionally, there are many complex, chemically zoned phenocrysts, as well as igneous textures, that reflect replenishment and magma mixing below the Canary Islands

(Stroncik et al., 2009). Thus, crustal contamination could potentially explain the higher oxygen fugacities of Canary Island samples compared to other OIB or MORB in this research.

Contamination at shallow crustal depths is thought to be a common process in

Canary Island magma chambers, especially during the shield building stage (Gurenko et al., 2011). Thirlwall et al. (1997) suggested that magmas from their study had assimilated

34 up to 8% of sediments from the northwest African passive margin. However, Gurenko et al. (2011) show that there is no evidence for magma contamination by lower crustal rocks in the samples they studied; therefore they reflect the true characteristics of their mantle source regions. It has been shown that the geochemistry of lavas from the eastern Canary

Islands (Gran Canaria, Fuerteventura, and Lanzarote, and certain massifs on Tenerife) can be explained as a mixture between Depleted MORB Mantle (DMM) with HIMU-type sources (“high μ”; i.e., high 238U/204Pb ratios), and Enriched Mantle (EM) type asthenosphere or delaminated African subcontinental lithosphere (Gurenko et al., 2011).

Oxygen isotope data from olivines from their study are highly variable, with δ18O values ranging from 4.6 to 6.1‰. Values below 5‰ are attributed to assimilation with hydrothermally altered crust, whereas values higher than this are considered to represent a normal mantle source (Hemond et al., 1993). This suggests contributions from recycled lithosphere, not from lower crustal material assimilated upon magma ascent. Although some samples may have assimilated crustal material, other samples have compositions that reflect those of their mantle source regions. Therefore, the more oxidized nature of lavas from the Canaries is probably not due to contamination, but rather reflects the redox states of their mantle source regions, unless they result from other processes, such as degassing.

There is no doubt that assimilation of relatively oxidized crustal material could alter the redox state of a basaltic magma. However, degassing may also play a role in changing the redox state of magma upon ascent. Many basalts are vapor saturated, which results in extensive degassing before eruption (e.g. Moore, 1979; Dixon et al., 1988).

Christie et al. (1986) observed that the whole rock cores of pillow basalts were significantly more oxidized compared to the rapidly quenched pillow rims. This was attributed to degassing of the cores relative to the rims. Some research has suggested that the degassing of H2O and loss of H2 are responsible for oxidizing a melt (Sato and

Wright, 1966; Sato, 1978; Mathez, 1984). This is represented by the chemical reaction:

H2O + 2FeO ↔ Fe2O3 + H2……………………………………………(Eq. 12)

However, Waters and Lange (2016) argue that H2O degassing does not affect fO2 of silica-rich magmas in natural settings. This is supported by Moussallam et al. (2014;

2016) under the conditions investigated in their study. Moreover, MORB and OIB typically have very low bulk H2O content (< 0.40 wt%), resulting in negligible H2O activity. Therefore, H2 diffusion is an ineffective mechanism for oxidizing the melt

(Carmichael, 1991).

The degassing of carbon monoxide has also been proposed to oxidize the melt

(Mathez, 1984), whereas the degassing of sulfur species can either oxidize or reduce melts. This depends on the particular sulfur species in the melt (S2- or S6+), as well as the fluid phase (H2S or SO2) (Carmichael, 1991; Gaillard and Scaillet, 2009; Gaillard et al.,

2011; Kelley and Cottrell, 2012; Moussallam et al., 2014). These processes are not viable under mantle pressures, as they only operate within the crust near the surface (Gerlach,

1993).

Theoretically, sulfur degassing is capable of influencing the redox state of the melt (Carmichael and Ghiorso, 1990; Gaillard and Scaillet, 2009; Gaillard et al., 2011)).

This has also been shown in nature, where the degassing of sulfur was found to be

36 responsible for reduction of magma upon ascent at Mt. Eerebus (Moussallam et al.,

2014), as well as at Kilauea (Moussallam et al., 2016), represented by:

melt melt gas melt S + 2Fe2O3 = SO2 + 4FeO …………………………………………(Eq. 13)

Therefore, Moussallam et al. (2014; 2016) propose that the redox state of the mantle underneath most hotpot volcanoes is even more oxidized than that of lavas produced by them. Additionally, they suggest that the extent of sulfur degassing determines whether or not the oxygen fugacity of basalt truly represents the redox state of its mantle source region. MORB typically have slightly lower pre-eruptive dissolved sulfur contents compared to OIB (800 – 1200 ppm vs. 1000 – 2000 ppm respectively) at similar FeO contents (Oppenheimer et al., 2011). As a result, the redox states of OIB may be affected to a greater extent than those of MORB by sulfur degassing. If this is true, the mantle source regions of OIB in this research may be even more oxidized than calculated using the Olivine-Melt Equilibrium Method and other methods discussed earlier. This proposal has important ramifications regarding the calc-alkaline differentiation trend vs. the tholeiitic trend. If OIB mantle sources were relatively oxidized, then this would suggest that all crystallization to produce the tholeiitic differentiation trend would have to occur after degassing – a shallow level process.

It is important to note that the lavas from the Canaries can have unusually high sulfur concentrations for OIB (up to ~5,810 ppm; Gurenko and Schmincke, 2000).

Typically, these high sulfur magmas are the most oxidized, suggesting that sulfur reduction is the dominant degassing process that affects fO2 for lavas from the Canaries.

