GEOCHEMISTRY OF THE GALÁPAGOS ARCHIPELAGO: UNRAVELING DEEP MANTLE SIGNATURES FROM THE INFLUENCE OF SHALLOW LEVEL PROCESSES BY MARY E. PETERSON B.S. GEOLOGICAL SCIENCES, UNIVERSITY OF MICHIGAN, 2009 SC.M., GEOLOGICAL SCIENCES, BROWN UNIVERSITY, 2011 A DISSERTATION SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FRO THE DEGREE OF DOCTOR OF PHILOSOPHY IN THE DEPARTMENT OF EARTH, ENVIRONMENTAL AND PLANETARY SCIENCES AT BROWN UNIVERSITY PROVIDENCE, RHODE ISLAND MAY 2015 © Copyright 2015 by Mary E. Peterson This dissertation by Mary E. Peterson is accepted in its present form by the Department of Earth, Environmental, and Planetary Sciences as satisfying the dissertation requirement for the degree of Doctor of Philosophy. Date____________ ___________________________________ Alberto E. Saal, Ph.D, Advisor Recommended to the Graduate Council Date____________ ___________________________________ Reid F. Cooper, Ph.D, Reader Date____________ ___________________________________ Donald W. Forsyth, Ph.D, Reader Date____________ ___________________________________ Mark D. Kurz, Ph.D, Reader Date____________ ___________________________________ Stephen W. Parman, Ph.D, Reader Approved by the Graduate Council Date____________ ___________________________________ Peter M. Weber Dean of the Graduate School ! iii! CURRICULUM VITAE: MARY E. PETERSON Department of Earth, Environmental and Planetary Sciences, Brown University, Providence, RI, 02912 PROFESSIONAL PREPARATION Brown University – Geochemistry (Expected) Ph.D. May 2015 Brown University – Geochemistry MS May 2011 University of Michigan - Geology (honors) B.S. April 2009 Honors Thesis: Elevated water concentrations in olivine-hosted melt inclusions: Insight into the source of high melt volumes in the Iceland hotspot APPOINTMENTS Research Assistant, Brown University Sept 2014 – Present NSF Graduate Research Fellow Sept 2011 – Aug 2014 Teaching Assistant, Brown University Jan 2011 – May 2011 Charles Wilson Brown Fellowship, Brown University Sept 2009 – May 2010 MANUSCRIPTS • Peterson, M., Saal, A., Nakamura, E., Kitagawa, H., Kurz, M. (2014). Origin of the ‘ghost plagioclase’ signature in Galapagos melt inclusions: new evidence from lead isotopes. Journal of Petrology. 55(11), 2193-2216. doi: 10.1093/petrology/egu054 • Peterson, M., Saal, A., Kurz, M., Hauri, E., Blusztaijn, J., Harpp, K., Werner, R., Geist, D., (submitted to Journal of Petrology). Determining the volatile budget of the Galapagos Plume: separating deep and shallow signatures. • Peterson, M., Kelley, K., Cottrell, L., Saal, A., Kurz, M. (in prep). The oxidation state of Fe in glasses from the Galapagos Archipelago: variation in oxygen fugacity as a function of mantle source. • Peterson, M., Wang, Z., Saal, A., Kurz, M., Eiler, J. (in prep). Oxygen isotopes in lavas from the Galapagos Archipelago: evaluating contributions from ancient and present-day crustal sources. ABSTRACTS AND PRESENTATIONS • Peterson, M., Saal, A., Hauri, E., Werner, R., Hauff, F., Kurz, M., Geist, D., Harpp, K. (2013) Volatile budget of the Galapagos plume. Goldschmidt2013 Conference Abstracts Mineralogical Magazine, 77(5) 1956 (poster) • Invited talk: "Volatiles in submarine glasses from the Galapagos: Process vs. source” March 28, 2012 BU, Solid Earth Seminar • Peterson, M., Saal, A., Nakamura, E., Kitagawa, H., Kurz, M., (2011). Pb isotopes and the source of the “ghost plagioclase” signature in melt inclusions from the Galapagos archipelago. Goldschmidt2011 Conference Abstracts Mineralogical Magazine, 75(3) 1628 (poster) • Peterson, M., Saal, A., Hauri, E., Werner, R., Hauff, F., Kurz, M., Geist, D., Harpp, K. (2011) Volatile budget of the sources of the Galapagos plume. American Geophysical Union Chapman Conference: The Galapagos as a Laboratory for the Earth Sciences, (poster) • Peterson, M., Saal, A., Hauri, E., Werner, R., Hauff, F., Kurz, M., Geist, D., Harpp, K. (2010) Sources of volatiles in basalts from the Galapagos Archipelago: deep and shallow evidence. American Geophysical Union, Fall Meeting 2010 (oral) iv • Peterson, M., Mukasa, B., Stefano, C., Shimizu, N., Kent, A., (2009) Elevated water concentrations and evidence of small scale heterogeneity in olivine-hosted melt inclusions: Insight into the source of high melt volumes in the Iceland hotspot. American Geophysical Union, Fall Meeting 2009, vol. 1, p. 05. (oral) OTHER PRODUCTS • Mukasa, S., Loudin, L., Peterson, M., Dixon, E., (2013) Relationship between volatiles and noble gases in Icelandic Lavas: Evidence for crustal recycling. Goldschmidt2013 Conference Abstracts Mineralogical Magazine, 77(5) 1801 • Yao, L., Dygert, N., Peterson, M., Sun, C., Wetzel, D., Liang, Y., (2010) A “bundle of columns” model for trace element fractionation during melting and melt migration in a vertically upwelling, chemically and lithologically heterogeneous mantle. American Geophysical Union, Fall Meeting 2010, abstract# V11A-2258. AWARDS • 2011-2014 NSF Graduate Research Fellowship • 2009 Charles Wilson Brown University Fellowship • 2008-2009 Department of Geological Sciences Award for Academic Excellence, University of Michigan HONORS • 2011-2012 Graduate Student Council representative • 2010-2011 Geoclub President • 2007-2009 Member of Sigma Gamma Epsilon, University of Michigam TEACHING • Teaching Assistant, GEOL023–Geochemistry: Earth and Planetary Materials and Processes o Designed and ran weekly labs (focus on hand specimen identification), ran review sessions, assisted professors on field trips • Teaching Assistant, GEOL0010-Face of the Earth (Introductory Geology) o Designed and ran labs (focus on hand specimen identification and use of topography maps), ran review sessions and assisted professors on field trips • Sheridan Teaching Certificate I FIELD WORK • 2012 Department Field Trip: Expression of Ancient Magmatism, Vinalhaven, Maine, USA • 2010 Graduate Field Trip: Connecting Komatiites and Impact sites to Industry, Quebec and Ontario, Canada (Trip Leader) • 2009 Graduate Field Trip: Exploring the Bay of Islands Ophiolite, Newfoundland, Canada • 2009 Soft Rocks Field Trip: Sea level rise and fall and the expansion of the North American Craton, New Mexico and Texas, USA • 2008 Soft Rocks Field Trip: Comparing ancient and modern Coral Reefs, Florida, USA • 2007 Geology Field Course: University of Michigan Camp Davis 7 week field course, Jackson Hole, WY, USA v SYNERGISTIC ACTIVITIES • 2013-2014 Graduate Mentor • 2009-2014 Volunteer Science Teacher/developer of curriculum, Vartan Gregorian Elementary School, Providence, RI, Second grade PROFESSIONAL SOCIETY MEMBERSHIPS • American Geophysical Union • Mineralogical Society of America • Geochemical Sciences vi ACKNOWLEDGMENTS It’s difficult to know where to begin in acknowledging the people without which this dissertation may never have come into being. I’ve been on this path since I was a little kid on the beach in Michigan, determined to find fossils amongst the lake-smoothed rocks. Everyone who encouraged me to pursue science from then till now deserves some partial credit, but listing them here would be my own catalogue of ships in the epic poem that is my graduate life. Therefore, I have made an effort to avoid making my acknowledgements the part you skip. I’d like to start by acknowledging the members of my committee, who have helped to keep me on track as I waded through the complicated geochemistry of the Galapagos. Don Forsyth, I appreciate the patience you showed me in the discussions we’ve had about seismic data for hotspots and the very simple statistical models I’ve tried to incorporate into my research. You’ve inspired me to try harder to understand the why of my results. Steve Parman and Yan Liang, you gave me the tools I needed to connect the things I read to the things I do as well as pushing me to reach for the less obvious answer. Mark Kurz, I cannot express my gratitude at your willingness to discuss my explorations of Galapagos data, whether it is over the phone or through a long chain of emails. You’ve helped to elevate the quality of my research. Reid Cooper, your willingness to simply chat with me about things I read, or hear or learn has been invaluable to me. In addition to the breath of scientific knowledge you passed on to me, you helped me find my confidence both in my understanding of topics that initially felt out of reach and in my ability to pursue new knowledge. I will always try to live by your mantra, that I’m trying to become a doctor of Philosophy. Alberto Saal, as my advisor ! vii! you’ve challenged my understanding, battled me in my interpretations, pushed me to be more careful, thoughtful and rigorous, and in the end made me the scientist I am. Your willingness to listen to me and to go to bat for me as well as your obvious desire for my success helped me to find the drive to finish, for which I will always be grateful. Though I cannot guarantee it will always happen, I will continue to try not to shoot from the hip. Beyond my committee, there are many other people who have contributed to my development as a scientist. Marc Parmentier, Greg Hirth, Mac Rutherford and Paul Hess I greatly appreciate your willingness to be sounding board for my ideas and for the good advice you’ve given me when I ran head first into scientific brick walls. Katie Kelley and Liz Cottrel, thank you for your enthusiasm about redox studies in the Galapagos and your willingness to field my thousands of questions. Matt Jackson, for encouraging me to expand my sites beyond my comfort zone and make the big connections. I’d also like to thank the faculty at the University of Michigan who gave me the foundation with which to build my knowledge of geochemistry. In particular I’d like to acknowledge my undergraduate advisor, Sam Mukasa, who gave me my first taste of geologic research and who single-handedly turned me from my path towards paleontology to the geochemistry of hotspots. Going back further, I’d like to acknowledge Larry Kaye, Noelle Ochoa, and Kim Woodcroft who were the first adults to really encourage me to be curious about the world and to push me to go beyond memorizing the answers to gaining an understanding of the questions. I wouldn’t have been able to stay sane if it weren’t for the support of the DEEPS graduate student body. I have to acknowledge Chen Sun and Chris Havlin for the hours they’ve spent walking me through analytical problems, going through my code, and ! viii! explaining their ridiculously complicated research to me. Shannon Loomis, Amanda Getsinger and Lijing Yao for giving me someone I can talk through my ideas with. The current and past basement GMPers who have toiled through classes, reading groups, and lunch bunches with me, always giving the days a shine of amusement and the potential for a party. I have to thank Kei Shimizu for riding out the roller coaster of my emotions. I’ve seen more of you these past 5 years than any other person on Earth and yet you’re still willing to talk me back from the ledge and listen to my stupid ideas. I like to give you a hard time, but know that I couldn’t have asked for a better office mate. Tim Goudge and Tabb Prissel, the support and encouragement you’ve given me this past year has been invaluable. You’ve helped give me the push I needed to finish. Go Defenders! There have been a couple of groups of people who have been pivotal in letting me keep my life in addition to my work. The members of my pathfinder group, Nick, Kat, Sandra, Charlie, Kevin, Tim, Lauren, Jeff, it’s hard to express the impact you’ve had on me. I know in the end it’s just a game that we play together, but for me, having an outlet for creativity and silliness again was something that I didn’t realize I had lost and raised my base level of happiness when I got it back. I’d like to acknowledge my various soccer teams for giving me a place to let out my frustrations and keeping me from becoming a total slug. I’d like to thank Emily Hopper, Diane Wetzel, Stephanie Bouchey, Lauren Joswiak, and my two roommates, Rebecca Greenberger and Jenny Whitten. You went beyond what a person could ask for from friends. You became my family and saw me through the incredible highs and crushing lows that I’ve gone through in graduate school. ! ix! Finally I’d like to acknowledge my family. The unconditional love, support, advice, and phone time you’ve given me are the reason I was able to stick it out. I’m incredibly lucky to have you all in my life. -There are no mistakes, simply happy accidents -Bob Ross ! x! TABLE OF CONTENTS Title Page ............................................................................................................................ i! Copyright Page .................................................................................................................. ii! Signature Page .................................................................................................................. iii! Curriculum Vitae ............................................................................................................. iv! Acknowledgements ......................................................................................................... vii! Table of Contents ............................................................................................................. xi! INTRODUCTION ............................................................................................................ 1! References!..........................................................................................................................................!10! CHAPTER 1: Origin of the “ghost plagioclase” signature in Galapagos melt inclusions: new evidence from lead isotopes................................................................. 21! Abstract!..............................................................................................................................................!22! Introduction!.......................................................................................................................................!23! Previous Results!................................................................................................................................!25! Samples!..............................................................................................................................................................!25! Major and trace element compositions!....................................................................................................!26! The role of plagioclase: evidence from major and trace elements!.................................................!27! Santiago Inclusions!....................................................................................................................................................!27! Fernandina Inclusions!...............................................................................................................................................!28! Lead Isotope Analytical Methods!.................................................................................................!29! Results!.................................................................................................................................................!30! Discussion!...........................................................................................................................................!32! Evidence for contamination by plagioclase cumulates: relationship between major, trace elements, and Pb isotopes!............................................................................................................................!32! Evaluating the two theories for the origin of the Fernandina “ghost plagioclase” signature!33! Estimating the Pb isotope composition of ancient plagioclase cumulates!..............................................!33! Melting of an ancient plagioclase cumulate: adding Pb isotopes to a critical melting model!..........!36! Present-day formation of a ghost plagioclase signature!................................................................................!38! Presence of the ghost plagioclase signature in settings outside the Galapagos!.....................................!43! Conclusions!........................................................................................................................................!44! Funding!...............................................................................................................................................!44! Acknowledgements!..........................................................................................................................!45! References!..........................................................................................................................................!46! Figure Captions!................................................................................................................................!57! Figures (11)!......................................................................................................................................................!64! Tables (5)!............................................................................................................................................!76! CHAPTER 2: Determining the volatile budget of the Galapagos Plume: separating deep and shallow signatures........................................................................................... 86! Abstract!..............................................................................................................................................!87! Introduction!.......................................................................................................................................!89! Geochemical Background!...............................................................................................................!91! Analytical Methods!..........................................................................................................................!94! Results!.................................................................................................................................................!97! Array 1 between Fernandina and Sierra Negra/Cerro Azul: HHe Group and ITE Group!..!100! xi Array 2 between Genovesa (MORB-like) and Pinta/Wolf, Darwin Islands: ITD Group and WD Group!......................................................................................................................................................!102! Volatiles!..........................................................................................................................................................!104! Discussion!.........................................................................................................................................!105! Shallow level process!.................................................................................................................................!105! Degassing!....................................................................................................................................................................!105! Sulfide Saturation!.....................................................................................................................................................!108! Assimilation!...............................................................................................................................................................!109! Mantle Composition!...................................................................................................................................!113! The effect of H2O content on mantle rheology!...............................................................................................!117! C/3He ratios of the high 3He/4He mantle source!.............................................................................................!118! Conclusions!......................................................................................................................................!120! Acknowledgments!..........................................................................................................................!123! References!........................................................................................................................................!124! Figure Captions!..............................................................................................................................!145! Figures (13)!...................................................................................................................................................!154! Supplementary Figures (3)!.......................................................................................................................!167! Tables!(4)!........................................................................................................................................!170! CHAPTER 3: The oxidation state of Fe in glasses from the Galapagos Archipelago: variation in oxygen fugacity as a function of mantle source ..................................... 