It has been postulated that three different models could explain the origin of these sulfur-

37 rich magmas (Gurenko and Schmincke, 2000). The first is that these magmas are juvenile in origin, and that partial melting of the mantle source is responsible for producing such high sulfur contents. However, to dissolve such large amounts of sulfur, the mantle source of the magmas must be much more oxidized than that of typical MORB or OIB.

This is due to the fact that the solubility of sulfur is a function of oxygen fugacity, where higher oxygen fugacities are required to dissolve higher concentrations of sulfur

(Gurenko and Schmincke, 2000). The second is that the high sulfur concentrations may represent assimilation of old oceanic crust and sediments, as sulfate dissolved in seawater has been shown to form anhydrite at temperatures between 150 – 200°C. The third possible scenario involves interaction with seawater. These high sulfur content melt inclusions have relatively high Cl/K2O ratios, likely derived from seawater contamination.

Taking these factors into consideration, the calculated fO2 of OIB may not completely represent the redox states of their mantle source regions. However, the

Canary Islands are clearly derived from a more oxidized source than the other geological settings investigated in this research. Since some samples do not show chemical or isotopic evidence for contamination, and they do not show evidence of oxidation due to degassing, the oxidized nature of the Canaries likely does reflect that of its mantle source region. Therefore, it is still unclear as to why the Canary Island lavas are more oxidized, but it cannot be completely ruled out that the process of crustal recycling may contribute to the relatively oxidized nature of the Canaries.

7.4 Comparison of Olivine-Hosted Melt Inclusions and Pillow Glass

Most of the datasets processed with the Olivine-Melt Equilibrium Method consist of olivine-hosted melt inclusions. However, there are other forms of glass for every location. These mostly include pillow basalt rims, hyaloclastite glass, glass in gabbroic nodules, or matrix glass. It is important to determine if the olivine crystals coexisting with these glasses truly represent equilibrium compositions. To do this, the olivine-melt distribution coefficients (KD) were analyzed for each sample. If the calculated KD did not fall in the range of 0.30 ± 0.03, then the olivine-melt pair is not in equilibrium and the olivine potentially represents a xenocryst that was incorporated in the melt. Therefore, these data were excluded from this research.

The most reliable datasets were retrieved from published research articles that analyzed olivine compositions in more detail. For example, it is very useful to have an olivine zoning profile. Olivine phenocrysts are frequently chemically zoned, where the core is more forsteritic compared to the rim, which is typical for normally zoned crystals.

The KD values calculated in this research indicate that the olivine rim best represents equilibrium conditions, so these were selected and paired with glasses of the same sample in this research. Figure 22 shows how the calculated ΔFMQ and KD can vary across a normally zoned olivine crystal, and that the rim best represents equilibrium between the olivine and coexisting glass.

Published research articles have yielded many datasets for this research, however it can be difficult to acquire all of the required compositional data for volcanic glass and coexisting olivine crystals from some articles. This is a complication because many

39 research articles do not contain both analyses of olivine and glass, or only include analyses of olivine macrocryst cores. In many articles it is not specified where the olivine is being measured, whether or not it is a core or rim, or if normal or reverse zoning is present.

Figure 23 shows a comparison of the average ΔFMQ values, calculated using olivine-melt equilibrium, for both olivine-hosted melt inclusions and glasses paired with olivine crystals. This shows that on average, both analyses yield similar results, never exceeding a log unit in difference. However, it is apparent that samples of matrix or groundmass glass typically yield slightly higher values compared to melt inclusions. This could be due to the difficulties described above when pairing olivine and glass, especially if there was no careful analysis of olivine-zoning profiles. If only cores are reported, the olivine would likely be more forsteritic than the true equilibrium value, and this would yield anomalously high oxygen fugacities. Additionally, this observation could be the result of a sampling bias, as many datasets used in this research contain much more melt inclusions compared to other glass. Unlike the other datasets, the Iceland data contains more glasses than melt inclusions. These were retrieved from research articles that typically contain olivine zoning profiles, or specify where the olivine is being measured.

Consequently, results for olivine-rim matrix-glass pairs are nearly identical to the average of the melt inclusion results (ΔFMQ of −0.40 and −0.48, for 45 and 33 samples respectively). Therefore, the other glass data may likely yield results more consistent with melt inclusions if there were more samples available, and if there were more analyses of individual olivine crystals. However, it is also possible that this difference is due to the

40 method used to correct for post entrapment modifications to olivine-hosted melt inclusions, as they suffer from Fe-loss.

7.5 Post Entrapment Modifications to Olivine-Hosted Melt Inclusions

Another factor to consider is how the olivine-hosted melt inclusion was corrected for post entrapment equilibration with the host olivine. As a melt inclusion cools, after being trapped in olivine during crystallization, a rim of olivine crystallizes on the inclusion walls (Danyushevsky et al., 2002). As this olivine rim crystallizes, Fe diffuses into it from the melt inclusion, while the rim loses Mg into the inclusion. This results in a compositional gradient within the rim. This compositional gradient is responsible for re- equilibration of the melt inclusion with its host olivine. This process is completed by diffusion of Fe out of the inclusion, and diffusion of Mg into the inclusion. This process is referred to as “Fe-loss” by Danyushevsky et al. (2002). On the contrary, the process of

“Fe-gain” is caused by heating of the melt inclusion greater than the temperature of entrapment. This results in melting of the host olivine around the inclusion, which would increase the amount of Mg inside the inclusion, causing disequilibrium between the melt and olivine host. Therefore, the process of re-equilibration would occur, which is accomplished by diffusion of Fe into the inclusion and the diffusion of Mg out of the melt inclusion (Danyushevsky et al., 2002).

Various correction methods have been used to account for this process, such as iteratively adding olivine back into the melt at 0.1% increments or correcting to a known

FeO content. Figure 24 illustrates the importance of correcting for Fe-loss, as the uncorrected melt inclusions yield oxygen fugacities several log units higher than is 41 expected. Additionally, the calculated values exhibit much scatter, varying almost 5 orders of magnitude. Moreover, the KD values clearly indicate disequilibrium. All of the melt inclusions used in this research were corrected for these effects. The best melt inclusions are those that are larger, with lower surface area to volume ratios. These are less affected by Fe-loss. Additionally, it is favorable to use corrected melt inclusion data in which it is necessary to add less olivine back into the melt, because these were less affected by Fe-loss. As described above, the melt inclusion data is very comparable to the other glass data, suggesting that these have been appropriately corrected for post- entrapment modifications.

Chapter 8: Conclusions

This research project utilizes a new method to determine the oxygen fugacity at which basaltic rocks crystallize. The Olivine-Melt Equilibrium Method has yielded results for four different mid-ocean ridge localities and three ocean islands created via hotspot volcanism. The results clearly show that the Olivine-Melt Equilibrium Method calculates oxygen fugacities that are very consistent with other methods, including Fe-Ti oxide geothermometer-oxybarometers, as well as models based on Fe3+/Fe2+ ratios of the melt. However, this method has the advantage that it can be applied to basalts (which often lack Fe-Ti oxides). Additionally, this method does not require difficult or unusual analytical methods, and there are abundant datasets available for olivine-melt pairs from several different tectonic environments.

The oxygen fugacities of basalts calculated from MORB and OIB differ in this research, which may suggest variable redox states of their mantle source regions. This could potentially have major implications for models of mantle evolution and crustal recycling, as the latter process in principle could explain the slightly more oxidizing conditions of the Canary Islands. However, it is necessary to acquire more data in order to accurately address this question, as oxygen fugacity can be modified upon ascent of the magma. This research affirms that it is important to continue to investigate the redox

43 states of many different tectonic environments in order to better constrain the redox states of MORB and OIB.

Chapter 9: Future Work

9.1 Proposed Research in Iceland

Since the goal of this study is to apply the Olivine-Melt Equilibrium Method to natural samples, and compare its results with other methods, it is proposed to collect and analyze samples directly, instead of only obtaining compositional data from published research articles. Samples have been collected for analysis from Undirhlíðar quarry on the Reykjanes Peninsula, as well as Midfell in the Hengill Region of Iceland during the summer of 2016. Thin sections will be cut from these samples to include the pillow rim glass, as well as the glass halos and interstitial glass inside the gabbro xenoliths from the

Midfell samples. A preliminary analysis of these samples will include observing zoning profiles using a Scanning Electron Microscope (SEM) at The Ohio State University.

These samples will later be analyzed via Electron Probe Micro-Analysis (EPMA) to acquire compositional data of olivine and glass. Then these data will be processed with the Olivine-Melt Equilibrium Method in order to determine the oxygen fugacity at which the samples crystallized.

These samples will be used to better constrain the oxygen fugacity at which

Icelandic rocks crystallized, and provide information about the redox states of their mantle source regions. Furthermore, these samples could also be used to compare the 45

Olivine-Melt Equilibrium Method with models based on Fe3+/Fe2+ ratios. Moreover, the proportion of the different valence states of vanadium could potentially be analyzed. This would provide a comparison with methods involving V/Sc ratios, which suggest a higher redox state than is indicated by oxygen fugacities calculated using olivine-melt equilibrium. These samples could also be studied in more detail to provide information regarding the role of degassing.

9.2 Analysis of Different Tectonic Environments

It is proposed to acquire more data from published research articles and calculating fO2 from different MORB and OIB localities in order to develop a better global distribution of datasets. Particular mid-ocean ridges of interest include the

Southwest, Southeast, and Central Indian Ridges, the Gakkel Ridge, and the Pacific

Antarctic Ridge. More hotspot volcanoes will also be considered, including the Hawaiian

Islands, the Azores, Réunion, and possibly other ocean islands in the Pacific Ocean.

Oxygen fugacities have already been calculated from many MORB and OIB samples. However, the Olivine-Melt Equilibrium Method has been applied to only a few

Continental Flood Basalts (CFB) and Island Arc Basalts (IAB). Data from east Greenland and the Deccan Traps yield similar oxygen fugacities to OIB. The purpose of applying this method to IAB is to test the hypothesis that magmas from volcanic arcs are derived from a relatively oxidized source compared to MORB and OIB. Recent work based on

V/Sc ratios (Lee et al., 2005) has proposed that this hypothesis is incorrect, and that different tectonic environments have very similar redox states.

To apply this method to volcanic arcs, it is then necessary to take into account the amount of H2O present in the melt, as volcanic arcs contain substantially more water compared to MORB and OIB (Stolper and Newman, 1994). Furthermore, the pressures of crystallization must also be considered because most arc magmas begin crystallizing at mid-crustal depths. Therefore, it is essential to use experimental datasets to determine if there is any effect of dissolved H2O and high pressure on the olivine-melt distribution coefficient. If these do affect the KD, then the equation could be modified to account for these variables. Seth Bryson has already begun work to determine the potential effect of pressure and dissolved water on KD.