200! Abstract!............................................................................................................................................!201! Introduction!.....................................................................................................................................!203! Geologic background and previous work!.................................................................................!205! Sample preparation and methods!...............................................................................................!210! Results!..............................................................................................................................................!211! Discussion!.........................................................................................................................................!213! Magmatic oxidation during differentiation and partial melting!...................................................!213! The role of volatiles in magmatic oxidation!.......................................................................................!217! Magmatic oxidation during assimilation of hydrothermally altered material!.........................!220! Mantle source variation in Fe3+/ΣFe ratios!..........................................................................................!221! Conclusions!......................................................................................................................................!223! Acknowledgments!..........................................................................................................................!224! References!........................................................................................................................................!226! Figure Captions!..............................................................................................................................!242! Figures (7)!........................................................................................................................................!246! Tables (2)!..........................................................................................................................................!253! CHAPTER 4: Oxygen isotopes in lavas from the Galapagos Archipelago: evaluating contributions from ancient and present-day crustal sources .................................... 255! Abstract!............................................................................................................................................!256! Introduction!.....................................................................................................................................!257! Methods and Previous Results!....................................................................................................!259! Results!...............................................................................................................................................!261! Discussion!.........................................................................................................................................!262! Fractional crystallization and partial melting!.....................................................................................!262! δ18O!variation!of!the!main!isotopic!end>members!in!the!Galapagos!...................................!264! Assimilation!..................................................................................................................................................!265! δ O%values%greater%than%normal%upper%mantle!............................................................................................!265! 18 δ O%values%lower%than%normal%upper%mantle!................................................................................................!266! 18 xii Alternative Explanation for light δ18O in the Galapagos!...............................................................!269! Conclusions!......................................................................................................................................!270! Acknowledgements!........................................................................................................................!272! References!........................................................................................................................................!273! Figure Captions!..............................................................................................................................!282! Figures (6)!........................................................................................................................................!285! Tables (3)!..........................................................................................................................................!291! Synthesis and Future Directions .................................................................................. 294! References!........................................................................................................................................!304! xiii INTRODUCTION Plate tectonics have fueled almost continuous communication between the interior and the exterior geochemical reservoirs for most of the Earth’s history. Oceanic melts erupt at the surface (recording the fractionating effects of melting and differentiation), are imprinted by the chemistry of the lithosphere, atmosphere and ocean (systems that have evolved through time), and are subducted back into the mantle. During subduction the material goes through an unconstrained amount of dehydration and metamorphism and is then mixed to some degree back into the chemistry of the mantle, beginning the cycle anew. Mantle heterogeneity, as a result, indisputably exists in the Earth, but the causes, effects and scale of this heterogeneity are poorly resolved. Therefore, one of the motivations of geochemical study is to determine how material collected at the surface reflects the composition of the mantle. Ocean islands, or hotspots such as the Galapagos Archipelago, are ideal locations to study the mantle because these melting anomalies are thought to be the product of plumes of buoyantly upwelling material originating within the Earth from as deep as the core-mantle boundary (CMB) (Morgan, 1971). Ocean islands are generally associated with a large buoyancy flux, low velocity anomalies and thin transition zones (identified with seismic data) (e.g. Courtillot et al., 2003, Hooft et al., 2003, Villagomez et al., 2007, Villagomez et al., 2011, Rychert et al., 2014, Villagomez et al., 2014). These characteristics can be explained by the presence of greater than ambient mantle thermal anomalies (thus the term “hotspot”) originating at the CMB. Oceans islands are also unique in terms of their noble gas isotopic signatures with high 3He/4He ratios, solar-like Ne isotopes and unique heavier noble gas isotopes that are difficult to explain without the ! 1! presence of primitive mantle material (e.g. Kurz & Jenkins, 1981, Kurz et al., 1982, Graham et al., 1993, Kurz & Geist, 1999, Graham et al., 2001, Georgen et al., 2003, Kurz et al., 2009, Colin et al., 2011, Mukhopadhyay, 2012). In comparison to other geological settings, ocean island basalts in general exhibit extreme heterogeneity in their geochemical signatures. This is attributed to the presence of several isotopically unique reservoirs, hypothesized to be the result of both primitive mantle and recycled crustal material (Hofmann & White, 1982, Allègre et al., 1983, Zindler & Hart, 1986, White et al., 1993, Hofmann, 1997, White, 2010). Though it is hard to reconcile global mantle geochemical anomalies with a convecting mantle, material housed at the CMB is potentially isolated from the effects of convection, providing further support for ocean island basalts being the result of deep mantle plumes. Unequivocally identifying specific mantle signatures in OIBs is difficult, however, due to an inadequate understanding of the extent to which upwelling melt is affected by interaction with the shallow crust and lithosphere. The resulting ambiguity in geochemical signatures calls the nature of ocean island basalts and their hypothesized origin at the CMB into question. This has led to a rift in the community between those that support and oppose using the plume theory to explain the characteristics of ocean islands (e.g. Anderson, 2000, Foulger, 2011). In this dissertation, I have used a large geochemical data set of submarine glass, melt inclusions and whole rock, collected from across the Galapagos Archipelago, and a variety of techniques to explore the influence shallow level processes have on lavas erupted at the surface to then isolate the deep mantle signatures unique to the plume. ! 2! The Galapagos Archipelago, located < 200 km south of the Galapagos Spreading Center (GSC), contains 16 of 21 emergent volcanoes that have erupted since the Holocene, making it one of the most volcanically active areas in the world. Seismic studies have identified both a low-velocity anomaly and a thin transition zone in the mantle below the Galapagos Archipelago, characteristics attributed to a hot mantle plume column extending to at least as deep as the transition zone (Hooft et al., 2003, Villagomez et al., 2007, Villagomez et al., 2011, Villagomez et al., 2014). The plume column tilts towards the GSC, against the direction of plate motion (Villagomez et al., 2014). This in conjunction with geochemical studies conducted along the ridge show the strong influence the Galapagos plume has on the GSC (Detrick et al., 2002, Schilling et al., 2003, Cushman et al., 2004, Ingle et al., 2010, Gibson et al., 2015). There are four main compositional components attributed to Galapagos plume that are concentrated in different geographical regions (White et al., 1993, Hoernle et al., 2000, Harpp & White, 2001, Werner et al., 2003, Saal et al., 2007). The western component is characterized by the lavas of Fernandina and is distinguishable by high 3 He/4He isotopic ratios (Graham et al., 1993, Kurz & Geist, 1999, Geist et al., 2006, Kurz et al., 2009). This component is generally compared to a relatively volatile-rich primitive mantle component (Kurz et al., 1982, Farley et al., 1992, Hart et al., 1992, Hanan & Graham, 1996). There is a MORB-like component that, though it can be found across the archipelago, is focused in the central and eastern region and anchored by the composition of lavas from Genovesa (Harpp et al., 2002, Harpp et al., 2003, Geist et al., 2005). The northern and southern components (illustrated by the lavas of Pinta and Floreana respectively) display enrichments in trace elements in addition to isotopic signatures ! 3! generally associated with ancient recycled material in the source. Floreana is most 238 commonly compared to a HIMU component (high-μ = high U/204Pb) due to enrichments in Pb isotopes (e.g. Harpp et al., 2014). Pinta, in comparison, has low Nd 207 and Hf isotopes with elevated Pb/204Pb and 208 Pb/204Pb at a given 206 Pb/204Pb ratio, akin to the DUPAL anomaly (Hart, 1984) and generally associated with an Enriched Mantle (EM) component (White et al., 1993, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Schilling et al., 2003). Lavas collected from the other islands and the submarine platform in the Galapagos have intermediate compositions, leading to the assumption that there is widespread mixing of these 4 end-members throughout the archipelago (e.g. Harpp & White, 2001, Gibson et al., 2012). Chapter one [Peterson, M., Saal, A., Nakamura, E., Kitagawa, H., Kurz, M. (2014). Origin of the ‘ghost plagioclase’ signature in Galapagos melt inclusions: new evidence from lead isotopes. Journal of Petrology. 55(11), 2193-2216. doi: 10.1093/petrology/egu054] focuses on identifying the origin of the “ghost plagioclase” signature using new Pb isotope analyses of olivine-hosted melt inclusions from Fernandina and Santiago that both show a trace element signature consistent with the presence of plagioclase cumulates. In the case of Santiago inclusions, there is evidence for simple plagioclase assimilation. Fernandina inclusions, in contrast, possess trace element contents indicative of the assimilation of plagioclase cumulates, without the expected major element signature (i.e. a “ghost plagioclase” trace element signature). Two competing hypotheses have emerged to explain the ghost plagioclase signature: 1. Incorporation of ancient, recycled plagioclase-rich cumulate (now eclogite) into the mantle and 2. Shallow-level interaction with plagioclase cumulates in the present-day ! 4! lower oceanic crust. We explore both hypotheses using Pb isotope measurements of the olivine-hosted inclusions from both islands. We find that melting and mixing models of ancient (~0.5-1 Ga) recycled plagioclase-rich cumulates cannot reproduce the observed trace element and Pb isotope characteristics of the Fernandina inclusions with the ghost plagioclase signature. Shallow-level diffusive interactions between basalt and present-day plagioclase-rich cumulates, as a result, provide the simplest explanation for the observed trace element and isotopic compositions of melt inclusions from Fernandina and Santiago Islands. Chapter 2 [Peterson, M., Saal, A., Kurz, M., Hauri, E., Blusztaijn, J., Harpp, K., Werner, R., Geist, D., (submitted to Journal of Petrology). Determining the volatile budget of the Galapagos Plume: separating deep and shallow signatures] takes a broader approach by exploring the volatile contents of submarine glass collected from across the Galapagos Archipelago. Volatile elements assert a strong influence on processes such as mantle melting, magma crystallization, and style of eruption as well as the viscosity and rheology of the mantle (Hirth & Kohlstedt, 1996, Gaetani & Grove, 1998, Asimow & Langmuir, 2003, Asimow et al., 2004). Surface processes such as degassing, sulfide saturation, and melt-crustal interaction, however, can significantly modify the primitive volatile/refractory element ratios of the melt (Michael, 1989, Dixon et al., 1995, Dixon, 1997, Michael & Cornell, 1998, Dixon & Clague, 2001, Saal et al., 2002, Stroncik & Haase, 2004, le Roux et al., 2006, Workman et al., 2006). Therefore we use new major, trace and isotopic measurements in conjunction with volatile element contents (H2O, CO2, F, Cl, and S) of 118 glass chips to separate the signature of shallow level processes from the variation in volatile content of the Galapagos mantle plume. ! 5! Major, trace and isotopic measurements allow us to separate the samples into 2 broad arrays that range between the compositions of Fernandina and Sierra Negra/Cerro Azul (array 1) and between the composition of Genovesa and Pinta (array 2). Degassing has significantly decreased the CO2 content of the magmas while having a minor affect on H2O and S concentrations and no recognizable impact on Cl and F concentrations. Melts range from being just saturated with a sulfide liquid to being sulfide undersaturated at estimated pressures of crystal fractionation (~2 kbars). Assimilation of hydrothermally altered material, assessed with Cl/K ratios and the magnitude of the Sr anomaly (Sr/Sr* ≥ 1) in a primitive mantle normalized diagram, has significantly affected the samples of both arrays. The chemistry of array 2 further suggests that mixing of melts occurs after some degassing and contamination of the end-member compositions has taken place. Once shallow level processes are accounted for we use ratios of volatile and refractory elements with similar incompatibilities to estimate the volatile budget of the Galapagos mantle plume. Distinctive H2O/Ce and F/Nd ratios can differentiate the four main compositional components defined by the glasses of this study. Using the equations of Hirth and Kohlstedt (1996), the higher H2O contents of the high 3He/4He Fernandina source will result in a mantle viscosity 5-7x lower than what we would predict for the depleted upper mantle component. The high 3He/4He end-member also shows C/3He ratios an order of magnitude higher than the depleted upper mantle and among the highest yet reported for OIB samples; a characteristic we attribute to a more carbonated source (Marty & Jambon, 1987, Trull et al., 1993, Marty & Tolstikhin, 1998, Marty & Zimmermann, 1999, Shaw et al., 2004, Barry et al., 2014). ! 6! In Chapter 3 [Peterson, M., Kelley, K., Cottrell, L., Saal, A., Kurz, M. (in prep). The oxidation state of Fe in glasses from the Galapagos Archipelago: variation in oxygen fugacity as a function of mantle source] we build on the results of Chapter 2 by taking a subset of submarine glass chips measured for major, trace, and volatile contents and radiogenic isotopes from across the archipelago and using μ-XANES to measure new Fe3+/ΣFe ratios, proxy for ƒO2 of the lavas. Oxygen fugacity (ƒO2) is an intensive thermodynamic property that controls the speciation and partitioning of multi-valent elements and affects igneous processes such as magmatic phase relations during melting and crystallization, depth of melt initiation and the composition of volcanic gases (e.g. Ballhaus, 1993, Holloway & Blank, 1994, Canil, 1999). Our analysis reveals that Fe3+/ΣFe ratios vary with the influence of shallow level processes, generally increasing with magmatic differentiation, decreasing with sulfur degassing and increasing with assimilation of hydrothermally altered material. After taking these processes into account, however, there is still variability in Fe3+/ΣFe ratios of the isotopically different groups identified in Chapter 2. This translates into a magmatic ƒO2 that ranges from ΔQFM = +0.16 to +0.74, between the ranges identified for MORB (-0.25 - +0.4) and arc basalts (+0.2 - +1.8) (Kelley & Cottrell, 2009, Cottrell & Kelley, 2011, Kelley & Cottrell, 2012, Cottrell & Kelley, 2013, Brounce et al., 2014). In general, the Sierra Negra/Cerro Azul component is more oxidized than the Pinta/Wolf, Darwin component. Both the Siera Negra/Cerro Azul and Pinta mantle components are hypothesized to contain recycled material. The difference in oxidation state between these two sources and the correlation between Fe3+/ΣFe ratios and Pb isotopes suggest a difference in mantle lithology and therefore a difference in the type or influence of the recycled material on ƒO2. The high ! 7! 3 He/4He end-member is shown to have the most oxidized glasses (HHe = 0.175 ± 0.006) of Chapter 3. The combination of the high C/3He ratios identified in Chapter 2 and the high Fe3+/ΣFe ratios indicates that the high 3He/4He mantle source is both more carbonated and oxidized than previously thought for a primitive mantle component. In Chapter 4 [Peterson, M., Wang, Z., Saal, A., Kurz, M., Eiler, J. (in prep). Oxygen isotopes in lavas from the Galapagos Archipelago: evaluating contributions from ancient and present-day crustal sources], we apply oxygen isotopes in olivine phenocrysts by laser fluorination in lavas collected from across the Galapagos Archipelago to isolate the deep mantle signatures from the effects of shallow level processes. Oxygen isotopes, in comparison to methods employed in Chapters 1-3, are uniquely suited to separate deep from shallow signatures due to the fact that measurable oxygen isotope fractionations only occur at low temperatures. As a result, deviations in δ18O from the mantle source value (average upper mantle olivine = 5.0-5.4‰, Mattey et al., 1994, Eiler, 2001) are inferred to be the result of the presence of material that at one time resided on or near the surface. We find that Fernandina, Floreana, and Pinta, which represent three isotopic end- members of the Galapagos hotspot, have δ18O values = 5.02 ± 0.08‰, within range of the oxygen isotope value of olivine in the upper mantle (Mattey et al., 1994, Eiler, 2001). δ18O values do not correlate with radiogenic isotopes, but there is an overall trend of increasing Sr/Sr* and Ba/Nb ratios (indicators of the assimilation of plagioclase cumulates) with δ18O values for the islands of San Cristobal, Sierra Negra, Darwin, and Santa Cruz. This suggests that the highest δ18O values measured in olivines from San Cristobal (5.34 ± 0.05‰) and the large ranges in δ18O values identified for Sierra Negra, Darwin Volcano and Santa Cruz are the result of interaction with the shallow crust. Sierra ! 8! Negra, Alcedo, Darwin Volcano, Wolf Volcano and Ecuador, which have lower than upper mantle δ18O values (4.85 ± 0.1‰), do not show any correlations with radiogenic isotopes. For other ocean islands lower than upper mantle values for oxygen isotopes attributed to hydrothermally altered material in the source is accompanied by extreme isotopic heterogeneity (Eiler, 2001, Skovgaard et al., 2001, Gurenko & Chaussidon, 2002, Cooper et al., 2004, Thirlwall et al., 2006, Day et al., 2009, Day et al., 2010, Gurenko et al., 2011). Therefore, the low δ18O values in the Galapagos are either being produced by a low δ18O component with intermediate isotopes, in contrast to what is observed elsewhere, or are produced by interaction with a low δ18O component in the shallow lithosphere. The olivines for these islands come from lavas with lower MgO that are more evolved. This suggests that the low δ18O values of the olivines are due to the assimilation by the melt of a low δ18O component in the volcanic edifice or shallow lithosphere. The lavas from these volcanoes, however, do not show unusual trace element or radiogenic isotope characteristics, making it difficult to unambiguously prove that assimilation is occurring in these samples. 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CHAPTER 1 Origin of the “ghost plagioclase” signature in Galapagos melt inclusions: new evidence from lead isotopes M. E. Peterson1,*, A. E. Saal1, E. Nakamura2, H. Kitagawa2, M. D. Kurz3, A. M. Koleszar4 1 Department of Earth, Environmental and Planetary Sciences, Brown University, Providence, Rhode Island, 02912, United States 2 The Pheasant Memorial Laboratory for Geochemistry & Cosmochemistry, Institute for Study of the Earth’s Interior, Okayama University, Misasa, 682-0193, Japan 3 Department of Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, 02543, United States 4 Department of Geological Sciences, University of Texas at Austin, Austin, Texas, 78712, United States Published in Journal of Petrology, 55 (11) 2193-2216, 2014 ! 21! ABSTRACT Olivine-hosted melt inclusions from both Fernandina and Santiago islands in the Galapagos Archipelago have compositions indicating that plagioclase played an important role in the magmatic evolution of these volcanic islands. Major and trace element chemistry of Santiago melt inclusions indicates simple plagioclase assimilation. In contrast, Fernandina inclusions have compositions where the plagioclase appears to be only present as a ‘ghost’ trace element signature (i.e. “ghost plagioclase” signature). Two competing hypotheses have been proposed to explain this unique composition: 1) incorporation of an ancient recycled plagioclase-rich cumulate (eclogite) into the mantle and 2) shallow-level interaction between melts and plagioclase-rich cumulates in the present-day lower oceanic crust. Here we present new Pb isotope measurements of olivine-hosted melt inclusions from Fernandina and Santiago islands to distinguish between the two models. The new Pb isotope ratios are within the range previously reported for whole rock basalts from those islands. Melting and mixing models of ancient (~0.5-1 Ga) recycled plagioclase-rich cumulates cannot reproduce the observed trace element and Pb isotopic characteristics of Fernandina melt inclusions with a ghost plagioclase signature. Shallow-level diffusive interactions between basalt and present- day plagioclase-rich cumulates provide the simplest explanation for the observed trace element compositions and Pb isotope ratios of melt inclusions from Fernandina and Santiago islands. Keywords: Galapagos; ghost plagioclase; mantle plume; melt inclusion; Pb isotopes ! 22! INTRODUCTION The magmatism of the Galapagos Archipelago (Fig. 1) has conventionally been divided into four end-member components based on Sr, Nd, Pb, Hf and noble gas (He, Ne) isotopes (Vicenzi et al., 1990, Bow & Geist, 1992, White et al., 1993, Reynolds & Geist, 1995, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Geist et al., 2002, Saal et al., 2007, Kurz et al., 2009, Gibson et al., 2012). The end-member compositions are expressed at individual islands although several of these end-members are partially expressed at multiple volcanoes in the Archipelago. The two enriched mantle components are best represented by lavas from Pinta and Floreana islands, the high 3 He/4He component is characteristic of Fernandina lavas and the depleted component, similar to MORB, is present in lavas from Genovesa Is. (Fig. 2). Previous geochemical studies of olivine-hosted melt inclusions and high-MgO basalts from the eastern central region of the Galapagos Archipelago show major, trace and isotope signatures consistent with interaction between basalt and plagioclase-rich cumulates during melt percolation through the oceanic lithosphere (Saal et al., 2007, Koleszar et al., 2009). In contrast, geochemical data of Fernandina and Isabela Island basalts in the western region of the archipelago do not show evidence of plagioclase assimilation. However, some Fernandina melt inclusions show strong trace element evidence of plagioclase assimilation, but the associated major-element signature of this process is absent. Detailed trace element studies of basalts and melt inclusions from several oceanic islands (including the Galapagos Archipelago) and mid-ocean ridges have shown the presence of an unusual composition characterized by Sr, Ba, and Eu anomalies greater than one on primitive mantle normalized diagrams (Gurenko & Chaussidon, 1995, ! 23! Hofmann & Jochum, 1996, Kamenetsky et al., 1998, Yang et al., 1998, Chauvel & Hemond, 2000, Sobolev et al., 2000, Huang et al., 2005, Kent et al., 2002, Danyushevsky et al., 2003, Huang et al., 2005; Ren et al., 2005, Maclennan, 2008). Melting or crystallization will not significantly fractionate these elements from elements of comparable incompatibility unless plagioclase is present during these processes. Although the observed geochemical signature is typical of plagioclase-rich cumulates, many of these lavas and associated olivine-hosted melt inclusions show no indication of plagioclase accumulation or assimilation in their major element concentrations (e.g., uncommon enrichment of Al2O3 and depletion of FeO with decreasing MgO contents). Thus, this unusual composition was named the “ghost plagioclase” signature (Sobolev et al., 2000). Most lavas and melt inclusions with a ghost plagioclase signature are characterized by very primitive compositions (e.g., high-MgO contents). This suggests that a plagioclase-rich component is required either in the melt source region or in the crust through which the melt percolates. Two main hypotheses have been proposed to explain this signature. The first considers the presence of an ancient recycled oceanic crust with plagioclase-rich cumulate gabbro (now eclogite) as a component intrinsic to the convecting mantle (Hofmann & Jochum, 1996, Chauvel & Hemond, 2000, Sobolev et al., 2000, Kent et al., 2002, Ren et al., 2005). The second argues for the interaction of melts with a plagioclase-rich cumulate during melt percolation through the present-day oceanic lithosphere (Danyushevsky et al., 2003, Danyushevsky et al., 2004, Saal et al., 2007, Maclennan, 2008). Trace element contents of a 500-m gabbroic section of the oceanic crust drilled at site 735, Leg 118, have a clear plagioclase-rich cumulate ! 24! signature (“protolith” in Hart et al. (1999)). If this composition represents the lower oceanic crust, then both hypotheses give a reasonable explanation for the ghost plagioclase signature observed in oceanic basalts. Pb isotopes provide an effective tool for determining which model best explains the ghost plagioclase signature. Basalts generated from a source containing an ancient (~0.5-1 Ga) recycled plagioclase-rich cumulate gabbro (now eclogite) will have a much higher 207Pb/206Pb and 208Pb/206Pb than present-day end-member MORB. This is due to the lower (U, Th)/Pb ratio of a plagioclase-rich cumulate (0.05 U/Pb and 0.114 Th/Pb based on average ‘protolith’ of Hart et al. (1999)) compared to that of the Earth’s depleted upper mantle (0.18 U/Pb and 0.439 Th/Pb based on DMM composition of Workman and Hart (2005)). In contrast, basaltic magmas percolating through the present-day oceanic 207 lithosphere will obtain, at most, similar Pb/206Pb and 208 Pb/206Pb ratios to those of the plagioclase cumulate with which they interact. In this report we present new Pb isotope data from geochemically well- characterized olivine-hosted melt inclusions from Santiago and Fernandina islands (representing the central/eastern and western regions of the Galapagos Archipelago, respectively) to test the origin of their ghost plagioclase trace element signature. Our Pb isotope data are consistent with the interaction between melts and a present day plagioclase-rich cumulate in the oceanic lithosphere. Previous results Samples ! 25! We used olivine-hosted melt inclusions picked from 2 samples collected from Fernandina and Santiago island basalts (Fig. 1). The Fernandina sample is AHA D25C, dredged from the western submarine flank of Fernandina (00°27.2'S, 91°44.6'W, average 2750 meters depth) during the AHA-NEMO2 cruise in 2000. This dredge is part of the normal series lavas described in Geist et al. (2006) with a composition similar to the aphyric subaerial lavas erupted in Fernandina (Allan & Simkin, 2000). Both the subaerial and submarine lavas of Fernandina are noted for their compositional homogeneity, which is thought to be the result of mixing in a common magma chamber within the crust below Fernandina (Allan & Simkin, 2000, Geist et al., 2006). The Santiago sample, STG06-29, was collected from a subaerial hornito (0°19’5.2”S, 90°35’1.7”W) on the eastern side of Santiago, which is characterized by the MORB-like depleted component (Baitis & Swanson, 1976, White et al., 1993, Koleszar et al., 2009, Gibson et al., 2012). Olivine- hosted melt inclusions were hand picked from crushed basalts under a binocular microscope, mounted in indium metal, exposed and then polished using diamond pastes down to 0.25 µm (Fig. 3). The melt inclusions are naturally quenched to a homogeneous glass and have shapes that indicate a primary origin (Sobolev, 1996). They range in size from 50 to 250 µm in diameter, with Santiago’s inclusions being, on average, smaller (~60 µm) than those from Fernandina (~125 µm). Major and Trace Element Compositions The major and selected trace element contents of the melt inclusions, matrix glass and olivine host were reported in Koleszar et al. (2009). Complementary trace element data are reported here. The trace element data corrected for host olivine crystallization are ! 26! included in Supplementary Data Electronic Appendix S1 (supplementary data are available for downloading at http://www.petrology.oxfordjournals.org). Analytical errors (2σ) for triplicate analyses of representative melt inclusions are <3% for SiO2, CaO, Al2O3, <5% for MgO, TiO2, and FeO and <9% for Na2O. For the olivine hosts the error is <2% for SiO2, MgO, and FeO. Errors for the trace elements, determined from replicate analyses of matrix glasses and large melt inclusions, are typically <5% (2σ). Major and trace elements are corrected for olivine crystallization on the interior wall of the inclusion by numerically adding host olivine back into the melt inclusion composition in 0.1% increments until the inclusion and olivine host are in MgO-FeO equilibrium. Samples required only 2.6-16 %, averaging 7.0% olivine addition. For further information on the analytical techniques and corrections refer to Koleszar et al. (2009). Olivine phenocrysts hosting the melt inclusions have Mg# (molar, MgO/[MgO+FeO]*100) averaging 85±2 in the Fernandina sample and 86.6±0.8, with only two exceptions having Mg# of 82 and 85, in the Santiago lava (Koleszar et al., 2009). In both samples, the Mg# of the olivine in equilibrium with the matrix glass is 78 and 82 for Fernandina and Santiago olivines respectively, indicating FeO–MgO disequilibrium between the phenocrysts and the host glasses. The role of plagioclase: evidence from major and trace elements Santiago Inclusions: After correction for olivine crystallization, both the major and trace element compositions of the Santiago inclusions evidence direct assimilation of plagioclase. Assimilation can be defined by a large decrease in FeO* (FeO* indicating total Fe) and a large increase in Al2O3 with decreasing MgO contents, trends that cannot ! 27! be reproduced by crystal fractionation processes (Fig. 4) (Koleszar et al., 2009). Furthermore, almost all inclusions possess positive Sr, Ba, and Eu anomalies in a primitive mantle normalized diagram (Fig. 5) (Koleszar et al., 2009). The magnitude of these anomalies correlates with Al2O3 contents, however they do not show any clear trend with very incompatible trace element ratios such as Nb/La suggesting a common source region (Fig. 6). If the variation in Al2O3 were the result of a different process from simple plagioclase assimilation, for example different degrees of melting, we would expect a large range in trace element ratios of highly to moderately incompatible elements (e.g., La/Yb) associated with the Al2O3 content. For most of the Santiago inclusions this correlation is not observed (Koleszar et al., 2009). Inclusion STG06-29-19 represents an extreme case, with the lowest incompatible trace element contents and the strongest Sr and Eu anomalies. Only one Santiago inclusion (STG06-29-10, labeled Santiago enriched composition in the figures) has low Al2O3 and high FeO contents at low MgO coupled with a Sr anomaly lower than one, indicating possible plagioclase fractionation prior to entrapment. Fernandina Inclusions: Although major element compositions of Fernandina melt inclusions indicate fractional crystallization, trace element compositions from a subgroup of inclusions have a ghost plagioclase signature. The major element contents of the inclusions, when combined with a compilation of Fernandina submarine glasses (Geist et al., 2006), can be simply explained by olivine, olivine-plagioclase, and olivine- plagioclase-clinopyroxene crystallization (Fig. 4). However, trace element contents of the inclusions define two distinct groups: one with an enriched incompatible trace element composition typical of Fernandina lavas and a second group with an incompatible trace ! 28! element depleted composition. Eight of the eleven inclusions with depleted trace element compositions have positive Sr, Ba, and Eu anomalies in primitive mantle normalized diagram (Fig. 5 and 6), but their major element contents are inconsistent with the assimilation of plagioclase (i.e. ghost plagioclase signature). The magnitudes of Sr, Ba, and Eu anomalies do not correlate with Al2O3 contents (Fig. 6). However, a weak negative correlation exists between Sr/Sr* (SrPM/[(CePM+NdPM)/2]) and Nb/La ratios, which indicates the presence of a trace element depleted melt in the creation of the ghost plagioclase signature. LEAD ISOTOPE ANALYTICAL METHODS Pb isotope analyses (Table 1) were performed using the Cameca 1270 SIMS at the Institute for the Study of the Earth’s Interior, Okayama University, Misasa, Japan. In preparation for using the ion probe, mounts were submerged in 0.5M HNO3 for 5 minutes to remove any sources of contamination from the surface. The samples were rinsed with MilliQ water three times, left under vacuum overnight to dry and coated with 300 Å of gold. The samples were sputtered using a 30 nA focused beam of O- ions resulting in a 30 µm diameter pit. The primary beam (acceleration) energy was 23 keV, and the secondary accelerating voltage was 10 keV. Samples were analyzed at a mass resolution of 3000. Pb isotopes were collected using a multi-collection system equipped with five electron multipliers following the methods described in Kobayashi et al. (2004). Each analysis consisted of 3 minutes of pre-sputtering followed by 400 cycles of 10 seconds on 206 Pb peaks. Mass positions were set by centering the Pb peak at the beginning of each ! 29! analysis due to the high concentration of Pb in the standard. Background runs were measured using the settings of the run, but without the primary beam emission, and routinely checked throughout the session. We used the synthetic basaltic glass standard FMSA (5 ppm Pb) to correct for the instrumental mass discrimination (Kobayashi et al., 2004). The standard was measured repeatedly (≥9 times/day) over the course of the session. The daily external precision for the standard (2σ standard deviation of daily group of standard measurements) ranges 207 from ± 0.27 (n = 16) to 0.35% (n = 9) for Pb/206Pb and 0.24 (n = 9) to 0.35% (n = 16) for 208Pb/206Pb. The Pb concentrations of the melt inclusions are all <2ppm, leading to low precision in the measured isotopic ratios. In the case of Santiago inclusions, with Pb concentrations of 0.6 to 0.9 ppm, the 2σ standard error in Pb isotope ranges from ± 0.6 to 1.5% for 207Pb/206Pb and ± 0.5 to 1.1% for 208Pb/206Pb. For Fernandina inclusions, with Pb concentrations of 0.5 to 1.5 ppm, the errors are ± 0.4 to 2% and ± 0.3 to 1.6% 207Pb/206Pb 208 and Pb/206Pb respectively. The low 204 Pb ion current yielded large uncertainties in the 204 Pb-normalized isotope ratios, rendering these ratios unusable. RESULTS The Pb isotopic compositions of the melt inclusions are consistent with those previously reported on whole rock samples from Fernandina and Santiago islands (White et al., 1993, Harpp & White, 2001, Saal et al., 2007, Gibson et al., 2012) (Fig. 7). 207 For Santiago inclusions, Pb/206Pb ranges from 0.8261 to 0.8473, and 208 Pb/206Pb from 2.0335 to 2.0678, which is within error of the range seen for Santiago whole rock ! 30! samples (207Pb/206Pb = 0.8133-0.8381 and 208 Pb/206Pb = 2.0237-2.0569) (White et al., 1993, Saal et al., 2007, Gibson et al., 2012). There is no significant correlation between the Pb isotopes and the major and trace element contents indicative of plagioclase assimilation typical of these inclusions (Fig. 8). Inclusion STG06-29-10 (Santiago enriched inclusion with major and trace element evidence for plagioclase fractionation) has Pb isotope ratios within the range of the typical Santiago inclusions. STG06-29-19 (highest Sr anomaly) was too small in size to obtain reliable Pb isotope data. In the case of Fernandina inclusions, 207Pb/206Pb ranges from 0.8113 to 0.8240 and 208 Pb/206Pb from 2.010 to 2.042, which is within error of the range for Fernandina whole rock isotopic data reported in the literature (207Pb/206Pb = 0.8136-0.8168 and 208 Pb/206Pb = 2.0281-2.0316) (White et al., 1993, Kurz & Geist, 1999, Saal et al., 2007). The average of the melt inclusion range (207Pb/206Pb = 0.8174, 208 Pb/206Pb = 2.0315, weighted for Pb concentration) falls close to the value of the host lava AHA D25C (Fig. 7). This is consistent with the host lava being an aggregate of a heterogeneous mixture with a varied geochemical composition that melt inclusions are sampling more fully (Saal et al., 2005, Maclennan, 2008). There is no discernable correlation between the Pb isotopes, major, or trace element compositions of the inclusions (Fig. 8). The inclusions with depleted incompatible trace element contents we were able to measure have, at similar 207Pb/206Pb, unusually high 208Pb/206Pb. However, the large errors in these two analyses, due to the low Pb concentrations of the inclusions, do not allow us to discuss the implications of these results any further. From the eight inclusions with ghost plagioclase signature only four were large enough to measure the Pb isotopes. These four inclusions have 207Pb/206Pb and 208 Pb/206Pb within the range of the other Fernandina inclusions. ! 31! DISCUSSION Evidence for contamination by plagioclase cumulates: relationship between major, trace elements and Pb isotopes Santiago melt inclusions (normal and enriched) have Pb isotopes that are within the range of those previously measured on whole rock lavas from this island indicating the presence of both trace element and Pb isotope enriched and depleted melt components (White et al., 1993, Gibson et al., 2012). Although the melt inclusions have major and trace element evidence for the assimilation and crystal fractionation of plagioclase, their Pb isotope ratios do not significantly correlate with any chemical indicators of these processes. There is, if any, only a very weak positive trend between Pb isotopes and both Al2O3 content and Sr/Sr* (quantified by SrPM/[(CePM+NdPM)/2]) (Fig. 8). This indicates that the inclusions with depleted trace element and high 207Pb/206Pb, 208Pb/206Pb have been more affected by assimilation of plagioclase cumulates than the enriched inclusions, consistent with the hypothesis proposed by Saal et al., (2007) using the whole rock geochemical data of basalts from the Galapagos Archipelago. The assimilation of plagioclase is responsible for the observed range in Sr/Sr* and Nb/La seen in the normal Santiago inclusions (Fig. 6); this process changes the Sr/Sr* without significant modification of the Nb/La ratio in the melt. The exception is the depleted Santiago inclusion that has a Nb/La ratio much lower than the normal Santiago inclusions (0.4 compared to 0.8, respectively). This ratio cannot be explained by a normal Santiago melt undergoing simple plagioclase assimilation and therefore requires the presence of a more depleted melt (that will possess a low Nb/La) within the magmatic ! 32! plumbing system beneath Santiago. It is unfortunate that the depleted Santiago inclusion (STG06-29-19) was too small in size to obtain reliable Pb isotope data. The Pb isotopic compositions of Fernandina inclusions fall within error of the range defined by the lavas from this island. This is expected for inclusions having major and trace element compositions analogous to Fernandina whole rock lavas. However, inclusions with a ghost plagioclase signature have Pb isotopes indistinguishable from those of normal Fernandina inclusions (Fig. 7 and 8). This suggests the process producing ghost plagioclase signatures in Fernandina inclusions also does not significantly affect Pb isotope ratios of the melt. Evaluating the two theories for the origin of the Fernandina “ghost plagioclase signature” With the new Pb isotopic constraint it is possible to evaluate the 2 proposed models for the origin for the ghost plagioclase signature: ancient recycled versus interaction with present-day plagioclase-rich cumulate. While both models are indistinguishable with respect to their trace/major element concentrations, their Pb-isotopes will be significantly different. Here, we evaluate both models for their ability to reproduce both the expected trace element and Pb isotopic signatures of the Fernandina inclusions. Estimating the Pb isotope composition of ancient plagioclase cumulates To estimate the Pb isotope ratios of an ancient (~ 0.5-1 Ga) recycled plagioclase cumulate, we assume the cumulate is formed in the oceanic lithosphere during the crystallization of a range of MORB compositions. Using present-day Pb isotopes of ! 