9.3 Determination of Temporal Variations in fO2

Not only will this research focus on spatial variations in oxygen fugacity, but also on possible temporal variations in the oxygen fugacities of erupted magmas. The Olivine-

Melt Equilibrium Method could be applied to Archean and Proterozoic komatiites, and the results compared with modern samples to discern if there has been any overall change in mantle redox state through time. There is debate regarding the redox state of the

Archean mantle, which has implications for understanding fundamental Earth processes.

For example, this research could provide insight into the role of crustal recycling during the Archean. Additionally, the redox state of the mantle influences the speciation of volatiles, and therefore impacts the composition of Earth’s atmosphere and hydrosphere that formed through the process of volcanic degassing (Kasting 1993). There has been debate regarding the composition of Earth’s early atmosphere, and this has ramifications for our understanding of the development of life on Earth (Kasting 1993). 47

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Appendix A: Figures

Figure 1: Oxygen Buffer Diagram. log(fO2) vs. temperature for various buffers at one bar based on specific algorithms. These buffers include: MnH: manganosite-hausmannite (3MnO + ½ O2

↔ Mn3O4) equation from Chou (1978); MH: magnetite-hematite (4Fe3O4 + O2 ↔ 6Fe2O3) equation from Heubner (1971); NiNiO: nickel-nickel oxide (2Ni + O2 ↔ 2NiO) equation from

Heubner and Sato (1970); FMQ: fayalite-magnetite-(β)quartz (3Fe2SiO4 + O2 ↔ 2Fe3O4 + 3SiO2) equation from Heubner (1971); WM: wüstite-magnetite (3Fe1-xO + O2 ↔ Fe3O4) equation from

Heubner (1971); IW: iron-wüstite (2(1-x)Fe + O2 ↔ 2Fe1-xO) equation from Heubner (1971);

QIF: (β)quartz-iron-fayalite (2Fe + SiO2 + O2 ↔ Fe2SiO4) equation from Heubner (1971).

Figure 2: KD as a Function of Temperature. KD is plotted against the molar fractions

(X) of SiO2 minus alkalis (Na2O + K2O) divided by SiO2 for two different temperature conditions (1200 and 1100°C) at fO2 near that of the FMQ buffer. This demonstrates that although KD varies as a function of silica and alkalis in the melt, the experimental values plot on the predicted KD line even though temperatures vary. Therefore, this shows little effect of temperature on KD. (Barton, personal data, 2014).

Figure 3: KD as a Function of Oxygen Fugacity. ΔFMQ vs. ΔKD (experimental KD – average KD for the same bulk composition). This shows that KD does not vary over a large range of fO2 (>10 log units). (Barton, personal data, 2014).

Figure 4: Illustration of the Olivine-Melt Equilibrium Method. Relationship between

Mg# and Fe3+/ΣFe in the melt (A) and between predicted equilibrium olivine composition and Mg# (B). (C) shows the relationship between the olivine and Fe3+/ΣFe in the melt, as well as the true analyzed olivine composition, which is used to constrain the Fe3+/ΣFe ratio. This in turn is used to calculate fO2, as shown in (D). This example uses data from an experimental run (#RE 46-8) from Yang et al. (1996). 65

Figure 5: Nd and Sr Isotope Compositions of Each Geological Setting. Data for the

Juan de Fuca Ridge (JdFR), Iceland, the Galápagos Islands, and the Canary Islands were retrieved from a compilation of literature sources from the petrological database: PetDB

(http://www.earthchem.org/petdb; Lehnert et al., 2000). The Mid-Atlantic Ridge (MAR) and East Pacific Rise (EPR) fields were developed after N-MORB data from Ito et al.

(1987) and LeRoex et al. (1983). On average, the three OIB localities are enriched in

87Sr/86Sr and depleted in 143Nd/144Nd relative to MORB. However, Iceland largely falls into the range of MORB, whereas the Galápagos Spreading Center forms the top-left part of the Galápagos data array, and the Galápagos Islands (bottom-right part of array) range to more typical OIB.

Figure 6: Map of the Mid-Atlantic Ridge (MAR), also showing the location of the

FAMOUS region, the Azores, the Charlie-Gibbs Fracture Zone (CGFZ), which separates the MAR from the Reykjanes Ridge (RR). Created with GeoMap

(http://www.geomapapp.org) (Ryan et al., 2009).

Figure 7: Map of the East Pacific Rise (North and South). Also shown are the Pacific-

Antarctic Ridge (PAR), the Chile Rise (CR), the Galápagos Islands (GI), and the

Galápagos Spreading Center (GSC). Created with GeoMap (http://www.geomapapp.org)

(Ryan et al., 2009).

Figure 8: Map of the Juan de Fuca Ridge (JdFR). Also shown are nearby ridges and fracture zones, including the Mendocino Fracture Zone (MFZ), the Gorda Ridge (GR), the Blanco Fracture Zone (BFZ), Endeavour Ridge (ER), Sovanco Fracture Zone (SFZ),

Explorer Ridge (ExR). Created with GeoMap (http://www.geomapapp.org) (Ryan et al.,

2009).

Figure 9: Map of the Galápagos Spreading Center (GSC). Also shown are nearby

Galápagos Islands (GI) and East Pacific Rise (EPR). Created with GeoMap

(http://www.geomapapp.org) (Ryan et al., 2009).

Figure 10: Map of the Galápagos Islands. IF and IM refer to Isla Floreana and Isla

Marchena. Created with GeoMap (http://www.geomapapp.org) (Ryan et al., 2009).