33! Pacific MORB and the U, Th, Pb contents estimated for the MORB mantle (E-DMM, N- DMM and D-DMM Donnelly et al., 2004, Stracke et al., 2005, Workman & Hart, 2005), we calculate Pb isotopes of the plagioclase-rich cumulate at the time of formation (~ 0.5- 1 Ga). We then use these calculated Pb isotopes with the measured U, Th and Pb concentrations of plagioclase-rich cumulates in the oceanic crust (Hart et al. 1999, Zimmer et al., 1995) to calculate its present-day Pb isotopic compositions. Present day Pb isotope ratios of the ancient plagioclase cumulate calculated by this model are inconsistent with the values measured in Fernandina inclusions with ghost plagioclase signatures (Fig. 9). The main variable parameters that control the isotopic evolution of the recycled gabbro in the calculation described above are the present-day isotopic and trace element composition of the MORB source, the trace element composition of the gabbro, and the time. The boundaries of the isotopic composition of the MORB source are defined by a compilation of Pacific MORB isotope data (207Pb/206Pb = 0.82-0.88, 208 Pb/206Pb = 2.026- 2.120) (Stracke et al., 2005). These boundaries contain enriched (E), normal (N), and depleted (D) MORB isotopic values. As such, we change the U, Th, and Pb contents with the isotopic values to reflect the difference in trace element enrichment of EMORB versus N and D-MORB sources (Donnelly et al, 2004, Workman and Hart, 2005). For the trace element composition of the gabbro we use reported U, Th, and Pb contents measured for what are considered pristine, unaltered protolith (i.e., Hp) and variably hydrothermally altered strip samples (i.e., Ps, Qs, Os) of lower crustal gabbros that show trace element evidence for plagioclase-rich cumulates reported in Hart et al., (1999) and an average Gabbal Gerf plagioclase-rich gabbro composition (GG) as an additional ! 34! composition to test (Zimmer et. al., 1995). Therefore both a range in U/Pb (0.018-0.176) and Th/Pb ratios (0.032-0.484) were input into the calculation. We test ages of 500 Ma and 1Ga for the recycled gabbro. The 500 Ma minimum time of evolution is a conservative average minimum recycling time found through simultaneous analysis of multiple isotopic systems of other hotspots (e.g. McKenzie et al., 2004), while material older than 1Ga did not show significant result improvement. All parameters are listed in Table 2 and results are shown in Fig. 9. Though subduction may affect U, Th, and Pb contents of plagioclase-rich cumulate gabbros, what precisely this affect is remains largely unconstrained. The study of Becker et al. (2000) approached this problem by studying high-pressure eclogite bearing terranes. Those samples with a gabbroic protolith were comparable to samples gathered from the lower oceanic crust, which suggests subduction does not have a significant affect on U, Th, and Pb contents of subducted plagioclase cumulates. Other studies have tried to quantify trace element flux in a subduction zone using mass balance calculations (e.g. Kelley et al., 2005, Chauvel et al., 2009, Porter & White, 2009). Porter and White, (2009), using data from 8 different subduction zones for their trace element flux calculation found that, on average, 73-79% of U, Th, and Pb survived in to the deep mantle making changes to U/Pb, Th/Pb and Th/U small and not systematic. Therefore, for our model, we consider the U/Pb and Th/Pb ratios of the recycled plagioclase cumulate to be not significantly affected by subduction processes, which is consistent with the model of Sobolev et. al., (2000). Using the parameters described above, the results of the model considering an ancient recycled plagioclase-rich cumulate do not produce isotopic values within what ! 35! has been found for Fernandina melt inclusions (Fig. 9). In order for a recycled gabbro to evolve to a Fernandina isotopic composition, the present-day MORB isotopic value has to be equal to the most enriched EMORB in the Pacific Ocean (which has a similar isotopic composition to Fernandina lavas; Fig. 9) and experience an incredibly short recycling time (≤152 Ma). This would avoid large changes in the initial isotopic value, but requires an improbably young recycled material. Alternatively, to stay within the more reasonable minimum age of recycling (500 Ma) we would have to expand the range in present-day MORB isotopic values to include the most enriched samples measured for Atlantic EMORB, samples which have been affected by the Sierra Leone plume (Schilling et al., 1994). The very young age and/or the extreme isotopic composition necessary to reproduce the Fernandina Pb isotope values makes it unlikely that melt inclusions with the ghost plagioclase signature are sampling an ancient recycled plagioclase-rich cumulate in the source of the Galapagos plume. An additional melting/mixing step is needed, which will require us to expand our model to include trace elements in addition to Pb isotopes. Melting of an ancient plagioclase cumulate: adding Pb isotopes to a critical melting model To test whether melting of a plume mantle containing an ancient recycled plagioclase- rich cumulate could reproduce the trace element and Pb isotopic composition found in Fernandina melt inclusions with the ghost plagioclase signature, we use the model presented by Sobolev et al. (2000). The melting reactions, phase proportions and partition coefficients are presented in Table 3. This model considers that a quartz-eclogite ! 36! (recycled plagioclase-rich cumulate) entrained in a mantle plume starts to melt before the surrounding mantle during adiabatic upwelling. The resulting Si-rich melt infiltrates and reacts with the surrounding peridotite to form pyroxene (Yaxley & Green, 1998) transferring the trace element signature of the recycled plagioclase-rich cumulate to the peridotite. The subsequent melting of this hybrid mantle in the garnet stability field could then produce melts characterized by a ghost plagioclase trace element signature (Sobolev et al., 2000). Although Fernandina mantle has noble gasses consistent with a less degassed mantle composition (FOZO), we consider for the modeling two different mantle compositions, a trace element depleted DMM (Workman & Hart, 2005) and a primitive mantle (PM, Mcdonough & Sun, 1995) to better simulate a large range of possibilities. For the isotopes of the DMM mantle, we use the Pb isotopes of DMM from Stracke et al. (2005). For the PM mantle we use the average Pb isotopic composition for lavas from Fernandina (Saal et al., 2007). For the remaining parameters (composition of ancient plagioclase, degree of melting for the eclogite, fraction of eclogite melt to peridotites used for mantle hybridization, and the degree of melting of the hybrid mantle) we test a range of values and present the results of the best-fit models in Fig. 10 and Table 4. The ancient recycled plagioclase-rich cumulate isotopic composition is calculated using the method and compositional ranges described in the previous section. The parameters for the cumulate calculation that yield the isotopic composition closest to the Fernandina 207 melt inclusions isotopic composition are the extreme Pb/206Pb and 208 Pb/206Pb of the EMORB Pacific ridge samples and the composition of either the gabbro strip Os (Hart et. al., 1999) or the average Gabal Gerf gabbro composition GG (Zimmer et al., 1995) at an ! 37! age of 500 Ma and 1Ga. Following Sobolev et al. (2000) we consider that the eclogite melting has a residual porosity of 8%, which is reasonable for a high-Si melt (Yaxley and Green, 1998). Any fraction of melt above 8% will be removed to hybridize the surrounding mantle (either DMM or PM) at a ratio 0.1 melt to 0.9 peridotite. Though we tested our model using 9-50% instantaneous and aggregate eclogite melt, the results presented in Fig. 10 use the first weight percent past the porosity (F = 9%) of instantaneous melt for hybridization because this produces the largest Sr anomaly. We also varied the proportion of eclogite melt to peridotites from 0.1:0.9 to 0.03:0.97, though this does not produce a significant change from what we show in Fig. 10. Following hybridization, we consider 5-15% accumulated critical melting of the hybrid mantle with 2% porosity. The resulting melt extracted from the mantle is mixed with typical Fernandina melt, in 10% increments. For the composition of the modeled melt originated from the hybrid mantle, the differences between 5 and 15% melting will not significantly affect the Sr/Sr* ratios, therefore we present an average value of 10% melting in Fig. 10. Results of each step of this model for the ideal parameters indicated above can be found in Table 4. None of the models were able to reproduce both the trace element and Pb isotope compositions of the ghost plagioclase inclusions. Over the range of tested parameters, the percentage of the eclogite needed to create the ghost plagioclase trace element signature force the Pb isotopes outside what we measured in the Fernandina inclusions with ghost plagioclase signature. Present-day formation of a ghost plagioclase signature ! 38! In order to explain the similar Pb isotopic composition of the Fernandina melt inclusions with and without the ghost plagioclase, the process forming the ghost plagioclase signature must create both the observed trace element anomalies and isotopic ratios. For example, a plagioclase cumulate formed in the oceanic crust beneath Fernandina within the last few million years would have the present-day Pb isotope composition of Fernandina lavas. Therefore, a shallow-level diffusive interaction of a Fernandina melt with this plagioclase-rich cumulate would create the needed trace element anomalies while not affecting the Pb isotopes and the major element composition of the melt. To illustrate this, we’ve included a simple diffusion model that considers the approximate solution to the diffusion equations in a finite crystal-melt geometry (Hart, 1993) presented in Saal et al., (2007) (Fig. 11). This model considers an end-member case of melt percolating through a troctolite mush with equal proportions of plagioclase and olivine. For simplicity, we neglect the effects of crystal dissolution and precipitation that may occur during diffusive re-equilibration of a partially molten system (Hart, 1993, Liang, 2003), namely we assume the melt is in major element equilibrium with the plagioclase cumulate (Saal et al., 2007). As a result, because of the low trace element budget of olivine, variations in the trace element content of the melt are due solely to diffusive exchange between plagioclase and melt. There are 2 values for melt concentration that we must consider: the melt from which the plagioclase cumulate crystalizes and the composition of the melt interacting with the cumulate. The geochemistry of the Fernandina melt inclusions with ghost plagioclase signature (large positive Sr/Sr*, large range in Nb/La ratios, and Pb isotopes equal to those of Fernandina whole rock) limit what these melt compositions can be. A ! 39! large positive Sr anomaly, as is observed in the Fernandina melt inclusions with the ghost plagioclase signature, can be created by diffusion if the melt from which the plagioclase crystallized is much more enriched in trace element concentrations than the percolating melt. Therefore, for the cumulate composition, we considered the crystallization of plagioclase from an evolved Fernandina melt (AHA31A, Geist et al., 2006) and we assume a Pb isotopic composition equivalent to that of Fernandina whole rock (Saal et al., 2007). For the percolating melt we considered trace elements depleted melt because the Nb/La ratio in the depleted Fernandina melt inclusions with and without plagioclase ghost signature is much lower than those of the “normal” inclusions (Fig. 6 and 11). Although part of the Nb/La variation can be reproduced by the diffusive interaction of a percolating melt with the plagioclase cumulate due to the different diffusivity of Nb and La in plagioclase (Cherniak, 2003, Saal and Van Orman, 2004), the total variation in Nb/La ratios cannot be reproduced using reasonable parameters, suggesting that the percolating melt is depleted (i.e., low Nb/La). Therefore, for the percolating melt composition, we assumed a 12% aggregated fractional melt of the DMM composition (Workman and Hart, 2005) with a Nb/La ratio of 0.773 and with the average isotopic composition of Pacific MORB (Stracke et al., 2005) (Fig. 9). Although we do not see erupted lavas with trace element depleted composition in Fernandina, the trace element depleted melt inclusions suggest the presence of low volume, trace element depleted melt composition in the plumbing system of this volcano. In a crystal mush zone, either in a shallow crustal magma chamber or shallow regions of the mantle, the melt is expected to percolate by both reactive porous flow (low melt/rock ratio more susceptible to diffusive interaction with the crystal cumulate) and ! 40! channelized melt flow (high melt/rock ratio where the composition of the melt is similar to the erupted lavas) (e.g. Jaupart and Tait, 1995; Kelemen et al., 1995). Therefore in a mush zone beneath Fernandina we would expect the presence of these two melts (channelized and reactive porous flow), which will progressively aggregate as presented in Fig. 11 reproducing the correlation between Nb/La and Sr/Sr* that is observed in the ghost plagioclase inclusions. Other parameters include temperature, crystal size and melt to rock ratio. Based on the results of previous studies of the crust-mantle transition zone we set a temperature of 1200°C, grain size of 1 mm and use a melt to rock ratio of 1% (Boudier et al., 1996, Dunn et al., 2001). We further assume the temperature is constant, which is consistent with inefficient heat removal from the crust mantle transition zone (Dunn et al., 2001). For the model we assume the solid is homogenous at the onset of diffusion. We also assume that the concentration in the middle of the grains is controlled by diffusion and that the mineral grains do not deform. Diffusion in the melt is considered instantaneous and the boundary between solid and melt is assumed to be in chemical equilibrium at all times (see Saal et al., 2007 for further considerations). The values of all parameters used can be found in Table 5. The results of this model (Fig. 11) demonstrate how the ghost plagioclase signature can be produced by diffusive interaction between a depleted melt and plagioclase cumulates and then the subsequent mixing of reacted melt with a normal composition Fernandina melt. The diffusivities of Sr and Pb, along with those of Ba, K and Eu+2, are several orders of magnitude faster than the other trace elements (e.g., Th, Nb and REE), allowing a melt to develop a plagioclase-rich signature within a short ! 41! amount of time (Behrens et al., 1990, Cherniak and Watson, 1994, Giletti and Casserly, 1994, Cherniak, 1995, Giletti and Shanahan, 1997, Cherniak, 2002, Cherniak, 2003, Saal and Van Orman, 2004). The fast diffusion of Pb also causes the rapid change in the Pb isotopes of the melt from the initial DMM value to a value within range of Fernandina. The rapid re-equilibration of Pb isotopes of the melt, however, necessitates that plagioclase cumulates have the Pb isotope composition of Fernandina. If we consider the opposite situation, where the plagioclase cumulate originated from a trace and isotopic composition similar to typical NMORB (e.g., crust created at the GSC), and the percolating melt has the isotopic and trace element composition of normal Fernandina lavas, the result will be a melt with the isotopic composition of NMORB and a negative Sr anomaly, thus not reproducing either the Pb isotopes or the trace element characteristics of the inclusions with ghost plagioclase signature. Alternatively, although quite unlikely, the plagioclase-rich cumulates formed at the ridge now under Fernandina might have had the extreme EMORB composition found in the Pacific Ocean, as this is similar to Fernandina whole rock. However, there is no evidence to support the existence of an extreme EMORB composition in the lithosphere beneath Fernandina. It is clear that our simple model cannot reproduce the signature of the melt inclusion with the highest Sr/Sr* (D25C-2-29), but our model is limited by only considering diffusion. Burgess et al. (in preparation), explore an expansion of this model that includes dissolution/precipitation of plagioclase in addition to kinetic equilibration. The additional processes expand the number of starting compositions that allow a present-day interaction between melt and plagioclase cumulates to result in a ghost plagioclase signature and further highlights how interaction provides the simplest ! 42! explanation for the creation of a ghost plagioclase signature in the Galapagos Archipelago. Presence of the ghost plagioclase signature in settings outside the Galapagos The ghost plagioclase signature is a well-documented phenomenon that, despite only occurring in a small percentage of samples, can be found in suites of samples from multiple OIBs (Gurenko & Chaussidon, 1995, Hofmann & Jochum, 1996, Yang et al., 1998, Chauvel & Hemond, 2000, Sobolev et al., 2000, Kent et al., 2002, Huang et al., 2005, Ren et al., 2005, Maclennan, 2008) and MORB (Kamenetsky et al., 1998, Danyushevsky et al., 2003, Danyushevsky et al., 2004). However, very few studies report Pb isotopes in addition to trace and major elements. The study of Maclennan, (2008) reports Pb isotopes for a suite of Icelandic melt inclusions with several inclusions identified as possessing the ghost plagioclase signature. The Pb isotopes of the anomalous inclusions are within the range of isotope ratios measured for the normal composition inclusions. This led the authors to conclude that the ghost plagioclase inclusions were sampling a present-day process as opposed to an ancient recycled gabbro in the mantle source, consistent with the findings of our study. The study of Sobolev et al., (2011) reports both Sr and Pb isotopes for a suite of melt inclusions from Hawaii that include several inclusions with large positive Sr anomalies on a primitive mantle normalized trace element diagram. The inclusions with the Sr anomalies did not fall outside of the Sr and Pb isotopic range defined by the inclusions with normal compositions, but the Sr anomaly correlates positively with Al2O3 content suggesting shallow-level interaction between melts and plagioclase-rich cumulates in the present-day ! 43! oceanic lithosphere beneath Hawaii. These few examples indicate that without the Pb or any other isotope data for the melt inclusions, it is difficult to unequivocally prove a recycled origin for the ghost plagioclase signature. CONCLUSIONS The major and trace element concentrations unique to the ghost plagioclase signature have been recreated equally well by models that invoke plagioclase-rich cumulate either as an ancient recycled component in the mantle or as present-day component in the oceanic lithosphere (Sobolev et al., 2000, Danyushevsky et al., 2003, Danyushevsky et al., 2004,). Because these two processes require vastly different ages, the Pb isotopes provide an excellent tool for evaluating these distinct models. Pb isotopes from the Galapagos basalts and melt inclusions suggest that the origin of a ghost plagioclase signature in basalts from the Galapagos Archipelago is not produced by the sampling of an ancient (~0.5-1 Ga) recycled plagioclase-rich cumulate within the mantle. Instead the trace elements and Pb isotopes are most consistent with a modern interaction between mantle melts and plagioclase-rich cumulates in the oceanic lithosphere. These results demonstrate that careful evaluation of the isotopic compositions of basalts and melt inclusions is required before unusual trace element anomalies such as Sr, Ba and Eu can be used to argue for the presence of ancient recycled crust in oceanic basalts. FUNDING This work was supported by the National Science Foundation Graduate Research Fellowship [Grant No. DGE-1058262 to M.E.P] and the National Science Foundation ! 44! Division of Ocean Sciences [Grant No. 0962195]. ACKNOWLEDGMENTS We thank K. Kobayashi for his help in understanding the details associated with Pb isotope analysis performed on the SIMS at Okayama University. We thank J. Blusztajn for providing the isotopic measurements of AHA D25C host glass and Tabb Prissel for the reading of the manuscript. We also thank D. Geist and S. Huang for their thoughtful reviews. ! 45! REFERENCES Allan, J.F., Simkin, T., 2000. 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FOZO is the approximate position for where the highest 3He/4He samples from the Galapagos plot in Pb-Pb space. Light grey fields represent most Galapagos volcanoes, while dark grey fields represent the data for Santiago and Fernandina islands. Figure 3. Photomicrographs of representative olivine-hosted melt inclusions from Fernandina and Santiago basalts taken in reflected light. In olivine grains with multiple exposed inclusions, a black arrow indicates the inclusion that was analyzed. Notice that there is no significant difference between the Fernandina melt inclusions with the ghost plagioclase signature, the depleted, and the normal inclusions (similar to matrix glass). Also notice ! 57! the same is true for the Santiago normal, enriched, and depleted inclusions. The small size of STG06-29-19 (extreme depletion) prevented us from analyzing this inclusion for Pb isotopes. Figure 4. Al2O3 and FeO* (reported as total iron) versus MgO wt% for (a) Fernandina inclusions (Koleszar et al., 2009) and compiled glass data (black plus sign, Geist et al., 2006). Dark blue circles, melt inclusions with trace element compositions similar to Fernandina host glass (Fernandina normal composition); light grey circles, depleted trace element compositions; cyan circles, ghost plagioclase signature; violet star, Fernandina host glass composition. (b) Santiago inclusions (Koleszar et al., 2009) and a compilation of Santiago subaerial lavas (black cross, data from http://georoc.mpch- mainz.gwdg.de/georoc). Dark green diamonds, melt inclusions with trace element composition similar to Santiago host glass (Santiago normal composition); yellow diamond, enriched trace element composition (STG06-29-10); brown diamond, extreme trace element depletion and plagioclase signature (STG06-29-19), pink star, Santiago host glass composition. Average 2σ error shown by grey plus symbol. Crystal fractionation trends have been calculated using PETROLOG (Danyushevsky, 2001, Danyushevsky and Plechov, 2011) for both Santiago and Fernandina inclusions at a pressure of 1.5 kbars for two different concentrations of water: dashed line uses 0.1 wt% and solid line 0.4 wt% water. Starting compositions for calculations were chosen to be high MgO content inclusions D25C-3-20 and STG06-29-06 for Fernandina and Santiago inclusions respectively. Note that, in general, Fernandina inclusions define a trend that ! 58! can be produced by magmatic differentiation of olivine, followed by olivine+plagioclase, followed by olivine+plagioclase+clinopyroxene. The inclusions with the ghost plagioclase signature also fall along this trend at major element concentrations indistinguishable from the normal Fernandina inclusions. In contrast, Santiago inclusions cannot be reproduced by simple magmatic differentiation; their increase in Al2O3 and decrease in FeO* with decreasing MgO wt% cannot be reproduced by normal crystal fractionation and represents the assimilation of plagioclase of approximately An50. Figure 5. Primitive mantle-normalized trace element diagram for representative samples. (a) Fernandina: normal composition melt inclusion D25C-3-12 (MgO wt% = 8.51), depleted trace element composition D25C-2-24 (MgO wt% = 8.34), and ghost plagioclase composition D25C-3-34 (MgO wt% = 8.58). (b) Santiago: normal composition melt inclusion STG-06-29-17 (MgO wt% = 9.79), enriched composition STG06-29-10 (MgO wt% = 9.47), and most depleted trace element composition STG06-29-19 (MgO wt% = 9.10). Normalizing values are from Mcdonough and Sun (1995). Included in (a) and (b) is the lower crustal gabbro protolith (Hp) reported by Hart et al. (1999), which has one of the highest U/Pb ratios (magenta diamond). Figure 6. (a) Sr/Sr* (SrPM/[(CePM+NdPM)/2]) versus Al2O3 wt% for individual melt inclusions from both Fernandina and Santiago islands. Included inset shows a zoomed in view to better illustrate the correlation between Sr/Sr* and Al2O3 in the Santiago inclusions. (b) Sr/Sr* ! 59! vs Nb/La for Fernandina and Santiago inclusions. Included plagioclase-rich gabbro symbol (magenta diamond) is the value of protolith gabbro sample Hp (Hart et al., 1999) Figure 7. 208 Pb/206Pb versus 207Pb/206Pb ratios for individual melt inclusions from both Santiago and Fernandina islands. Results overlay the fields for Fernandina and Santiago islands for whole rock lavas (references as in Fig. 2). Also included is the Fernandina host glass AHA D25C (violet star), an average weighted for Pb concentration of both the Fernandina inclusions and Santiago inclusions (white star and labeled). Inset shows a zoomed out view of the results. Note that the inclusions with the ghost plagioclase signature plot in the same isotopic space as the inclusions from Fernandina with normal trace element compositions. Figure 8. 207 (a) Sr/Sr* (SrPM/[(CePM+NdPM)/2]) versus Pb/206Pb ratios for Fernandina and Santiago 207 inclusions. A zoomed in view of Sr/Sr* versus Pb/206Pb is also included. Note that Sr/Sr*, an indicator of plagioclase signature, does not correlate with 207Pb/206Pb ratios for 207 either inclusions from Fernandina or Santiago Islands. (b) Al2O3 wt% versus Pb/206Pb for Fernandina and Santiago inclusions. Note that only a very weak correlation might 208 exist for only the Santiago inclusions. (c) Ba/Th ratios versus Pb/206Pb ratios for Fernandina and Santiago inclusions. Notice again the lack of trend between the Ba/Th ratios and isotopic ratios. Included is the plagioclase-rich gabbro composition of protolith ! 60! gabbro Hp from Hart et al., (1999) with the isotopic composition of average DMM (Stracke et al., 2005). Figure 9. 208 Pb/206Pb versus 207 Pb/206Pb ratios for individual melt inclusions from Fernandina Island with calculated possible trends for the isotopic evolution of ancient plagioclase cumulates. The field for MORB isotopes represents a compilation for the Pacific MORB (Stracke et al., 2005 and references therein). Dashed lines in the field separates the Pacific compilation into fields of EMORB, NMORB, and DMORB compositions. Dotted lines 208 show the back-calculation of the present-day Pb/206Pb and 207 Pb/206Pb for the selected values of present-day Pacific MORB field (black square), to an initial isotopic value at 0.5 Ga (a, c, e) and 1.0 Ga (b, d, f) (white square), assumed to be the possible age of formation of the recycled gabbro. Starting Pacific MORB compositions (black square) for this calculation are separated between the panels into EMORB (a, b), NMORB (c, d) and DMORB (e, f). Solid lines show the evolution of the gabbro to its present-day isotopic composition for several different gabbro compositions. The gabbros chosen represent the extreme compositions for (U, Th)/Pb ratios from the compositions reported in Hart et., al., (1999). Also included is the average gabbro composition of the Gabal Gerf ophiolite (Zimmer et al., 1995). Individual gabbro compositions are indicated by diamonds of different colors. In our isotopic calculation, for the back calculation to the isotopic composition of the gabbro at the time of formation, we tested all of the present-day isotopic values of Pacific MORB marked by the field, but have only included the results 207 of calculation using the extreme minimum value of Pb/206Pb and 208 Pb/206Pb ratios for ! 61! 207 the field (EMORB), the extreme maximum Pb/206Pb and 208 Pb/206Pb of the field (DMORB) and an average NMORB isotopic ratio, as these represent the boundaries of the calculation. Notice the results of all 6 models are well outside the field established by the Fernandina inclusions. Values of inputs can be found in Table 2. Figure 10. Comparison of Fernandina melt inclusion compositions with trends produced by mixing of a typical Fernandina basalt with lavas generated by partial melting of a mantle plume 208 containing an ancient (~0.5-1Ga) recycled plagioclase-rich cumulate. (a-b) Pb/206Pb 207 versus Pb/206Pb ratios. Large open circles are the end-members from the mantle tetrahedron projected onto the 208Pb/206Pb -207Pb/206Pb plane (Hart, 1988, Hart et al., 1992, Hauri et al., 1994, Workman et al., 2004, Stracke et al., 2005). (c-d) Sr/Sr* versus 207 Pb/206Pb ratios. Results of critical melting model using DMM (black open symbols) and PM compositions (black filled symbols) hybridized with ancient plagioclase cumulates (now eclogite) formed at 0.5 Ga (triangles) and 1 Ga (circles) (results presented in Table 4). Ancient cumulate is generated using the calculations described in the text. The parameters of the calculation are, for the starting composition, the extreme 207Pb/206Pb and 208 Pb/206Pb of the EMORB Pacific samples and the two gabbro compositions that resulted in the isotope ratios closest to Fernandina: (a, c) ancient cumulate composition calculated using the Gabal Gerf average gabbro composition GG (Zimmer et al., 1995), (b, d) ancient cumulate using the Os strip gabbro composition (Hart et al., 1999). The model trends represent mixing of 10% aggregated critical melt of the hybrid mantle (Ag_EM_DM or Ag_EM_PM) with normal Fernandina melts in 10% increments (see ! 62! text for explanation of the model). End members of the hybrid mantle melt used for mixing (results of the critical melting model) are indicated by a larger symbol while smaller symbols denote 10% mixing steps. Note that the Pb isotope ratios and the Sr/Sr* cannot be simultaneously reproduced by this model. Figure 11. Comparison of Fernandina melt inclusion compositions with trends produced by diffusively reacting a 12% aggregated, fractional melt generated from a DMM source (Workman and Hart, 2005) with a plagioclase cumulate created from an evolved Fernandina melt (AHA31A Geist et al., 2006) at 1200°C (see text for the description of the model and Table 5 for the model parameters). (a) 208Pb/206Pb versus 207Pb/206Pb ratios. 207 (b) Sr/Sr* versus Pb/206Pb ratios. (c) Sr/Sr* versus Nb/La ratios. Large purple circle indicates the composition of the percolating melt. Small purple circles represent the composition of the melt as diffusive equilibration with the plagioclase-rich cumulate progresses within a range of 1 to 400 years indicated by the numbers next to the small circles. Grey dashed lines represent mixing between the normal Fernandina melt (D25C- 2-19B) and the percolating melt that reacted over 20 and 400 years; the grey diamonds denote 10% mixing steps. ! 63! FIGURES Figure 1-1 Figure 1 Galapagos Archipelago Pinta 0°30’N Roca Redonda Marchena Genovesa Ecuador Vol Wolf Vol 0° Santiago Darwin Vol Fernandina Subaerial Hornito Submarine Alcedo Vol 0°30’S Dredge Sierra Negra Vol Santa Cruz Isabela Cerro Azul Vol San Cristobal 1°S Floreana 92°W 91°W 90°W 89°W 20°N 0°N 20°S 120°W 90°W 60°W ! 64! Figure 1-2 2.2 EMI 2.1 Santiago EMII DMM Darwin/Wolf Is Genovesa 208Pb/206Pb Pinta Wolf Vol 2.0 Floreana -Darwin Vol -Santa Cruz -San Cristobal FOZO -Marchena Fernandina 1.9 Sierra Negra -Cerro Azul -Ecuador HIMU -Alcedo -Roca Rendonda 1.8 0.70 0.75 0.80 0.85 0.90 207Pb/206Pb ! 65! Figure 1-3 Ghost Plagioclase D25C-2-29 150 D25C-2-46 150 D25C-3-55 150 Normal Fernandina D25C-2-9 150 D25C-3-49 150 D25C-3-41 150 Depleted Fernandina D25C-2-24 150 ! 66! Figure 1-3 continued Normal Santiago STG06-29-7 150 STG06-29-17 150 STG06-29-20 150 Enriched Santiago Ultra-depleted Santiago STG06-29-10 150 STG06-29-19 150 ! 67! Figure 1-4 (a) Fernandina 0.1 wt% H2O (b) Santiago 0.1 wt% H2O 0.4 wt% H2O 0.4 wt% H2O 17 18 Al2O3 (wt%) 16 17 ~An50 Al2O3 (wt%) 15 16 14 15 13 14 15 13 14 13 12 FeO* (wt%) FeO* (wt%) 12 11 11 10 10 9 9 ~An50 8 6 7 8 9 7 8 9 10 11 MgO (wt%) MgO (wt%) Fernandina normal composition Santiago normal composition Fernandina ghost plagioclase Santiago depleted composition Fernandina depleted composition Santiago enriched composition Fernandina host glass Santiago host glass Fernandina glass Santiago subaerial lavas ! 68! Figure 1-5 100 (a) Fernandina Sample/Primtive Mantle 10 1 Fernandina normal composition 0.1 Fernandina ghost plagioclase Fernandina depleted composition Plagioclase-rich gabbro 0.01 (b) Santiago 10 Sample/Primtive Mantle 1 Santiago normal composition Santiago depleted composition Santiago enriched composition Plagioclase-rich gabbro 0.1 Ba U K Ce Pr Nd Zr Sm Ti Dy Er Lu Rb Th Nb La Pb Sr P Hf Eu Gd Y Yb ! 69! Figure 1-6 11 11 (a) (b) 10 10 Fernandina normal composition 9 2 9 Fernandina ghost plagioclase Sr/Sr* 8 8 Fernandina depleted composition Santiago normal composition 7 1 7 Santiago depleted composition Santiago enriched composition 6 6 Sr/Sr* Plagioclase-rich gabbro 5 0 13 14 15 16 17 18 19 5 4 Al2O3 (wt%) 4 3 3 2 2 1 1 0 0 13 14 15 16 17 18 19 20 21 22 0.0 0.5 1.0 1.5 2.0 Al2O3 (wt%) Nb/La ! 70! Figure 1-7 2.10 Fernandina normal composition 2.09 Genovesa Fernandina ghost plagioclase Fernandina depleted composition 2.08 Santiago normal composition EMII DMM Santiago enriched composition 2.07 2.06 208Pb/206Pb EMI 2.05 Pinta 2.2 Santiago g Floreana EMII 2.04 Avg. Santiago inclusion 208Pb/206Pb 2.1 Pinta Santiago DMM 2.03 Floreana Avg. Fernandina inclusion 2.0 Genovesa 2.02 FOZO AHA D25C (glass) 1.9 2.01 HIMU Fernandina Fernandina 1.8 2.00 0.70 0.75 0.80 0.85 0.90 0.95 0.79 0.80 0.81 0.82 0.83 0.84 0.85 0.86 0.87 207Pb/206Pb 207Pb/206Pb ! 71! Figure 1-8 12 11 (a) 2.5 10 Fernandina normal composition Fernandina ghost plagioclase 2.0 9 Santiago normal composition 8 Santiago enriched composition Plagioclase-rich gabbro 1.5 7 Sr/Sr* Sr/Sr* 6 5 1.0 4 3 0.5 2 1 0.0 0.81 0.82 0.83 0.84 0.85 0 0.805 0.815 0.825 0.835 0.845 0.855 0.865 207Pb/206Pb 207Pb/206Pb 19 350 (b) (c) 18 300 250 17 Al2O3 (wt%) 200 Ba/Th 16 150 15 100 14 50 13 0 0.805 0.815 0.825 0.835 0.845 0.855 0.865 2.00 2.01 2.02 2.03 2.04 2.05 2.06 2.07 207Pb/206Pb 208Pb/206Pb ! 72! Figure 1-9 2.25 2.25 Ps (lowest U/Pb) (a) Fernandina normal composition (b) Qs (lowest (Th/Pb) Fernandina ghost plagioclase Os (highest (U,Th)/Pb) 2.20 2.20 Hp (highest (U,Th)/Pb of protoliths) GG (average ophiolite gabbro composition) Present Day MORB composition Gabbro composition at time of formation 2.15 208Pb/206Pb 2.15 2.10 2.10 EMORB EMORB 2.05 2.05 0.5 Ga 1 Ga 2.00 2.00 0.80 0.82 0.84 0.86 0.88 0.90 0.92 0.94 0.96 0.98 0.80 0.82 0.84 0.86 0.88 0.90 0.92 0.94 0.96 0.98 2.25 2.25 (c) (d) 2.20 2.20 208Pb/206Pb 2.15 2.15 NMORB NMORB 2.10 2.10 2.05 2.05 0.5 Ga 1 Ga 2.00 2.00 0.80 0.82 0.84 0.86 0.88 0.90 0.92 0.94 0.96 0.98 0.80 0.82 0.84 0.86 0.88 0.90 0.92 0.94 0.96 0.98 2.25 2.25 (e) (f) 2.20 2.20 208Pb/206Pb 2.15 2.15 DMORB DMORB 2.10 2.10 2.05 2.05 0.5 Ga 1 Ga 2.00 2.00 0.80 0.82 0.84 0.86 0.88 0.90 0.92 0.94 0.96 0.98 0.80 0.82 0.84 0.86 0.88 0.90 0.92 0.94 0.96 0.98 207Pb/206Pb 207Pb/206Pb ! 73! Figure 1-10 2.20 2.20 Fernandina normal composition EMI EMI 2.18 (a) 2.18 (b) Fernandina ghost plagioclase 2.16 2.16 2.14 2.14 208Pb/206Pb 2.12 2.12 2.10 2.10 EMII DMM EMII DMM 2.08 2.08 2.06 2.06 2.04 2.04 2.02 GG 2.02 Os 2.00 2.00 0.80 0.82 0.84 0.86 0.88 0.90 0.92 0.80 0.82 0.84 0.86 0.88 0.90 0.92 14.0 14.0 (c) (d) 12.0 12.0 10.0 10.0 8.0 8.0 Sr/Sr* 6.0 6.0 4.0 4.0 2.0 2.0 GG Os 0.0 0.0 0.80 0.82 0.84 0.86 0.88 0.90 0.92 0.80 0.82 0.84 0.86 0.88 0.90 0.92 207Pb/206Pb 207Pb/206Pb Ag_EM_DM 1.0 Ga Ag_EM_PM 1.0 Ga Ag_EM_DM 0.5 Ga Ag_EM_PM 0.5 Ga ! 74! Figure 1-11 2.09 11 (a) (b) 2.08 10 9 20 2.07 8 50 10 2.06 7 100 208Pb/206Pb 1 200 5 2.05 Sr/Sr* 6 400 2.04 5 5 4 2.03 10 3 2.02 20 1 50-400 2 2.01 1 2.00 0 0.81 0.82 0.83 0.84 0.85 0.86 0.81 0.82 0.83 0.84 0.85 0.86 207Pb/206Pb 207Pb/206Pb 11 10 (c) Fernandina normal composition 9 20 Fernandina ghost plagioclase 8 50 10 100 Diffusion model (0-400 years) 7 200 Mixture (10% increments) 5 Sr/Sr* 6 5 400 4 3 1 2 1 0 0.0 0.5 1.0 1.5 2.0 Nb/La ! 75! TABLES Table 1. Pb isotope data of Santiago and Fernandina melt inclusion and Fernandina host glass 207 Sample Pb/206Pb 2σ 208 Pb/206Pb 2σ Santiago STG06-29-02A 0.8262 0.0090 2.052 0.019 STG06-29-02B 0.8343 0.0067 2.035 0.014 STG06-29-03 0.8405 0.0065 2.051 0.013 STG06-29-7 0.8473 0.0063 2.065 0.013 STG06-29-13 0.8268 0.0065 2.034 0.013 STG06-29-16B 0.8406 0.0124 2.068 0.024 STG06-29-16C 0.8335 0.0068 2.064 0.015 STG06-29-17 0.8389 0.0062 2.058 0.012 STG06-29-20 0.8353 0.0078 2.045 0.016 STG06-29-24 0.8324 0.0077 2.046 0.015 STG06-29-25 0.8261 0.0068 2.048 0.015 STG06-29-47 0.8332 0.0072 2.054 0.014 Santiago Enriched STG06-29-10 0.8298 0.0052 2.049 0.010 Fernandina Normal D25C-2-9 0.8194 0.0034 2.032 0.007 D25C-2-19A 0.8169 0.0038 2.027 0.008 D25C-2-19B 0.8143 0.0058 2.010 0.012 D25C-2-25A 0.8150 0.0069 2.016 0.015 D25C-2-27 0.8177 0.0036 2.031 0.007 D25C-2-30 0.8164 0.0035 2.036 0.007 D25C-2-52A 0.8113 0.0044 2.031 0.009 D25C-2-54 0.8129 0.0036 2.020 0.007 D25C-2-55 0.8179 0.0053 2.034 0.011 D25C-2-59 0.8206 0.0040 2.034 0.008 D25C-2-62 0.8146 0.0037 2.031 0.008 D25C-3-07 0.8228 0.0043 2.042 0.009 D25C-3-10 0.8240 0.0041 2.037 0.008 D25C-3-12 0.8213 0.0039 2.038 0.008 D25C-3-14 0.8192 0.0034 2.041 0.007 D25C-3-31 0.8193 0.0038 2.030 0.008 D25C-3-41A 0.8168 0.0038 2.024 0.008 D25C-3-42 0.8127 0.0039 2.028 0.009 D25C-3-43 0.8209 0.0043 2.042 0.008 D25C-3-49 0.8166 0.0047 2.026 0.010 D25C-3-52 0.8143 0.0043 2.027 0.009 ! 76! Table 1 (continued) 207 Sample Pb/206Pb 2σ 208 Pb/206Pb 2σ Fernandina Ghost Plagioclase D25C-2-29 0.8211 0.0086 2.028 0.017 D25C-2-46 0.8284 0.0169 2.040 0.033 D25C-3-34 0.8181 0.0076 2.032 0.016 D25C-3-55 0.8149 0.0067 2.037 0.015 Fernandina Depleted D25C-2-24 0.8192 0.0135 2.050 0.028 D25C-2-57 0.8120 0.0153 2.052 0.034 Host glass AHA D25C glass 0.8155 2.031 Pb isotopic ratios and error are 2σ standard error for Santiago melt inclusions and Fernandina melt inclusions. Fernandina inclusions separated into normal, ghost plagioclase, and depleted based on the differing trace element signatures of the inclusions. The Santiago inclusion with the trace element enriched composition is also separated from the normal composition Santiago inclusions. ! 77! Table 2. Inputs for Recycled Gabbro Isotopic Evolution Trace Elements‡ E-DMM* N-DMM** D-DMM** U 0.023 0.0032 0.0018 Th 0.086 0.0079 0.004 Pb 0.075 0.018 0.014 U/Pb 0.3067 0.1778 0.1286 Th/Pb 1.1467 0.4389 0.2857 Isotopes EMORB° NMORB° DMORB° 206 Pb/204Pb 18.911 18.287 17.462 207 Pb/204Pb 15.514 15.613 15.341 208 Pb/204Pb 38.319 38.174 37.025 207 Pb/206Pb 0.8204 0.8538 0.8785 208 Pb/206Pb 2.0263 2.0875 2.1203 Gabbro Trace Elements† Ps°° Qs°° Os°° Hp°° GGˆ U 0.003 0.008 0.043 0.010 0.026 Th 0.013 0.009 0.118 0.030 0.043 Pb 0.163 0.294 0.244 0.160 0.308 U/Pb 0.0184 0.0279 0.1762 0.0625 0.0831 Th/Pb 0.0798 0.0316 0.4836 0.1875 0.1404 Time (Ga) 0.5 1 ‡ In our isotopic calculation, for the back calculation to the isotopic composition of the gabbro at the time of formation, we tested all of the present-day isotopic values of Pacific MORB marked by the field in Fig. 9. Values below 0.8386 we considered EMORB and used trace element values listed for E-DMM for the calculation (bound is equal to the average Pacific EMORB value presented in Donnelly et. al., 2004). Values above 0.8766 we considered DMORB and used trace element values listed for D-DMM for the calculation (Workman and Hart, 2005). Fig. 9 shows the results of calculation using the extreme minimum 207 Pb/206Pb and 208Pb/206Pb ratios for the field (EMORB), the extreme maximum 207Pb/206Pb and 208Pb/206Pb of the field (DMORB) and an average NMORB isotopic ratio. * Donnelly et al., 2004 ** Workman and Hart, 2005 ° Stracke et al., 2005 °° Hart 1999- Hp highest U/Pb ratio protolith, Ps, lowest U/Pb ratio strip sample, Qs, lowest Th/Pb ratio strip sample, Os, highest U/Pb and Th/Pb ratio strip sample ˆ Zimmer, 1995 † at time of formation ! 78! Table 3. Partition coefficients, mineral modes and parameters used in critical melting model Kd solid-melt* Element Ol Opx Cpx Gt Co Ba 0.000005 0.000006 0.0003 0.00007 0 Th 0.000007 0.00002 0.0021 0.0021 0 Nb 0.00005 0.003 0.0089 0.011 0 K 0.00002 0.0001 0.001 0.01 0 La 0.0002 0.0031 0.054 0.0007 0 Sr 0.00004 0.0007 0.091 0.0007 0 Ce 0.00007 0.0021 0.086 0.0026 0 Nd 0.0003 0.0023 0.19 0.027 0 Zr 0.001 0.012 0.26 0.2 0 Sm 0.0009 0.0037 0.27 0.22 0 Eu 0.0005 0.009 0.43 0.61 0 Ti 0.015 0.086 0.4 0.6 0 Dy 0.0027 0.011 0.44 2 0 Y 0.0082 0.015 0.47 2 0 Er 0.0109 0.021 0.39 3.3 0 Yb 0.024 0.038 0.43 6.4 0 Pb 0.0003 0.0014 0.0075 0.0003 0 Eclogite Melting! ! ! ! ! Start Mode a 0 ! 0 ! 0.4 ! 0.5 ! 0.1 Melting 0 0 0.6 0.1 0.3 Modeb ! α** ! 0.08 ! ! ! ! F(instant)‡ 0.09 ! ! ! ! ! ! ! ! !Lherzolite Melting ! c ! ! ! ! Reaction -1.1 ! 0.5 ! 0.8 ! -0.2 ! 0 Start Mode 0.53 0.14 0.265 0.065 0 Melting 0.1 -0.3 1 0.2 0 Mode ! α** ! 0.02 ! ! ! ! F(aggre)¥ 0.1 !!! !!! !!! !!! * Halliday et al. (1995) except for quartz (coesite) where partition coefficients were assumed to be zero. a Estimated using the thermodynamic model of Sobolev and Babeyko (1994). b Taken from the experiments of Yaxley and Green (1998). ** Porosity of the melting model. Represents the amount of melt retained by the solid matrix to prevent compaction ! 79! ‡ Amount of instantaneous, eclogite melt extracted that hybridizes the overlying mantle c The reaction coefficients represent the factors for each mineral that are multiplied by the fraction of eclogite melt to determine the amount those minerals are added or subtracted from the original mantle peridotite to obtain the mineral proportions of the hybridized mantle (Yaxley & Green, 1998, Sobolev et al., 2000, Sobolev and Babeyko, 1994). The model results presented in Fig. 10 use 0.1 as the fraction of eclogite melt to represent a 0.1:0.9 proportion of eclogite melt to peridotite. Quantitative modeling of the phase change that occurs during this reaction can be found in Sobolev et al. (2000). The starting mode for the lherzolite melting is taken from Hofmann (1988). This starting mode will change depending on the amount of melt created during eclogite melting. The melting mode was taken from Sobolev et al. (2000). ¥ Amount of aggregate melt extracted from the hybrid mantle, hypothesized by Sobolev et al. (2000) to ultimately form the ghost plagioclase signature. ! ! ! 80! Table 4a. Starting compositions and results of critical melting model, Gabbro Os. Starting Compositions Modeling Results 10/90 10/90 10% 10% Element Gabbro: Os 9% EM DM PM EM_DM EM_PM Ag_EM_DM Ag_EM_PM Ba 2.46 26.70 0.56 6.60 3.18 8.61 38.23 103.63 Th 0.12 1.26 0.01 0.08 0.13 0.20 1.60 2.38 Nb 0.76 7.58 0.15 0.66 0.89 1.35 10.60 16.05 K 500.04 5148.60 49.81 240.00 559.70 730.90 6725.00 8781.80 La 1.19 10.76 0.19 0.65 1.25 1.66 13.78 18.31 Sr 155.64 1264.10 7.66 19.90 133.30 144.30 1365.10 1477.90 Ce 4.50 36.78 0.55 1.68 4.17 5.19 43.09 53.54 Nd 3.51 20.67 0.58 1.25 2.59 3.19 21.13 26.04 Zr 19.00 68.20 5.08 10.50 11.39 16.27 76.79 109.65 Sm 1.23 4.21 0.24 0.41 0.64 0.79 4.24 5.25 Eu 0.53 0.98 0.10 0.15 0.18 0.24 0.89 1.14 Ti 1865.53 3565.70 716.30 1205.00 1001.20 1441.10 4547.80 6545.60 Dy 0.00 0.82 0.51 0.67 0.54 0.69 2.10 2.69 Y 12.70 10.29 3.33 4.30 4.02 4.90 15.07 18.34 Er 0.00 0.33 0.35 0.44 0.35 0.43 1.20 1.48 Yb 1.28 0.38 0.37 0.44 0.37 0.43 0.90 1.06 Pb 0.24 2.58 0.02 0.15 0.27 0.39 3.27 4.69 0.5 Ga 0.5 Ga 0.5 Ga 207 206 Pb/ Pb 0.8479 0.8626 0.8147 0.8488 0.8362 208 Pb/206Pb 2.0405 2.0687 2.0291 2.0422 2.0365 1 Ga 1 Ga 1 Ga 207 Pb/206Pb 0.8788 0.8778 0.8556 208 Pb/206Pb 2.0606 2.0611 2.0491 ! 