Figure 11: Map of Iceland. Distribution of active volcanic zones (fissure swarms – grey, central volcanoes – black, summit crater/caldera – white). Reykjanes Volcanic Zone: (1) Reykjanes–

Svartsengi, (2) Krýsuvík, (3) Brennisteinsfjöll; West Volcanic Zone: (4) Hengill, (5)

Hrómundartindur, (6) Grímsnes, (7) Geysir, (8) Prestahnjúkur, (9) Langjökull; Mid-Iceland Belt:

(10) Hofsjökull, (11) Tungnafellsjökull; East Volcanic Zone: (12) Vestmannaeyjar, (13)

Eyjafjallajökull, (14) Katla, (15) Tindfjöll, (16) Hekla–Vatnafjöll, (17) Torfajökull, (18)

Bárðarbunga–Veiðivötn, (19) Grímsvötn; North Volcanic Zone: (20) Kverkfjöll, (21) Askja, (22)

Fremrinámur, (23) Krafla, (24) Þeistareykir; Öræfajökull Volcanic Belt: (25) Öræfajökull, (26)

Esjufjöll, (27) Snæfell; Snæfellsnes Volcanic Belt: (28) Ljósufjöll, (29) Helgrindur, (30)

Snæfellsjökull. Modified after Thordarson and Larsen (2007) with GeoMap

(http://www.geomapapp.org) (Ryan et al., 2009). 72

Figure 12: Map of the Canary Islands. Created with GeoMap

(http://www.geomapapp.org) (Ryan et al., 2009).

Figure 13: Histograms of Results for the Mid-Atlantic Ridge, showing the distribution of fO2 (relative to the Fayalite-Magnetite-Quartz buffer) calculated using the Olivine-

Melt Method, Fe-Ti oxide oxybarometers, and Fe3+/Fe2+ ratios. 74

Figure 14: Histograms of Results for the East Pacific Rise. Also shown is the range reported in Sano et al. (2011) using plagioclase-liquid equilibrium.

Figure 15: Histograms of Results for the Juan de Fuca Ridge.

Figure 16: Histograms of Results for the Galápagos Spreading Center.

Figure 17: Histogram of Results for the Galápagos Islands.

Figure 18: Histograms of Results for Iceland.

Figure 19: Histograms of Results for the Canary Islands. Also shown is the range

3+ 2+ reported by Klügel et al. (2000) using Fe /Fe ratios to determine fO2.

Figure 20: Distribution of fO2 Calculated using the Olivine-Melt Equilibrium

Method, for MORB (blue) and OIB (red) samples. The minimum and maximum values are represented by the extent of the arrows, and the average values are shown as black circles.

Figure 21: Illustration of Crustal Recycling as a Source for Ocean Island Basalts.

Simplified illustration showing the formation and subduction of oceanic lithosphere. Old, subducted slabs may later be recycled and incorporated into mantle plumes, with a contrast in redox state compared to the surrounding ambient mantle.

Figure 22: Comparison of KD and ΔFMQ across and Olivine Zoning Profile.

Illustration of using an olivine zoning profile from Hansteen (1991) paired with the average of the reported composition of the high MgO pillow glass. The dashed line and shaded area shows the acceptable range of KD (0.30 ± 0.03), as well as the appropriate range of fO2 determined using other methods for this geological setting - Iceland.

Figure 23: Comparison of Average Olivine-Hosted Melt Inclusions and Glass.

Comparison of fO2 calculated using Olivine-Melt Equilibrium. The average ΔFMQ for both melt inclusions (MI) and glass are plotted against each other for comparison for each geological setting.

Figure 24: Comparison of Corrected and Uncorrected Melt Inclusions. Line graph showing calculated ΔFMQ (circles) and raw KD for 102 samples of melt inclusions reported in Wanless and Shaw (2012). Both corrected melt inclusions (black, solid symbols) and uncorrected melt inclusions (red, hollow symbols) are plotted on the same graph. The range of fO2 reported in literature and calculated from other models is shown on the top graph, while the acceptable equilibrium KD values are represented on the bottom graph, as a range of dashed lines.

Appendix B: Tables

The following tables consist of the average ΔFMQ ± 1σ using the Olivine-Melt

Equilibrium Method, Fe3+/Fe2+ ratios, and Fe-Ti oxide geothermometer-oxybarometers for each of the four MORB and three OIB localities chosen for this research. Each table shows the reference in which the data were retrieved, the number of samples from each location, as well as a more specific location for each geological setting. For the Olivine-

Melt Equilibrium Method, the type of glass is indicated. For the Fe3+/Fe2+ ratio model, the analytical method of determining the melt Fe3+/Fe2+ ratio is shown. For the Fe-Ti oxide methods, the specific geothermometer-oxybarometer is referenced.

Table 1: Olivine-Melt Equilibrium Data for the Mid-Atlantic Ridge.

MAR: Olivine-Melt Equilibrium Method (Average ΔFMQ = −1.55 ± 0.75) Location Reference Glass Type # Samples

FAMOUS Hekinian et al., 1976 Pillow Glass 4 Nabelek and Langmuir, 1986 (Ol) FAMOUS and Bryan, 1977 (Gl) Glass 1 Nabelek and Langmuir, 1986 (Ol) FAMOUS and Langmuir, 1977 (Gl) Whole Rock 3 Nabelek and Langmuir, 1986 (Ol) FAMOUS and Langmuir, 1977 (Gl) Glass 1 Nabelek and Langmuir, 1986 (Ol) FAMOUS and Langmuir, 1977 (Gl) Glass 1

FAMOUS Laubier, et al., 2011 Melt Inclusion 315

FAMOUS Shimizu, 1998 Melt Inclusion 22 Lucky Strike Wanless et al., 2015 Melt Inclusion 62 Total 409

Table 2: Fe3+/Fe2+ Ratio Data for the Mid-Atlantic Ridge.

MAR: Fe3+/Fe2+ Ratios (Average ΔFMQ = −0.76 ± 0.63) Location Reference Fe3+ Analytical Method # Samples

FAMOUS Bézos and Humler, 2005 Wet Chemical Analyses 5 Equatorial MAR Bézos and Humler, 2005 Wet Chemical Analyses 10 Reykjanes Ridge Bézos and Humler, 2005 Wet Chemical Analyses 4 Southern MAR Bézos and Humler, 2005 Wet Chemical Analyses 1 Fe K-edge μ-XANES MAR Cottrell and Kelley, 2011 Spectroscopy 22

FAMOUS Miyashiro et al., 1969 Wet Chemical Analyses 3

FAMOUS Schilling et al., 1983 Wet Chemical Analyses 24 Total 69

Table 3: Fe-Ti Oxide Data for the Mid-Atlantic Ridge.

MAR: Fe-Ti Oxides (ΔFMQ = −0.02) Location Reference Geothermometer-Oxybarometer # Samples Equatorial MAR Prinz et al., 1976 Buddington and Lindsley, 1964 1 Total 1

Table 4: Olivine-Melt Equilibrium Data for the East Pacific Rise.

EPR: Olivine-Melt Equilibrium Method (Average ΔFMQ = −0.65 ± 0.51) Location Reference Glass Type # Samples

Spreading Center B Hays, 2004 Pillow Glass 1

Spreading Center A Hays, 2004 Pillow Glass 1

A-B Fault Hays, 2004 Pillow Glass 7

9°30′N Pan and Batiza, 2003 Pillow Glass 4

10°30′N Pan and Batiza, 2003 Pillow Glass 4

11°20′N Pan and Batiza, 2003 Pillow Glass 7 9°50.320'N, 104°17.568'W (2005-2006 Eruption) Wanless and Shaw, 2012 Melt Inclusion 16 9°53.095'N, 104°17.502'W (2005-2006 Eruption) Wanless and Shaw, 2012 Melt Inclusion 9 9°52.965'N, 104°17.525'W (2005-2006 Eruption) Wanless and Shaw, 2012 Melt Inclusion 17 9°50.328'N, 104°17.568'W (2005-2006 Eruption) Wanless and Shaw, 2012 Melt Inclusion 15 9°50'55''N, 104°17'36''W (1991 Eruption) Wanless and Shaw, 2012 Melt Inclusion 18 12°48.687′N, 130°56.439'W Wanless and Shaw, 2012 Melt Inclusion 27 Total 126

Table 5: Fe3+/Fe2+ Ratio Data for the East Pacific Rise.

EPR: Fe3+/Fe2+ Ratios (Average ΔFMQ = −0.30 ± 0.60) Location Reference Fe3+ Analytical Method # Samples Bézos and Humler, EPR 2005 Wet Chemical Analyses 11 Cottrell and Kelley, Fe K-edge μ-XANES EPR 2011 Spectroscopy 17 East Pacific Rise Cottrell and Kelley, Fe K-edge μ-XANES (Siqueiros F.Z.) 2011 Spectroscopy 4 Gale et al., 2013 EPR Database Various Techniques 49 Puchelt and EPR Emmermann, 1983 Wet Chemical Analyses 6 Total 87

Table 6: Fe-Ti Oxide Data for the East Pacific Rise.

EPR: Fe-Ti Oxides (Average ΔFMQ = −0.35 ± 0.21) Location Reference Geothermometer-Oxybarometer # Samples Nazca Plate (off axis) Mazullo et al., 1976 Buddington and Lindsley, 1964 7 Total 7

Table 7: Olivine-Melt Equilibrium Data for the Juan de Fuca Ridge.

JdFR: Olivine-Melt Equilibrium Method (Average ΔFMQ = −0.77 ± 0.42) Locality Reference Sample # Samples

Southern JdFR Dixon et al., 1986 Interstitial Glass 4 Northern JdFR (W Valley N) Wagoner et al., 1991 Pillow Glass 1 Northern JdFR (W Valley S) Wagoner et al., 1991 Pillow Glass 1 Northern JdFR (W Valley Wall) Wagoner et al., 1991 Pillow Glass 1 Northern JdFR (W Ridge) Wagoner et al., 1991 Pillow Glass 1 Northern JdFR (Middle Ridge) Wagoner et al., 1991 Pillow Glass 1 Northern JdFR (Sovanco FZ) Wagoner et al., 1991 Pillow Glass 1 Endeavour Ridge Wagoner et al., 1991 Pillow Glass 3 N Symmetrical Ridge Wagoner et al., 1991 Pillow Glass 1

Vance Segment Wanless and Shaw, 2012 Melt Inclusion 37

Cleft Segment Wanless and Shaw, 2012 Melt Inclusion 8 Total 59

Table 8: Fe3+/Fe2+ Ratio Data for the Juan de Fuca Ridge.

JdFR: Fe3+/Fe2+ Ratios (Average ΔFMQ = −0.20 ± 0.58) Location Reference Fe3+ Analytical Method # Samples

JdFR Cottrell and Kelley, 2011 Fe K-edge μ-XANES Spectroscopy 15

JdFR Gale et al., 2013 Database Various Techniques 10 Total 25

Table 9: Fe-Ti Oxide Data for the Juan de Fuca Ridge.