81! Table 4b. Starting compositions and results of critical melting model, Gabbro GG. Starting Compositions Modeling Results 10/90 10/90 10% 10% Element Gabbro: GG 9% EM DM PM EM_DM EM_PM Ag_EM_DM Ag_EM_PM Ba 47.44 514.86 0.56 6.60 51.99 57.43 625.81 691.21 Th 0.04 0.46 0.01 0.08 0.05 0.12 0.64 1.42 Nb 0.73 7.26 0.15 0.66 0.86 1.32 10.22 15.67 K 1577.19 16239.30 49.81 240.00 1668.80 1839.90 20051.40 22108.10 La 1.33 12.06 0.19 0.65 1.38 1.79 15.21 19.74 Sr 561.77 4562.80 7.66 19.90 463.20 474.20 4742.90 4855.70 Ce 3.81 31.11 0.55 1.68 3.61 4.62 37.23 47.68 Nd 2.56 15.07 0.58 1.25 2.03 2.63 16.56 21.48 Zr 15.30 54.92 5.08 10.50 10.07 14.94 67.84 100.70 Sm 0.72 2.46 0.24 0.41 0.46 0.61 3.08 4.08 Eu 0.52 0.96 0.10 0.15 0.18 0.23 0.88 1.13 Ti 5214.78 9967.30 716.30 1205.00 1641.40 2081.20 7455.50 9453.30 Dy 0.95 0.78 0.51 0.67 0.53 0.68 2.08 2.68 Y 6.48 5.25 3.33 4.30 3.52 4.40 13.18 16.46 Er 0.53 0.29 0.35 0.44 0.34 0.42 1.18 1.47 Yb 0.45 0.13 0.37 0.44 0.34 0.41 0.84 1.00 Pb 0.31 3.25 0.02 0.15 0.34 0.46 4.08 5.49 0.5 Ga 0.5 Ga 0.5 Ga 207 Pb/206Pb 0.8685 0.8626 0.8147 0.8682 0.852 208 Pb/206Pb 2.0624 2.0687 2.0291 2.0627 2.0521 1 Ga 1 Ga 1 Ga 207 Pb/206Pb 0.9258 0.9225 0.8899 208 Pb/206Pb 2.1124 2.1101 2.0855 Gabbro- starting plagioclase cumulate composition used for eclogite melting (Os-Hart et al., 1999, GG- Zimmer et al, 1995), 9%EM- is the result of 9% instantaneous melting of eclogite, DM- depleted mantle composition (Workman & Hart, 2005), PM- primitive mantle composition (Mcdonough & Sun, 1995), ! 82! 10/90 EM_DM- composition of 10% eclogite melt, 90% DM forming a hybrid mantle 10/90 EM_PM- composition of 10% eclogite melt, 90% PM forming a hybrid mantle 10% Ag_EM_DM- results of 10% aggregate melting of hybrid eclogite melt/DM mantle 10% Ag_EM_PM- results of 10% aggregate melting of hybrid eclogite melt/PM mantle Also shown are the resulting Pb isotope ratios of the hybrid mantle for plagioclase cumulates formed at 0.5 and 1 Ga. The Pb isotopes of DM are from Stracke et al., (2005), while the Pb isotopes of the PM mantle are taken to be equivalent to typical Fernandina basalt (Saal et al., 2007). Results of the melting of the hybrid mantle (both DM and PM) shown in figure 10. ! ! ! ! ! ! ! ! ! ! ! ! ! ! 83! Table 5. Diffusion Model Kd (plag- Dplag Plagioclase Initial a 2 b melt) (1200°C,m /sec) meltc meltd PM normalized output data* 1 yr 5 yr 10 yr 20 yr 50 yr 100 yr 200 yr 400 yr Ba 0.230 8.09E-19 271 4.69 0.75 0.88 1.06 1.40 2.41 4.03 7.04 9.79 Th 0.050 3.10E-21 2.88 0.07 0.83 0.83 0.83 0.83 0.83 0.84 0.84 0.84 U 0.024 3.10E-21 0.98 0.03 1.31 1.31 1.31 1.31 1.31 1.32 1.32 1.32 Nb 0.030 3.10E-21 47.9 1.24 1.88 1.88 1.88 1.88 1.88 1.88 1.89 1.89 Ta 0.030 3.10E-21 3.42 0.08 2.16 2.16 2.16 2.16 2.16 2.17 2.17 2.17 K 0.171 4.38E-16 14693.6 408.33 18.79 47.21 53.76 54.84 54.86 54.86 54.86 54.86 La 0.106 3.87E-19 35.14 1.60 2.48 2.52 2.57 2.67 2.96 3.44 4.39 5.32 Ce 0.081 3.87E-19 79.68 4.57 2.73 2.76 2.79 2.86 3.05 3.38 4.01 4.64 Pb 0.511 1.41E-17 2.48 0.15 1.84 4.76 7.57 11.25 15.11 15.89 15.94 15.94 Pr 0.104 10.8 0.88 3.27 3.29 3.32 3.37 3.52 3.77 4.27 4.76 Sr 1.813 3.37E-17 561 63.48 9.34 21.99 26.50 27.83 27.91 27.91 27.91 27.91 Nd 0.084 2.92E-19 44.96 4.76 3.81 3.83 3.85 3.88 3.99 4.17 4.53 4.88 Zr 0.004 3.10E-21 350 41.47 3.95 3.95 3.95 3.95 3.95 3.95 3.95 3.95 Hf 0.015 3.10E-21 7.55 1.27 4.50 4.50 4.50 4.50 4.50 4.50 4.50 4.50 Sm 0.056 2.92E-19 10.27 1.88 4.62 4.63 4.63 4.65 4.70 4.77 4.93 5.08 Eu 0.534 2.92E-19 3.19 0.74 5.43 7.51 9.29 11.21 12.64 13.01 13.53 14.00 Ti 0.076 3.10E-21 24932.9 5310.42 4.41 4.41 4.41 4.41 4.41 4.41 4.41 4.41 Gd 0.017 9.65 2.68 5.04 5.04 5.05 5.05 5.08 5.12 5.19 5.27 Tb 0.038 1.43 0.49 5.25 5.25 5.25 5.25 5.27 5.29 5.32 5.36 Dy 0.019 3.24E-19 7.78 3.37 5.46 5.46 5.46 5.46 5.46 5.46 5.46 5.46 Er 0.012 3.88 2.12 5.04 5.04 5.04 5.04 5.04 5.05 5.05 5.05 Yb 0.015 5.06E-19 3.16 2.04 4.63 4.63 4.63 4.63 4.63 4.64 4.65 4.65 207 Pb/206Pb 0.8147 0.8538 0.8345 0.8208 0.8176 0.8159 0.8150 0.8148 0.8148 0.8148 208 Pb/207Pb 2.0291 2.0875 2.0587 2.0382 2.0334 2.0308 2.0294 2.0292 2.0292 2.0292 ! 84! Table 5 (continued) a (m) 0.001 %melt to solid 1% %Plag in solid 50% a Experimentally determined partition coefficients for plagioclase (An80) at 1200°C from Bedard (2006) b Diffusion coefficients from Cherniak and Watson (1994), Cherniak (1995), Cherniak (2002), Cherniak (2003), Giletti and Casserly (1994), and Giletti and Shanahan (1997) for plagioclase at 1200°C. For diffusion coefficients not determined experimentally we used estimates based on cation size and charge (Behrens et al., 1990, Saal and Van Orman, 2004). We assumed diffusion coefficients for Nb, Ta, Zr, Hf, and Ti are similar to those of Th and U. We assumed Eu is 50% Eu+2 and that the diffusion coefficient for Eu+2 is similar to that of Sr c Geist et al., (2006) composition AHA31A with isotopic composition of average Fernandina whole rock. d 12% aggregated fractional melt of DMM composition from Workman and Hart (2005) with the isotopic values of average pacific DMM from Stracke et al. (2005) * McDonough and Sun (1995) ! ! ! 85! CHAPTER 2 Determining the volatile budget of the Galapagos Plume: separating deep and shallow signatures. M. E. Peterson1,*, A. E. Saal1, M. D. Kurz2, E. H. Hauri3, J. S. Blusztajn2, K. S. Harpp4, R. Werner5, D. J. Geist6, 1 Department of Earth, Environmental and Planetary Sciences, Brown University, Providence, Rhode Island, 02912, United States 2 Department of Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, 02543, United States 3 Department of Terrestrial Magnetism, Carnegie Institution of Washington Washington, District of Columbia, 20015, United States 4 Geology Department, Colgate University, Hamilton, New York, 13346, USA 5 GEOMAR Helmholtz Centre for Ocean Research Kiel, Wischoftrasse 1-3 24148 Kiel, Germany 6 Department of Geological Sciences, University of Idaho Moscow, Idaho, 83844, USA To be submitted to Journal of Petrology 4/27 ! 86! ABSTRACT We report new major, trace and volatile element contents (H2O, CO2, F, S, Cl), and new Sr, Nd, Pb and He isotopes on 118 submarine glass chips from the Galapagos Archipelago to evaluate the volatile budget of the Galapagos mantle plume. Major, trace and isotopic measurements identify two mixing arrays: Array 1 characterized by submarine glasses from near Fernandina (HHe: high 3He/4He component) and Sierra Negra/Cerro Azul volcanoes (ITE: incompatible trace element enriched component), and Array 2 defined by dredged samples from near Pinta/Wolf-Darwin lineament (WD: low 3 He/4He ratios and high D7/4 D8/4 and Th/La ratios) and Genovesa Island (ITD: MORB- like component). Degassing has significantly lowered the CO2 content of the lavas and had some effect on the H2O and S concentrations, but no recognizable impact on the F and Cl contents. The submarine glasses range from being saturated with a sulfide liquid to sulfide undersaturated at the estimated pressures for crystal fractionation (2 Kbars). Assimilation of hydrothermally altered material, assessed using Cl/K ratios and the magnitude of the Sr anomaly (Sr/Sr*>1), affected the volatile contents of a significant number of samples in both arrays. The chemistry of the samples in array 2 suggests shallow mixing occurred, after some degassing and contamination took place in the end- members. Ratios between volatile and refractory elements with similar incompatibilities are used to estimate the mantle volatile budget of the Galapagos mantle plume. The Fernandina end-member shows a F/Nd ratio (20.1 ± 0.8) and H2O/Ce ratio (202 ± 26) similar to the value predicted for FOZO. The Sierra Negra/Cerro Azul end-member ! 87! shows a high F/Nd ratio (24.9 ± 0.7), and a slightly lower H2O/Ce ratio (174 ± 24), consistent with values previously considered for other mantle plumes with enriched mantle components. S/Dy and Cl/K ratios for these two end-members are very similar (HHe: S/Dy = 218 ± 21, Cl/K = 0.057 ± 0.009 ITE: S/Dy = 237 ± 20, Cl/K = 0.049 ± 0.010) consistent with ratios previously measured in lavas from other ocean islands. The Pinta-Wolf-Darwin end-member has a F/Nd (23.5 ± 0.4), Cl/K (0.066 ± 0.006) and S/Dy ratio (255 ± 40) similar to those of Sierra Negra/Cerro Azul but with a significantly higher H2O/Ce ratio (230 ± 35). The volatile content of the ITD glasses of the Genovesa end-member have been significantly affected by secondary processes with only few samples showing volatile/refractory element ratios close to those previously reported for MORB. C/3He ratios in Galapagos glasses from the HHe group are, on average, an order of magnitude higher than predicted values for DMM and ranges to the highest values measured for oceanic basalts, indicating that the Galapagos plume contains a mantle component that is more carbonated than the depleted upper mantle. ! 88! INTRODUCTION Volatile elements (H2O, CO2, F, S, and Cl) assert a strong influence on mantle melting, magma crystallization, style of eruption, and on the viscosity and rheology of the mantle (Hirth & Kohlstedt, 1996, Gaetani & Grove, 1998, Asimow & Langmuir, 2003, Asimow et al., 2004). A significant amount of research has been devoted to characterizing the volatile content of the mantle (e.g. Michael, 1989, Michael, 1995, Dixon et al., 2002, Saal et al., 2002, Stroncik & Haase, 2004). One approach is to assess variations in volatile to refractory element ratios that share similar incompatibility during mantle melting and fractional crystallization (Michael & Cornell, 1998, Saal et al., 2002, Hauri et al., 2006, e.g Beyer et al., 2012, Dalou et al., 2012, Rosenthal et al., 2015). However, surface processes, such as degassing, sulfide saturation and, melt-crustal interaction, can significantly modify magmatic volatile/refractory element ratios in the melt (Michael, 1989, Dixon et al., 1995, Dixon, 1997, Michael & Cornell, 1998, Dixon & Clague, 2001, Saal et al., 2002, Stroncik & Haase, 2004, le Roux et al., 2006, Workman et al., 2006). Therefore, it is important to evaluate these processes before extrapolating the observed variability of these ratios in the magmas to their mantle sources. Oceanic hotspots, such as the Galapagos Archipelago, are ideal locations to study the volatile budget of the mantle because such melting anomalies contain several isotopically unique reservoirs, hypothesized to be the result of both recycled crustal and deep mantle material (e.g. Zindler & Hart, 1986, Hart et al., 1992, Hofmann, 1997, White, 2010). However, only a limited number of studies have examined the volatile budget of mantle plumes in comparison to studies that have focused on mid-ocean ridge basalts (Schilling et al., 1980, Jambon et al., 1995, Dixon & Clague, 2001, Dixon et al., 2002, ! 89! Lassiter et al., 2002, Simons et al., 2002, Stroncik & Haase, 2004, Workman et al., 2006, Koleszar et al., 2009, Cabral et al., 2014, Kendrick et al., 2014). These studies suggest that the mantle component characterized by high 3He/4He signature (sometimes called FOZO, PHEM or C and indicative of a "relatively undegassed" lower mantle; e.g. Kurz et al., 1982b, Allègre et al., 1983, Farley et al., 1992, Hart et al., 1992, Hanan & Graham, 1996), has a higher volatile content (H2O, CO2, and Cl) relative to normal MORB (N- MORB) (Dixon et al., 2002, Saal et al., 2002, Simons et al., 2002, Cartigny et al., 2008, Koleszar et al., 2009), while material with an isotopic signature associated with crustal recycling (high Sr and Pb and low Nd and Hf isotope ratios) is thought to be depleted in H2O contents while possessing variable but higher Cl contents relative to N-MORB (Dixon & Clague, 2001, Dixon et al., 2002, Saal et al., 2002, Stroncik & Haase, 2004, Workman et al., 2006, Cabral et al., 2014, Kendrick et al., 2014). The data is limited for S and F contents in plume basalts, with highly variable S/Dy ratios (ranging from less than to greater than expected N-MORB values) being explained by both degassing and sulfide fractionation. Equally variable F/Nd ratios are attributed to the presence of depleted mantle (F/Nd ~ 18), primitive mantle (F/Nd ~ 21), and variable amounts of recycled material in the source (for values higher than the primitive mantle) (Saal et al., 2002, Workman et al., 2006, Koleszar et al., 2009, Cabral et al., 2014, Kendrick et al., 2014). Isotopic studies of the Galapagos Archipelago and associated hotspot tracks have found the plume to be heterogeneous with four main isotopic components: a high 3He/4He component and 2 components (characterized by the melts of Pinta and Floreana Island in the Galapagos) with high Sr and Pb, and low Nd and Hf isotope ratios in addition to a ! 90! MORB-like end-member that is hypothesized to either be depleted upper mantle or a depleted component intrinsic to the Galapagos plume (e.g. Graham et al., 1993, White et al., 1993, Kurz & Geist, 1999, Hoernle et al., 2000, Blichert-Toft & White, 2001, Harpp & White, 2001, Werner et al., 2003, Saal et al., 2007, Koleszar et al., 2009, Kurz et al., 2009, Vidito et al., 2013). Koleszar et al. (2009) presented volatile data of melt inclusions on a sample set from Fernandina (high 3He/4He source) and Santiago islands (MORB-like end member) and found that the measured volatile/refractory element ratios were generally within the range observed in other high 3He/4He ocean island basalts (OIB) (Dixon & Clague, 2001, Simons et al., 2002, Workman et al., 2006), and MORB (Saal et al., 2002), respectively. However, the data was was limited to olivine-hosted melt inclusions picked from 1 sample collected at each location. We present new geochemical information for 118 submarine glasses collected from across the Archipelago to obtain the volatile budget of the Galapagos plume. New major, trace and volatile element concentrations (H2O, CO2, F, Cl, S) and new Sr, Nd, Pb and He isotope data provide new insights into the mantle beneath the archipelago. We first consider the effects of shallow level processes affecting the volatile contents of the basalts (crystal fractionation, degassing, sulfide saturation, and contamination with hydrothermally altered material) allowing us to identify the most pristine samples that let us evaluate the volatile variations in the different mantle sources that contribute to the Galapagos plume. Geochemical Background ! 91! The Galapagos Archipelago is located in the equatorial Pacific <200 km south of the Galapagos Spreading Center (GSC). The archipelago contains 21 emergent volcanoes that sit on a shallow submarine platform. Of these volcanoes, 16 have erupted since the Holocene, making it one of the most active volcanic regions in the world. The Galapagos Archipelago displays four main compositional components, two distinct end-members (having long-term enrichment of incompatible trace elements) occurring in the southern and northern regions, a high 3He/4He component occurring in the west, and a MORB-like end member prevalent in the center and northeast region of the archipelago (Swanson et al., 1974, Vicenzi et al., 1990, White et al., 1993, Kurz & Geist, 1999, Hoernle et al., 2000, Harpp & White, 2001, Harpp et al., 2002, Harpp et al., 2003, Saal et al., 2007, Gibson et al., 2012). The MORB-like end-member has been proposed to be either the depleted upper mantle or a depleted component intrinsic to the plume. The western component, found primarily near Fernandina Island, is characterized by intermediate 87Sr/86Sr, 143 Nd/144Nd, 206 Pb/204Pb, 207 Pb/204Pb, 208 Pb/204Pb and 176 Hf/177Hf ratios and high 3He/4He ratios reaching up to 29 Ra (3He/4He ratio normalized to the atmospheric ratios, 1.4e-6) (Graham et al., 1993, White et al., 1993, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Kurz et al., 2009) corresponding to the FOZO component of Hart et al. (1992). The northern component, illustrated by the magmas collected from Pinta Island and the Wolf-Darwin lineaments’ seamounts and islands, is intermediate to the other end-members in terms of 87Sr/86Sr, but has elevated 207 Pb/204Pb and 208Pb/204Pb at a given 206Pb/204Pb and represents the lowest 143Nd/144Nd and 176 Hf/177Hf isotopic ratios yet measured in the Galapagos (Kurz & Geist, 1999, Blichert- Toft & White, 2001, Harpp & White, 2001, Saal et al., 2007). The low Nd and Hf ! 92! isotopes and the elevated Pb isotopes that deviate from the northern hemisphere reference line are within the range of the Dupal anomaly (Hart, 1984). The one 3He/4He measurement made for Pinta (6.9 Ra; (Kurz & Geist, 1999)) is at the lower values of the range defining MORB akin to the low variation seen in the southwest Indian Ridge (e.g. Kurz & Jenkins, 1981, Kurz et al., 1982a, Graham et al., 2001, Detrick et al., 2002, Georgen et al., 2003, Colin et al., 2011, Graham et al., 2013). Samples collected from Pinta have higher Pb and Sr isotope ratios than MORB, suggesting that the northern component of the Galapagos plume is not part of the MORB source but part of the plume. The southern component is defined mainly by the magmas of Floreana Island and is characterized by the highest 87Sr/86Sr, 206Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb isotopic ratios found in the Galapagos with intermediate 3He/4He ratios (up to 14 Ra), 143 Nd/144Nd 176 isotopic ratios and unusually elevated Hf/177Hf ratios for the measured Nd isotopes (Bow & Geist, 1992, White et al., 1993, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Kurz et al., 2009, Harpp, 2014). This end-member is isotopically most similar to the HIMU mantle component, though its high 87Sr/86Sr and 176 Hf/177Hf ratios make it atypical in comparison to other HIMU magmas (Bow & Geist, 1992, Harpp & White, 2001, Harpp et al., 2014). We note that, the radiogenic isotopes of this end-member are restricted to samples collected sub-aerially from the island itself, while the submarine platform near the island has an isotopic flavor similar to Sierra Negra and Cerro Azul Volcanoes (e.g. Harpp et al., 2014). Sierra Negra and Cerro Azul Volcanoes, located on the southern half of the Isabela Island are characterized by intermediate compositions among those of Fernandina, Floreana and Pinta (White et al., 1993, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Saal et ! 93! al., 2007, Harpp, 2014). The composition of Sierra Negra and Cerro Azul, is either the product of a mixture among several end-member compositions (e.g., Fernandina and Floreana), or a unique mantle source with an intermediate isotopic signature (e.g. White et al., 1993, Harpp & White, 2001). The glasses of this study cover almost all the major geochemical domains in the Galapagos, with the exception of the Floreana component. ANALYTICAL METHODS Multiple glass chips were hand-picked from 118 dredged pillow basalts collected during dredging operations of the PLUME02 cruise in 1990 (Christie et al., 1992), AHA-Nemo- 2 cruise using the R/V Melville in 2000 (Geist et al., 2006), the DRIFT04 cruise using the R/V Revelle in 2001 (Kurz et al., 2001), and the R/V SONNE cruise SO 158 in 2001 (Werner, 2002). Locations of dredges are reported in Table 1 and shown in Figure 1. Major, trace and volatile element concentrations were obtained by replicate analyses on 2-4 glass chips picked from the same sample. The glass chips were mounted in indium and analyzed for major, trace and volatile elements. Precision is reported as 2σ standard error. Major element concentrations were determined by the Cameca SX-100 electron microprobe at the Department of Geological Sciences, Brown University. Analyses were performed using 15 kV accelerating voltage, 15 nA beam intensity, and 5 μm beam diameter for the glasses and PAP correction procedures (Pouchou & Pichoir, 1991, Devine, 1995). Primary standards used for calibration are USGS glasses BCR-2G, BIR-1G and AGV-2, Wilberforce apatite, Amelia albite (Purdue), fayalite (NMNH85276), rhodonite (NMNH104738), and synthetic samples of forsterite (FO97), rutile, and orthoclase (OR1). Smithsonian basaltic glasse VG-2B (USNM 111240) was ! 94! monitored as a secondary standard (included in Table 1). Precision for major elements is usually better than 2% for SiO2, TiO2, Al2O3, FeO, MgO, and CaO, better than 7% for K2O and Na2O, better than 10% for P2O5 and better than 20% for MnO. Trace elements analyses (Rb, Ba, Th, U, Nb, K, Pb, Sr, P, Zr, Hf, Ti, and REE) were conducted in 3 sessions. The first session was performed on the Cameca 6F Ion Probe at the Department of Terrestrial Magnetism, Carnegie Institution of Washington using techniques of Shimizu and Hart (1982). The second and third sessions were performed at the Graduate School of Oceanography, University of Rhode Island using a Thermo X-Series II quadrupole ICP-MS coupled with a New Wave UP 213 Nd-YAG laser ablation system, using a spot size of 80μm, 60% energy output and 10 Hz repeat rate. The second session was done on the glasses having depleted incompatible trace element contents previously analyzed by SIMS; we found the data to be consistent between both techniques while acquiring several new elements (Pr, Tb, Ho, Tm, Lu) using the LA-ICPMS. The third session was performed to expand the data set. Techniques and procedures used for reducing trace element data, gathered during the LA-ICPMS sessions are described in Kelley et al. (2003) using 43Ca as an internal standard for the glasses. Calibration curves were generated using eight natural-composition glasses from the USGS (BIR-1G, BCR- 2G, BHVO-2G) and the Max Planck Institute (KL2-G, ML3BG, StHs6/80-G, GOR132-G, T1-G; (Jochum et al., 2006)). Precision for trace element analyses is better than 10%. Volatile element analyses (H2O, CO2, F, S, and Cl) were performed on the Cameca 6F Ion Probe at the Department of Terrestrial Magnetism, Carnegie Institution of Washington using a Cs+ primary beam. For volatile element analyses we used the technique and standards previously described in Hauri (2002). Precision is typically ! 