JdFR: Fe-Ti Oxides (Average ΔFMQ = −0.11 ± 0.35) Location Reference Geothermometer-Oxybarometer # Samples Southern JdFR Dixon et al., 1996 Lindsley and Spencer, 1982 7 Total 7

Table 10: Olivine-Melt Equilibrium Data for the Galápagos Spreading Center.

GSC: Olivine-Melt Equilibrium Method (Average ΔFMQ = 0.08 ± 0.48) Location Reference Glass Type # Samples

GSC Coleman et al., 2015 Melt Inclusion 79

GSC Rilling, 2005 Glass 17 Total 96

Table 11: Fe3+/Fe2+ Data for the Galápagos Spreading Center.

GSC: Fe3+/Fe2+ Ratios (Average ΔFMQ = −0.15 ± 0.50) Location Reference Fe3+ Analytical Method # Samples

GSC Byers et al., 1983 Wet Chemical Analyses 8

GSC Cottrell and Kelley, 2011 Fe K-edge μ-XANES Spectroscopy 8

GSC Fornari et al., 1983 Wet Chemical Analyses 33

GSC Gale et al., 2013 Database Various Techniques 38 Puchelt and Emmermann, GSC 1983 Wet Chemical Analyses 11

GSC Rilling, 2005 Wet Chemical Analyses 7 Total 105

Table 12: Fe-Ti Oxide Data for the Galápagos Spreading Center.

GSC: Fe-Ti Oxides (Average ΔFMQ = 0.22 ± 0.45) Location Reference Geothermometer-Oxybarometer # Samples Perfit and Fornari, Spencer and Lindsley (Modified from GSC 1983 Buddington and Lindsley, 1964) 7 Total 7

Table 13: Olivine-Melt Equilibrium Data for the Galápagos Islands.

Galápagos Islands: Olivine-Melt Equilibrium (Average ΔFMQ = −0.33 ± 0.35) Location Reference Glass Type # Samples Melt Fernandina Kolezar et al., 2009 Inclusion 76 Melt Santiago Kolezar et al., 2009 Inclusion 44

Fernandina Rilling, 2005 WR 17

Volcan Darwin Rilling, 2005 Glass 2

Roca Redonda Rilling, 2005 Glass 6 Total 145

Table 14: Olivine-Melt Equilibrium Data for Iceland.

Iceland: Olivine-Melt Equilibrium Method (Average ΔFMQ = −0.43 ± 0.71) Location Reference Glass Type # Samples

Kistufell Breddam, 2002 Hyaloclastite Glass 7

Kistufell Breddam, 2002 Pillow Glass 1

Laki Guilbaud et al., 2007 Matrix Glass 1

Heleyjabunga Gurenko and Chaussidon, 1995 Pillow Glass 1 Continued 92

Table 14 Continued

Midfell Gurenko and Chaussidon, 1995 Pillow Glass 2

Mælifell Gurenko and Chaussidon, 1995 Pillow Glass 2

Heleyjabunga Gurenko and Chaussidon, 1995 Melt Inclusion 5

Midfell Gurenko and Chaussidon, 1995 Melt Inclusion 6

Mælifell Gurenko and Chaussidon, 1995 Melt Inclusion 2

Midfell Gurenko and Sobolev, 2006 Pillow Glass 3 Glass Crusts around Midfell Gurenko and Sobolev, 2006 Gabbro Xenoliths 4 Interstitial Glass in Midfell Gurenko and Sobolev, 2006 Gabbro Xenoliths 4

Fontur Hansen and Gronveld, 2000 Glass in Nodule 1

Saxi Crater Hansen and Gronveld, 2000 Matrix Glass 1 Hrímalda (East side) Hansen and Gronveld, 2000 Matrix Glass 1 Kverkfjöll Pl Ultraphyric Pillow Hansen and Gronveld, 2000 Matrix Glass 1

Búrfell Hansen and Gronveld, 2000 Pillow Glass 1

Brandur Crater Hansen and Gronveld, 2000 Melt Inclusion 3

Gigoldur Tuff Hansen and Gronveld, 2000 Melt Inclusion 1 Gigoldur Hyaloclastite Hansen and Gronveld, 2000 Melt Inclusion 1

Búrfell Pillows Hansen and Gronveld, 2000 Melt Inclusion 1

Mælifell Hansteen, 1991 Melt Inclusion 5

Mælifell Hansteen, 1991 Pillow Glass 4

Eyjafjallajökull Keiding and Sigmarsson, 2012 Tephra Glass 6

Laki Passmore et al., 2012 Tephra Glass 3

Hekla Portnyagin et al., 2013 Melt Inclusion 9 Continued

Table 14 Continued

Bláfjall Schiellerup, 1995 Hyaloclastite Glass 2 Total 78

Table 15: Fe3+/Fe2+ Ratio Data for Iceland.

Iceland: Fe3+/Fe2+ Ratios (Average ΔFMQ = −0.17 ± 0.38) Location Reference Fe3+ Analytical Method # Samples Mössbauer analyses of Fe3+/ΣFe of Midfell Oskarsson et al., 1994 Pillow Glass 1 Mössbauer analyses of Fe3+/ΣFe of Hengill Oskarsson et al., 1994 Pillow Glass 1 Mössbauer analyses of Fe3+/ΣFe of Stapafell Oskarsson et al., 1994 Pillow Glass 1 North Rift Mössbauer analyses of Fe3+/ΣFe of Zone Oskarsson et al., 1994 Pillow Glass 7 Total 10

Table 16: Fe-Ti Oxide Data for the Iceland.

Iceland: Fe-Ti Oxides (Average ΔFMQ = −0.46 ± 0.35) Location Reference Geothermometer-Oxybarometer # Samples Markarfljót Silicic Domes Gunnarsson et al., 1998 Oxybarometer of Stormer, 1983 10 Laufafell Silicic Domes Gunnarsson et al., 1998 Oxybarometer of Stormer, 1983 12 Hrafntinnusker Unit Gunnarsson et al., 1998 Oxybarometer of Stormer, 1983 2 Crustal Xenolith Gunnarsson et al., 1998 Oxybarometer of Stormer, 1983 10 Hekla - H3 QUILF program of Andersen & eruption Portnyagin et al., 2013 Lindsley, 1988 32 Hekla - H4 QUILF program of Andersen & eruption Portnyagin et al., 2013 Lindsley, 1988 10 Frost and Lindsley, QUILF program of Frost and Thingmuli 1992 Lindsley, 1992 14 Total 90

Table 17: Olivine-Melt Equilibrium Data for the Canary Islands.

Canary Islands: Olivine-Melt Equilibrium Method (Average ΔFMQ = 0.68 ± 0.52) Location Reference Glass Type # Samples Gran Canaria (submarine) Gurenko et al., 1998 Melt Inclusion 28 La Palma (Taburiente Volcano) Nikogosian et al. 2002 Melt Inclusion 28 Hierro (submarine NE-Rift) Stroncik et al., 2009 Pillow Glass 6 Hierro (submarine NW-Rift) Stroncik et al., 2009 Pillow Glass 7 Total 69

Table 18: Fe3+/Fe2+ Ratio Data for the Canary Islands

Canary Islands: Fe3+/Fe2+ Ratios (Average ΔFMQ = 0.47 ± 0.56) Location Reference Fe3+ Analytical Method # Samples Calculated from Method of Gran Canaria Gurenko et al., 1996 Maurel and Maurel, 1982 9 Gurenko and Calculated from Method of Gran Canaria Schmincke, 2000 Maurel and Maurel, 1982 49 Total 58

Table 19: Fe-Ti Oxide Data for the Canary Islands.

Canary Islands: Fe-Ti Oxides (Average ΔFMQ = 0.83 ± 0.80) Geothermometer- Location Reference Oxybarometer # Samples

Tenerife Ablay et al., 1998 Sack and Ghiorso, 1991 17

Tenerife Andujar et al., 2013 Sauerzapf et al., 2008 1

Tenerife Bryan, 2006 Ghiorso and Sack, 1991 10 Tenerife (Las Canadas Edifice) Bryan et al., 2002 Ghiorso and Sack, 1991 16

Gran Canaria Freundt et al., 1995 Spencer and Lindsley, 1981 32 Buddington and Lindsley, Tenerife Wolf and Storey, 1983 1964 12 Total 88 95

Table 20: S6+/ΣS Ratio Data for the Canary Islands.

Canary Islands: S6+/ΣS Ratios (Average ΔFMQ = 0.86 ± 0.68) Location Reference Model # Samples Gurenko and Schmincke, Wallace and Gran Canaria 2000 Carmichael, 1994 49 Total 49

Appendix C: Supplementary Tables

All supplementary tables consist of results for each of the seven different locations represented by this research. The results include those obtained using the

Olivine-Melt Equilibrium Method, the Fe3+/Fe2+ ratio method (Equation 5), as well as Fe-

Ti Oxide Geothermometer-Oxybarometers. These tables include the reference in which the data were retrieved, the sample location, identification number, and glass and olivine compositional data for every individual sample. All data are available upon request.

List of Supplementary Tables:

Table S1: Results of the Olivine-Melt Equilibrium Method for the Mid-Atlantic Ridge

Table S2: Results of the Fe3+/Fe2+ Ratio Model for the Mid-Atlantic Ridge

Table S3: Results of Fe-Ti Oxide Oxybarometers for the Mid-Atlantic Ridge

Table S4: Results of the Olivine-Melt Equilibrium Method for the East Pacific Rise

Table S5: Results of the Fe3+/Fe2+ Ratio Model for the East Pacific Rise

Table S6: Results of Fe-Ti Oxide Oxybarometers for the East Pacific Rise

Table S7: Results of the Olivine-Melt Equilibrium Method for the Juan de Fuca Ridge

Table S8: Results of the Fe3+/Fe2+ Ratio Model for the Juan de Fuca Ridge

Table S9: Results of Fe-Ti Oxide Oxybarometers for the Juan de Fuca Ridge

Table S10: Results of the Olivine-Melt Equilibrium Method for the Galápagos Spreading

Center

Table S11: Results of the Fe3+/Fe2+ Ratio Model for the Galápagos Spreading Center

Table S12: Results of Fe-Ti Oxide Oxybarometers for the Galápagos Spreading Center

Table S13: Results of the Olivine-Melt Equilibrium Method for the Galápagos Islands

Table S14: Results of the Olivine-Melt Equilibrium Method for Iceland

Table S15: Results of the Fe3+/Fe2+ Ratio Model for Iceland

Table S16: Results of Fe-Ti Oxide Oxybarometers for Iceland

Table S17: Results of the Olivine-Melt Equilibrium Method for the Canary Islands

Table S18: Results of the Fe3+/Fe2+ Ratio Model for the Canary Islands

Table S19: Results of Fe-Ti Oxide Oxybarometers for the Canary Islands

Table S20: Results of the S6+/ΣS Ratio Model for the Canary Islands