95! better than 9% for H2O, F, S, and Cl and better than 15% for CO2. Many isotope ratios and noble gas measurements are from the literature (Graham et al., 1993, Harpp & White, 2001, Geist et al., 2008, Kurz et al., 2009). New isotope data are presented in Table 1 (new data in table are indicated by bold font). New Sr, Nd, Pb and He isotope analyses for samples collected during the DRIFT04 and AHA-Nemo-2 cruises were conducted at the Woods Hole Oceanographic Institution. Prior to dissolution for Sr-Nd-Pb isotopic analysis, chips were leached for 1 hour in 6N HCL. Conventional ion-exchange procedures were used and the isotope ratios were measured with a Thermo- Finnigan NEPTUNE MC-ICPMS. Typical procedural blanks are <100 pg for Sr analyses, < 30 pg for Nd analyses, and 30-50 pg for Pb analyses. Internal two sigma (2σ) precision for Sr and Nd isotopic ratios is approximately 5-10 ppm and between 15 and 60 ppm for Pb. The external reproducibility, after normalization for NBS987 (87Sr/86Sr = 0.710240) and La Jolla standards (143Nd/144Nd = 0.511847), is ~25 ppm and ~15 ppm (2σ) respectively. External reproducibility for Pb isotopes ranges from 150-200 ppm (2σ). Two USGS standards (AGV-1: 18.9414, 15.6548, 38.5615 for 206Pb/204Pb, 207Pb/204Pb, and 208 Pb/204Pb respectively and BCR-1: 18.8215, 15.6356, 38.7309 for 206Pb/204Pb, 207Pb/204Pb, 208 and Pb/204Pb respectively) were used to verify the Pb procedure and yielded good agreement with literature values. Further information on isotope analytical procedures are given by Hart and Blusztajn (2006). Instrumental mass fractionation for isotopic analyses is corrected for by normalizing to 86Sr/88Sr = 0.1194, 146 Nd/144Nd = 0.7219. Pb isotope 206 analyses were normalized to Pb/204Pb = 16.9356, 206 Pb/204Pb = 15.4891, 206 Pb/204Pb = 36.7006. For He isotopic analyses, glasses are cleaned sonically once in distilled water, weak nitric acid and hydrogen peroxide, followed by a second cleaning in distilled water ! 96! and acetone. Chips were selected from cleaned glasses to be crushed in vacuum for measurement of helium abundance and isotopic composition using simultaneous collection of 3He and 4He in a fully automated mass spectrometer dedicated to helium measurements at the Woods Hole Oceanographic Institution, A subset of the powders that remained after crushing was melted in vacuum, which yielded the total helium concentration in the glasses and also allows a comparison with ion probe measurements in the glass. A description methods can be found in Kurz et al. (2004). Helium blanks measured during the course of the samples run were typically less than 5x10-11 cm3 STP. RESULTS The major elements, trace element and isotopic compositions of the glasses in this study cover almost the entire compositional range reported for the Galapagos with the exception of the Floreana component (Bow & Geist, 1992, Graham et al., 1993, White et al., 1993, Geist et al., 1998, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Saal et al., 2007, Kurz et al., 2009, Gibson et al., 2012). The majority of the samples, collected from submarine dredges, fall on 2 distinct arrays that are distinguishable from each other both in terms of geographic location and in terms of major element, trace element and isotopic variations (see below). The first array (Array 1) consists of samples with compositions ranging between those from Fernandina Island and Sierra Negra/Cerro Azul volcanoes. The array is further divided into the high 3 He/4He ratio (HHe) group and the incompatible trace element enriched (ITE) group using 3He/4He ratios above and below 22 Ra respectively. Samples of array 1 were collected exclusively from the western and southwestern edges of the Galapagos platform, ! 97! with the HHe group samples being concentrated around Fernandina Island (Fig. 1). The second array (Array 2) is comprised of samples that compositionally vary from incompatible trace element depleted samples (ITD group) similar to Genovesa Is (i.e., MORB), to incompatible trace element enriched glasses similar to samples from Pinta- Wolf-Darwin Is. (WD group) characterized by low 3He/4He ratios, high Th/La ratios and high D7/4, D8/4. ITD group samples from array 2 were mostly collected within the northeastern seamounts, with the samples of the WD group being exclusively collected in the area of the Wolf-Darwin lineaments. There are four samples collected just to the north of Wolf Volcano on Isabela Island that have unique compositions not matching those of either array; instead, these samples have geochemical characteristics similar to the lavas from Wolf Volcano. The major element compositions of the glasses in this study indicate that the bulk of the samples are tholeiitic basalt with only a small subset of samples with an alkalic composition (Macdonald & Katsura, 1964). Glasses have MgO contents ranging between 5 and 9.85 wt%. Lavas from array 1 that fall between the compositions of Fernandina and Sierra Negra/Cerro Azul Volcanoes are generally more evolved (average MgO content of 6.16 wt%) than samples from array 2 that fall between the compositions of Genovesa and Pinta-Wolf-Darwin Is. (average MgO content of 8.00 wt%) (Fig. 2). This is consistent with an inferred higher magma flux near the plume center leading to the formation of well-developed magma chambers that produce more differentiated and homogenous lavas in the western sector of the archipelago (Geist et al., 2014). In contrast, lavas from the north and northeastern sectors are generated under a lower magma flux that precludes the ! 98! formation of large magma chambers and promotes the eruption of less differentiated lavas. In order to evaluate the effects of magma differentiation in array 1 and array 2, we calculated the liquid lines of descent (LLD) using PetroLog3 for progressive crystallization of olivine, plagioclase, and clinopyroxene (Danyushevsky & Plechov, 2011). We considered the compositions and water content of a trace element enriched (Fernandina inclusion) and a depleted melt inclusion (Santiago inclusion), compositions both reported in Table 2 (Koleszar et al., 2009), because the isotope signatures of the two arrays are very different (which is discussed below). This indicates that several mantle sources are needed to produce the geochemical variability found in the glasses. Modeling was done for isobaric fractional crystallization at 2 Kbars. Pressure was chosen to be an average of the entrapment pressure of the Fernandina and Santiago melt inclusions that were found to be the least degassed. This is consistent with the average pressure of crystallization of Fernandina lavas from Geist et al. (2006) and provides a good fit to the data, though varying this pressure from 500-4000 bars does not significantly change the interpretation of the LLDs. The results of the modeling (Fig. 2) show that the major elements for arrays 1 and 2 cannot be produced by a single LLD. The same conclusion is reached if parameters such as the water content of the starting compositions are varied. Therefore modeled LLDs in addition to highly variable trace elements and isotopic signatures suggests that the glasses of the array 1 and of the array 2 are produced by melting different mantle sources followed by mixing of the melts. A notable feature of the ITD glasses in array 2 is that, at high MgO contents, these glasses have much higher Al2O3, lower TiO2 and FeO contents than is predicted by the ! 99! LLD. These samples also overlap the range of the Santiago melt inclusions affected by plagioclase cumulate assimilation (Koleszar et al., 2009, Peterson et al., 2014) suggesting that the same process affected these glasses. This is consistent with the strong correlation between Sr/Sr* (SrPM/[(PrPM+NdPM)/2]) and Al2O3 contents shown by the samples (Fig. 3). Sr/Sr* is the magnitude of the Sr anomaly on a primitive mantle (PM) normalized diagram and is a metric of plagioclase assimilation (Sr/Sr* > 1) and fractionation (Sr/Sr* < 1) (Fig. 4). The Santiago melt inclusions show progressive assimilation of plagioclase cumulates indicated by a negative correlation between Sr/Sr* and MgO contents. The ITD samples have the opposite relationship between MgO and Sr/Sr*. The highest MgO glasses in this study start out with very high Sr/Sr* ratios that then decrease with decreasing MgO contents. This indicates that the glasses are most likely not primitive, uncontaminated submarine glasses and the least evolved samples underwent significant plagioclase assimilation (causing the high Sr/Sr* of the highest MgO samples) followed by crystallization of plagioclase during subsequent differentiation (causing the positive correlation between MgO and Sr/Sr*). Array 1 between Fernandina and Sierra Negra/Cerro Azul: HHe Group and ITE Group The samples that fall on array 1 between the compositions of Fernandina and Sierra Negra/Cerro Azul are enriched in incompatible trace elements (Fig. 4) and have high Nb/Zr (0.11-0.16) and Th/La ratios (0.07-0.10) (Supplementary Fig. S1). Nb can potentially be fractionated from Zr by low degrees of melting. The presence of several alkalic composition glasses implies that there is a range in the degree of melting for the ! 100! glasses of this study. However, the alkalic samples also possess distinct isotopic signatures, which suggest they are produced from a unique mantle source. Additionally, Nb/Zr and Th/La ratios are also correlated with isotopic ratios (Supplementary Fig. S2). Therefore variations in the degree of melting cannot entirely explain the range in compositions. As a result, both Nb/Zr and Th/La ratios, in conjunction with their isotopic compositions, reveal the incompatible trace element signatures of the mantle to be long- term characteristics. The high Sm/Yb ratios (1.85-2.75) in these lavas indicate deep melt generation within the garnet-lherzolite stability field. High-Ti anomalies on a primitive mantle normalized trace element diagram are cited as a characteristic of the FOZO mantle source (e.g. Saal et al., 2007, Jackson et al., 2008) therefore the large range in Ti/Gd ratios (2480-3216) could be reflecting the presence of FOZO end-member in the Fernandina mantle source. A recent study of Peters and Day (2014), questioned the significance of the Ti anomalies as a global indicator of a high 3He/4He mantle source, citing instead a combination of unique pressure-temperature melting conditions of OIB’s and AFC processes. In the case of the Galapagos lavas, it is very difficult to explain the observed correlation (Supplemental Fig. S3) between 3He/4He and the Ti anomaly by magmatic processes. Therefore the Ti anomalies in the Fernandina samples most likely reflect a characteristic of the Fernandina mantle source. The end-member compositions of the array 1 are distinctive isotopically, with samples that possess 3He/4He ratios between 22 Ra and 28.82 Ra (HHe group) showing, on average, lower Pb and Sr at similar Nd isotope ratios than samples with 3He/4He ratios less than 22 Ra (ITE group), consistent with previously measured isotopes for Fernandina ! 101! and Sierra Negra/Cerro Azul whole rock samples (Fig. 5) (White et al., 1993, Kurz & Geist, 1999, Saal et al., 2007, Kurz et al., 2009). Array 2 between Genovesa (MORB-like) and Pinta/Wolf, Darwin Islands: ITD Group and WD Group The samples that fall on array 2 range between the compositions similar to Genovesa and Pinta-Wolf-Darwin Is., with incompatible trace element contents that span from depleted (ITD group), equivalent to those observed in MORB, to enriched compositions (WD group being slightly different from those measured in the ITE group) (Fig. 4, 5, S1 and S2). The maximum Nb/Zr (0.01-0.15) and Ti/Gd (1553-2648) ratios of array 2 overlap values seen in the HHe and ITE groups (Fig. S1). Sm/Yb ratios (0.73-1.77), however, are almost universally lower, which is most likely indicative of melting in the shallower spinel-lherzolite stability zone. The range in Th/La ratios (0.02-0.127) is the most distinctive characteristic of samples from array 2, reaching significantly higher (WD group) and lower (ITD group) than those of the HHe and ITE groups. The high Th/La ratio end-member is a characteristic of lavas from Pinta (Saal et al., 2007), while the low Th/La ratio component is typical of MORB. Samples from array 2 have 3He/4He ratios (6.87-8.81 Ra) within the range for MORB, with the lowest value found in samples from the WD group (Kurz & Jenkins, 1981, Kurz et al., 1982b, Graham et al., 1993, Kurz & Geist, 1999, Graham et al., 2001, Detrick et al., 2002, Georgen et al., 2003, Kurz et al., 2009, Colin et al., 2011, Graham et al., 2013). Glasses from the ITD group have Sr, Nd and Pb similar to those measured in MORB and Genovesa lavas (Fig. 5) (White et al., 1993). Glasses from the WD group ! 102! have Sr, Nd and Pb isotope ratios that overlap with those measured in whole rock samples from Pinta Island and the Wolf-Darwin lineaments. The Pb isotopes are the most 207 distinctive characteristic in samples from WD group with higher Pb/204Pb = 15.561- 208 15.620, and Pb/204Pb = 38.532-39.282 for a given 206 Pb/204Pb range = 18.852-19.366 (high D7/4 and D8/4 as defined by Hart (1984)) (Cullen et al., 1989a, White et al., 1993, Harpp & White, 2001, Saal et al., 2007). High Th/La ratios and D7/4 and D8/4 values are characteristics commonly been attributed to the presence of an older recycled crustal component in the source of OIB (e.g. Hofmann, 1997). A greater crustal component in the WD group over the lavas from the other groups is consistent with its low 3He/4He, though, more data is required to confirm that the WD group has higher Δ7/4, D8/4 and lower 3He/4He than MORB. These characteristics are also consistent with plumes that contain an Enriched Mantle (EM) component such as the Samoa, Pitcairn and Society plumes (e.g. Stracke et al., 2003, Workman et al., 2006, Kendrick et al., 2014). The range in trace element and isotopic compositions of array 2 completely overlaps with the composition demarcated by samples from the GSC (Figures 5, S1 and S2). The compositions of submarine samples collected just north of Isabela island do not match the compositional range defined by either array, but they are consistent with those measured in whole rock lavas from Wolf Volcano. The glasses collected from north of Isabela have Th/La ratios and Sr, Nd and Pb isotopic compositions similar to MORB, but with higher Nb/Zr, Sm/Yb and Ti/Gd ratios consistent with a low degree of melting in the garnet-lherzolite stability field of a depleted mantle as previously proposed for Wolf Volcano lavas (Geist et al., 2005, Saal et al., 2007). ! 103! The 2 distinct arrays based on major and trace element contents and Sr, Nd, Pb and He isotope data implies that the glasses of this study define four end-members. The end-members of array 1 correspond to the trace element and isotopic compositions of Fernandina Island (HHe group) and Sierra Negra/Cerro Azul volcanoes (ITE group), while the end-members of array 2 correspond to Pinta-Wolf-Darwin Is. (WD group) and Genovesa Is. (ITD group). Thus the data will define volatile budget variations of these mantle source compositions contributing to the Galapagos plume. Volatiles There is a large range in volatile concentrations exhibited by the Galapagos submarine glasses, with H2O concentrations varying from 0.10 to 1.23 wt%, CO2 from 10.7 to 193.7 ppm, F from 61.4 to 840.7 ppm, S from 715 to 2185 ppm and Cl from 3.80 to 875 ppm. These concentrations are within the range found at other OIBs (Dixon & Clague, 2001, Dixon et al., 2002, Hauri, 2002, Simons et al., 2002, Workman et al., 2006, Cabral et al., 2014, Kendrick et al., 2014). Concentrations of H2O, F, Cl, and S generally increase with decreasing MgO wt% as is predicted by calculated LLDs from primitive melt inclusions representing the HHe and the ITD end member component (Koleszar et al., 2009) (Fig. 6; parameters such as partition coefficients for volatile elements for LLD models included in Table 2), though the data do to deviate from the LLD models. This deviation is most likely due to a combination of shallow level processes and source heterogeneity. To overcome the effect of crystal fractionation we consider ratios of volatile to refractory elements with similar incompatibility during melting and crystallization (Nb, Ce, Nd, K, and Dy with CO2, H2O, F, Cl, and S, respectively, Fig. 6) (Schilling et al., 1980, Michael, ! 104! 1995, Michael & Cornell, 1998, Danyushevsky et al., 2000, Saal et al., 2002, Hauri et al., 2006, Workman et al., 2006). Once we account for the effect of shallow level processes such as degassing, sulfide saturation and assimilation of hydrothermally altered material, the volatile to refractory element ratios provide a straightforward approach to determine the volatile budgets of the different mantle components forming the Galapagos plume. DISCUSSION Shallow level process Degassing The partitioning of H2O, CO2, S, F, Cl from melt to a gas phase is controlled by the melt and gas composition, ƒO2, solubility and speciation of the volatiles, temperature and pressure (e.g. Unni & Schilling, 1978, Rowe & Schilling, 1979, Stolper & Holloway, 1988, Carroll & Webster, 1994, Dixon et al., 1995, Dixon, 1997, Rosenthal et al., 2015). Primitive CO2/Nb ratios are predicted to range from ~300 - 600 (Saal et al., 2002, Cartigny et al., 2008). Only a few samples from the ITD group, those having the lowest Nb contents, are in this CO2/Nb range, consistent with undegassed melt, with all other samples having much lower ratios (Fig. 6B). To assess the pressure of CO2-H2O saturation and the effects of degassing on the CO2 and H2O contents of the samples, we used the model of Dixon et al. (1995) to determine the pressure of CO2-H2O saturation and VolatileCalc (Newman & Lowenstern, 2002) to generate degassing paths that pass through the compositions of the samples. For the primitive compositions of the model, we used samples from the ITD and ITE groups with the highest concentrations of CO2 (ITD group: PL02-9-29, ITE group: D49A). Calculated degassing paths for both open ! 105! and closed system degassing predict extensive degassing of CO2 with only minor degassing of H2O occurring during closed system degassing (Fig. 7A). To get changes in H2O contents on the order of 26% (relative to the pressure), as is shown by the model, requires a melt to have a large H2O content initially and be erupted at a lower pressure than is typical of the high H2O content samples. Therefore, the samples are extensively degassed for CO2 and minimally degassed for H2O. The few samples in the ITD group with CO2/Nb ratios ≥300 can potentially be used to assess the CO2 content of the depleted component of the Galapagos mantle. The pressure of CO2-H2O vapor saturation ranges from 22 to 472 bars. F, Cl and S are usually considered not to be affected by degassing processes in hotspot and ridge settings for samples erupted at pressures higher than ~ 50 bars (Unni & Schilling, 1978, Rowe & Schilling, 1979, Carroll & Webster, 1994). None of the Galapagos samples have unusually low F/Nd or Cl/K ratios, suggesting that they have not degassed significant amounts of these volatile elements. However, there are a number of samples that were collected over a range of pressures that have unusually low S/Dy ratios. By comparing the calculated CO2-H2O saturation pressure to S/Dy ratios of both the melt inclusion data from Koleszar et al. (2009) and the glass data of this study it becomes apparent that for pressure of saturation below ~ 400 bars S/Dy ratios switch from being relatively constant to decreasing with decreasing saturation pressure (Fig. 7B; samples of this study filtered for samples with Sr/Sr* >1, see section on Assimilation). The trend in the glasses and melt inclusions is not due to mixing of melts with variable S/Dy ratios because both melt inclusions from Santiago (MORB component) and Fernandina (high 3 He/4He component) show relatively constant S/Dy ratios at high saturation pressures. ! 106! This suggests that S begins degassing at higher pressures than originally thought and that the low S/Dy ratios (<175) in the glasses of this study are the result of S degassing. An interesting observation can be made when comparing the calculated CO2-H2O saturation pressure with the collection pressure for glasses from the Galapagos and other OIBs with MORB samples (Fig. 7C). Most samples erupted at mid-ocean ridges (EPR, GSC etc.) on average fall above a 1:1 line indicating that they are supersaturated with CO2-H2O vapor at the depth of eruption, representing melt that migrated from the magma chambers to the surface quickly enough to prevent melt-vapor diffusive equilibration (le Roux et al., 2006 and references therein). This is also shown by the samples from the ITD group with all but one of the samples having CO2/Nb ratios ≥300 being supersaturated at the pressure of eruption. In contrast, samples analyzed from different OIBs, including the Galapagos glasses in this study, fall along the 1:1 line, indicative of melts that were able to equilibrate at every pressure until eruption and quenching. This suggests a fundamental difference between OIB and MORB in the ratio of the melt transport rate to the CO2-H2O diffusive gas-melt equilibration rate, the implication being that melts are upwelling in OIBs at a rate that allows for continuous gas-melt re- equilibration or that the composition of OIB melt allows for faster CO2 – H2O diffusion. Samples that fall below a 1:1 line between pressure of saturation and pressure of collection are thought to have either erupted at a shallower level and rolled down slope or are undersaturated in CO2-H2O. Of the samples from this study that fall below the 1:1 line, none of them are likely undersaturated in volatiles. All were either collected from local topographic lows or were collected from sections of steep topography, making it likely that these rolled downhill. ! 107! Sulfide Saturation Sulfide saturation in basaltic melt is dependent on temperature, pressure, oxygen fugacity, sulfur fugacity and bulk composition (Wallace & Carmichael, 1992, Carroll & Webster, 1994, Mavrogenes & O'Neill, 1999, Holzheid & Grove, 2002, Liu et al., 2007, Moune et al., 2009). To evaluate if the glasses reached sulfide saturation, we calculated the concentration of sulfur needed to saturate with a sulfide liquid using the model of Liu et al. (2007). The input parameters of this model include temperature, pressure, H2O content in the melt and a compositional parameter referred to as MFM that describes the melt bulk composition using cation mole fractions to represent the ratio of network modifiers to network formers. To calculate the pre-eruptive temperature of the melt, we use an equation from Wallace and Carmichael (1992), which utilizes the predictable variation of Al2O3 and MgO content of a melt undergoing crystal fractionation, with an incorporated correction for the effect of H2O content on the liquidus (Medard & Grove, 2008). Calculated temperatures vary from 1092-1255 °C with an estimated error in the calculation of 9 °C. We are interested in the pressure at which fractional crystallization is occurring and where sulfides could be separated from the melt. Therefore, we use 2 Kbars of pressure, which is the pressure used to calculate the LLDs in the results section above. Varying pressure between 2 Kbars and the pressure of vapor saturation (Dixon et al. (1995)) for individual samples does not significantly changes the results of the model and does not change our interpretation of those results. Model input H2O contents are taken directly from measurements. We use an average value of Fe+3/ΣFe measured by μ- ! 108! XANES (Fe+3/ΣFe = 0.17) for the glasses (Peterson et al., in prep) to determine the amount of ferric and ferrous iron to input for the calculations. The total variation in MFM is 7.07-8.62. The sulfide saturation model of Liu et al. (2007) was independently verified by the experiments of Moune et al. (2009). We consider this the best approximation because it is the only available model that incorporates the effect of water on sulfide saturation. The samples range from just sulfide saturated to under-saturated in a sulfide liquid at 2 Kbars of pressure (Fig. 8). None of the samples identified as having partitioned S into a gas phase are sulfide saturated. It seems that samples from array 1 are closer to sulfide liquid saturation than those from Array 2. Assimilation The variation in volatile elements with refractory elements in Fig. 6 indicates that while H2O, Cl, and S are correlated with Ce, K, and Dy, respectively, there is a large dispersion in the volatile data at a given refractory element value. Only F shows very little dispersion at a given value of Nd. Because degassing of H2O, Cl, and S (for the majority of the glasses) has been shown to be negligible, the dispersion is likely indicative of a secondary process affecting the volatile element contents and/or variation in the volatile/refractory element ratios of primary melts. High Cl/K and Cl/Nb ratios are used as indicators of assimilation of hydrothermally altered material due to the high concentration of Cl in seawater and brines percolating through the oceanic crust (Quinby-Hunt, 1983) compared to magmas. Uncontaminated mantle sources have Cl/K ratios that range from approximately 0.01 (MORB source) to 0.08 (enriched mantle source) (Michael, 1989, Michael & Cornell, ! 109! 1998, Stroncik & Haase, 2004, Workman et al., 2006). Oceanic lithosphere that has interacted with seawater tends to have large amounts of Cl, and thus, a much higher Cl/K ratio than uncontaminated melts (Michael, 1989, Jambon et al., 1995, Michael & Cornell, 1998, Kent et al., 1999a, Kent et al., 1999b, Coogan et al., 2003, Gillis et al., 2003, Straub & Layne, 2003, Stroncik & Haase, 2004, le Roux et al., 2006, Bonifacie et al., 2008, John et al., 2011). Therefore the assimilation of seawater-altered material is strongly reflected in the Cl/K ratios of the lavas. All of the samples in this study show Cl enrichment to some degree beyond what would be expected for an uncontaminated melt. The ITD group (Cl/K = 0.024-0.772) and the WD group (Cl/K = 0.032-0.211) show large ranges in Cl/K. The ITE and HHe groups have overlapping ranges in Cl/K ratios (Cl/K = 0.033-0.149). While a Cl/K ratio up to 0.09 has been attributed to the mantle source of other ocean islands (Workman et al., 2006), the total variations in Cl/K ratios of glasses analyzed for each compositional group do not correlate with any indicators of source or extent of melting (such as incompatible trace element and radiogenic isotope ratios). When considering only the samples with the lowest Cl/K ratios for each compositional group, then a very weak correlation is apparent between Cl/K ratios and Sr and Nd isotopes (Fig. 9). This suggests that the extreme variation in Cl/K within the different groups is not due to variation in the Cl concentration of the melt source, but rather is due to variable contamination by hydrothermally altered material, such as assimilation of 1.6% of a 50% NaCl brine, which could explain the highest Cl/K sample of the ITD group (Fig. 10). There is a statistically significant correlation between Cl/K and H2O/Ce for the samples from array 1 (ITE and HHe groups) (Fig. 10), suggesting that hydrothermally ! 110! altered material might have affected not only the Cl/K but also the H2O/Ce ratios of the melts. The range in H2O/Ce, like Cl/K for these groups overlaps (H2O/Ce = 109-328) with the maximum measured ratios for the ITE and HHe group glasses being higher than previously observed in other trace element enriched OIB (Dixon et al., 2002, Saal et al., 2002, Simons et al., 2002, Workman et al., 2006). H2O/Ce also does not correlate with isotopic ratios within a given group, as would be expected if the range in H2O/Ce ratios within each group is due only to mixing between a volatile rich and volatile poor melt. Thus, the correlation of H2O/Ce with Cl/K is most likely caused by the assimilation of hydrothermally altered rock in addition to any mixing that may be occurring. The samples in the ITE and HHe groups with Cl/K ratios lower than 0.08 (highest Cl/K ratio attributed to a plume source (Stroncik & Haase, 2004)), as a result, represent the least contaminated samples for Cl and H2O. Cl/K ratios do not correlate with F/Nd or S/Dy ratios for these groups. In the ITD group, CO2/Nb, H2O/Ce, F/Nd and S/Dy ratios weakly correlate with values of Sr/Sr* ≥ 1, while Cl/K ratios are hugely variable reaching values up to 0.772, ten times higher than expected for uncontaminated samples (Fig. 11). The high Sr/Sr* values in the ITD group indicate that these samples have experienced interaction with a plagioclase-rich cumulate before subsequent crystallization of plagioclase (as is discussed in Results and shown in Fig. 3) (Saal et al., 2007). Products of hydrothermal alteration, such as amphibole and serpentine, serve to concentrate volatile elements in the oceanic crust (e.g. Gillis & Meyer, 2001, Adam & Green, 2011, John et al., 2011, Alt et al., 2012). Although the mode produced by hydrothermal alteration is highly variable, plagioclase is ubiquitous in the oceanic crust, and occurs as both a primary (magmatic) ! 111! and secondary (hydrothermal) phase (Alt et al., 1996a, Alt et al., 1996b, Bach et al., 2003, Coogan et al., 2003, Barker et al., 2008, Heft et al., 2008, Kirchner & Gillis, 2012). The assimilation of plagioclase and associated hydrothermally altered cumulates, therefore, can produce Sr/Sr* values > 1 accompanied by high volatile/refractory element ratios. The high volatile refractory element ratios with the highest Sr/Sr* values in the ITD group lavas having high MgO content are, therefore, most likely caused by the early assimilation of hydrothermally altered crust followed by crystal fractionation. These results are supported by the extremely high Cl/K ratios an order of magnitude above the 0.02 expected for a MORB source (Michael, 1989, Michael & Cornell, 1998, Stroncik & Haase, 2004). The broadly positive correlation between Sr/Sr* and CO2/Nb, H2O/Ce, F/Nd and S/Dy, however, cannot be explained by crystal fractionation (which will only decrease Sr/Sr*), rather this correlation suggests the higher MgO melts that interacted with plagioclase cumulates are mixing with more evolved ITD melts less affected by contamination. The dispersion in the volatile to refractory trace element ratios defined by the glasses from the WD group most likely are the result of mixing with an ITD melt after the ITD samples are contaminated and/or before they are degassed. The WD group end- member has Cl/K, H2O/Ce, F/Nd and S/Dy ratios essentially consistent with those previously reported for other OIB (Cl/K ≤ 0.08, H2O/Ce ~200, F/Nd ≥ 21, S/Dy ~260) (Dixon et al., 2002, Simons et al., 2002, Stroncik & Haase, 2004, Workman et al., 2006, Koleszar et al., 2009, Cabral et al., 2014, Kendrick et al., 2014). Mixing of this enriched end-member with melts from the ITD group would cause the increasing CO2/Nb ratios and decreasing F/Nd ratios as WD samples approach the isotopic signature of the ITD ! 112! end-member (Fig. 13 and 15). Because the highest CO2/Nb ratio of the ITD group represents the least degassed of the ITD glasses and may be the product of contamination (due to the high Sr/Sr*) this is evidence that the mixing between the WD and ITD groups must have occurred after the ITD group was contaminated and/or before it degassed. The range in the WD group also overlaps the volatile variation of lavas from the GSC (Fig. 5, 9, S1 and S2). Mantle Composition A number of volatile studies in OIB indicate that isotopic heterogeneities in the mantle have well-defined, associated volatile/refractory element ratios (Dixon & Clague, 2001, Dixon et al., 2002, Simons et al., 2002, Workman et al., 2006, Koleszar et al., 2009, Cabral et al., 2014, Kendrick et al., 2014). In this study, after accounting for variability caused by shallow level processes, unique volatile/refractory element ratios are apparent in the end-member compositions of array 1 and 2 (Table 3). For array 1, the largest contrast is in F/Nd ratios (Fig. 12) and H2O/Ce ratios (Fig. 13). The HHe group end- member (Fernandina mantle component) possesses an H2O/Ce ratio (202 ± 26) and a F/Nd ratio (20.1 ± 0.8) that is close to that generally predicted for the FOZO mantle component based on a limited number of studies (F/Nd ~ 21, H2O/Ce ~ 200, McDonough and Sun, 1995, Dixon et al., 2002, Simons et al., 2002, Koleszar et al., 2009). The ITE group end-member (Sierra Negra/Cerro Azul mantle component), in contrast, shows a low H2O/Ce ratio (H2O/Ce = 174 ± 24) and the highest F/Nd ratio (24.9 ± 0.7) of all the samples analyzed. S/Dy ratios and Cl/K ratios are similar between the HHe group (S/Dy ! 113! = 218 ± 21, Cl/K = 0.057 ± 0.009) and the ITE group (S/Dy = 237 ± 20, Cl/K = 0.049 ± 0.010) end-members. For the ITE group samples, there is a correlation between 3He/4He ratios and Δ7/4 (Fig. 12), which suggests that melts from the WD and HHe groups are influencing the composition of the ITE group samples. This is consistent with the lower H2O/Ce ratio of the ITE group, which is at the upper limit of what has been measured in OIB with EM mantle components, such as Samoa (58-157) (Workman et al., 2006), Pitcairn (80) and Society (60) (Kendrick et al., 2014), and EM-influenced MORB such as the Shona, Discovery and Great Meteor anomalies (<100) (Dixon et al., 2002). The F/Nd ratio of the ITE end-member, however, is distinct from that of either the WD or HHe groups. The sample with the highest F/Nd ratio in the ITE group (F/Nd = 26.5), has a 3He/4He ratio = 13.1 Ra, whereas samples from this group having lower and higher 3He/4He ratios have lower F/Nd ratios (Fig. 12). It has also been suggested that the ITE mantle component could be the result of a mixture between the HHe group (Fernandina) and the HIMU mantle component of Floreana (e.g. Harpp and White, 2001). The F/Nd ratio of the ITE end-member, though higher than the HHe end-member, is lower than the value found for the HIMU end-member represented by Mangaia melt inclusions (Austral Islands) (F/Nd = 30) (Cabral et al., 2014). The volatile/refractory element ratios of the ITE end-member can therefore be due to either a mixture of melts from several sources or a unique, intermediate composition. The high 3He/4He ratio of the Fernandina group is thought to be due to the presence of a volatile-rich primitive mantle component. Using the volatile/refractory element ratios of the HHe end-member and the K, Ce, Nd, and Dy content of the ! 114! primitive mantle (240 ppm, 1.675 ppm, 1.25 ppm and 0.674 ppm respectively of (Mcdonough & Sun, 1995)) constrains the Cl, H2O, F and S content of the HHe mantle source to 13.7 ± 2 ppm, 338 ± 44 ppm 25.3 ± 1 ppm and 147 ± 14 ppm respectively. These values are generally lower than the estimated volatile content of the primitive mantle (17, 750, 25, 250 ppm) (Mcdonough & Sun, 1995). Understanding how the volatile/refractory element ratios of the ITE group are related to the volatile content of the ITE mantle component, in contrast, is more complex than the HHe group. The differences in F/Nd and H2O/Ce between the ITE and HHe groups are probably not the result of a source with homogenous volatile contents and heterogeneous trace element contents because the isotopic signature of the ITE group would suggest that the ITE mantle source is enriched in trace elements over the source of the HHe group. This would produce both lower F/Nd and H2O/Ce ratios, which is not observed. However, quantifying the difference in the volatile content of the ITE group mantle source from the HHe group mantle source is difficult because this end-member is either the product of a mixture (e.g., Fernandina and Floreana), or a unique intermediate mantle source, which makes the trace element composition of the source unconstrained. Therefore we limit the estimation of the volatile content of the ITE source to the volatile/refractory element ratios reported above. For the samples that fall on array 2, the volatile contents of the end-member mantle compositions are difficult to estimate due to the strong evidence of contamination in the ITD group samples and mixing with the WD melts. For the ITD group, only the minimum Cl/K ratios (0.026 ± 0.002) and the corresponding F/Nd (17.3 ± 1.5) and S/Dy ratios (228 ± 16) of the glasses are within range of the Santiago melt inclusions analyzed ! 115! by Koleszar et al. (2009) (Fig. 13) and yield F (12.5 ± 1 ppm), Cl (1.3 ± 0.1 ppm) and S contents (115 ± 8) of the mantle source (using mantle contents of Nd = 0.703, K = 50 and Dy = 0.505, for the DMM of Workman and Hart (2005)) similar to those estimated for the depleted upper mantle (F = 16 ± 3, Cl = 0.09 ± 0.63, S = 146 ± 35; Saal et al. (2002)). The highest CO2/Nb ratios of the ITD group are in the range found in MORB (Saal et al., 2002, Cartigny et al., 2008), but these glasses are not undersaturated in volatiles, so degassing will have decreased CO2/Nb ratios and these glasses are associated with the highest values of Sr/Sr*, so contamination may have increased CO2/Nb ratios (Fig. 11). In addition, all of H2O/Ce ratios of the ITD group exceed what is expected of the depleted upper mantle (150-200; e.g. Dixon et al., 2002, Saal et al., 2002, Koleszar et al., 2009). As a result, it is difficult to know how representative the CO2/Nb and H2O/Ce ratios of the ITD glasses are of the depleted mantle source in the Galapagos, therefore we restrict the values to what has previously been reported for the CO2/Nb and H2O/Ce ratios for this mantle component from melt inclusion studies (Koleszar et al., 2009). Samples from the WD group end-member that have the isotopic signature most similar to Pinta-Wolf-Darwin Is. are the least affected by mixing with contaminated melts from the ITD group. This yields S/Dy (255 ± 40) and Cl/K values (0.066 ± 0.006) similar to what has been reported for other ocean islands with evidence for recycled crustal material. The H2O/Ce ratio (230 ± 35), while similar to ratios found in Mangaia melt inclusions and samples collected from Easter Island and the Salas y Gomez seamount chain (Simons et al., 2002, Cabral et al., 2014), is higher than what was found for other EM type sources such as Samoa, Pitcairn and Society (Workman et al., 2006, Kendrick et al., 2014). The F/Nd ratio (23.5 ± 0.5) is close to the ITE end-member, and higher than ! 116! what is estimated for the primitive mantle. Though the isotope signature of the Pinta- Wolf-Darwin end-member is indicative of a recycled crustal source, the exact trace element composition of this source is difficult to constrain, thereby limiting the estimation of the volatile content of this mantle component to the reported volatile/refractory element ratios for the WD end-member. The effect of H2O content on mantle rheology The H2O content of the mantle source will have a strong effect on the viscosity of the mantle. Despite the HHe group having a lower H2O content than what is predicted for the primitive mantle, it is still 5x the concentration predicted for the MORB-source in the Galapagos (338 ppm versus 64 ppm based on a H2O/Ce ratio for the depleted source of 117 (Koleszar et al., 2009) and a DMM Ce content of 0.55 from Workman and Hart (2005)). These values can be used in conjunction with the power law form of the dependence of strain rate (!!) on differential stress (σ) of an olivine aggregate (Hirth & Kohlstedt, 1996, Hirth & Kohlstedt, 2003) ! ∗ !!! ∗ !! = !! ! ! !! ƒ!! !! exp !" (1) where A is a material parameter, n is the stress exponent, d is grain size, p is the grain size exponent, ƒH2O is water fugacity, r is the water fugacity exponent, E* is the activation energy, V* is the activation volume, R is the gas constant and T is the absolute temperature, to estimate the difference in viscosity we would expect between the HHe component of the plume and the ITD component. For both the diffusion creep regime and ! 117! dislocation creep regime we can neglect the effect of the presence of melt on the strain rate at pressures greater than the estimated onset of carbonated melting (~200 km; Villagomez et al., 2014). For a constant differential stress (σ = 0.3 MPa) and values for the wet diffusion creep regime listed in table 4, the viscosity of the HHe mantle component would be ~5x less than predicted for the ITD component. For the wet dislocation creep regime this number increases to ~7x less viscous. This has important implications for modeling plume ridge interaction, as a lower viscosity will increase the rate of solid flow, which will affect how plume material might flow towards the ridge, how melt is retained in the mantle, and melt production processes. C/3He ratios of the high 3He/4He mantle source It is difficult to estimate the primitive CO2 contents of the Galapagos glasses due to extensive degassing. Information about the mantle source, however, can be inferred by combining glass He concentrations of melt fractions, previously measured by Kurz et al. (2009) with glass CO2 contents measured by SIMS. This is due to the similar solubilities of CO2 and He in basaltic melt (Lux, 1987, Pan et al., 1991, Dixon et al., 1995, Jendrzejewski et al., 1997), which results in C/3He ratios of the glass that reflect the C/3He ratio of the mantle. Because CO2 measurements are obtained by SIMS, we are only able to calculate C/3He ratios for samples that have been measured for He isotopes utilizing a melting step, as this measures He retained in the glass matrix (as opposed to vesicles). This limits us mainly to glasses in the HHe group, with only a few samples from the ITE group and the one sample from the ITD group collected near Fernandina. C/3He ratios can be found in Table 1. ! 118! Calculated C/3He ratios of Galapagos glass average 4.37 ± 9 x 1010, though this number is skewed by one low He concentration glass (AHA19A). When this glass is excluded the average shifts to 3.10 ± 6 x 1010. This number is an order of magnitude higher than the canonical upper mantle value (DMM C/3He = ~2 ± 1 x 109; (Marty & Jambon, 1987, Marty & Tolstikhin, 1998, Shaw et al., 2004, Barry et al., 2014)), and values measured for mantle plumes (2-20 x 109 (Trull et al., 1993, Marty & Tolstikhin, 1998, Marty & Zimmermann, 1999, Barry et al., 2014)). Possible causes for variations in C/3He ratios include degassing, assimilation of shallow lithospheric material, or variation in the mantle source. Degassing could fractionate the C/3He of the residual melt due to the slightly higher solubility of He in comparison to CO2 (SHe = 6.4 x 10-4 cm3 STP/g while SCO2 = 2.7x10-4 cm3 STP/g experimentally derived solubility constants at 1 atm (Lux, 1987, Pan et al., 1991, Dixon et al., 1995, Jendrzejewski et al., 1997)). This would cause a decrease in the C/3He ratio of the melt as degassing progresses, which does not explain the higher C/3He ratios of Galapagos glasses. Assimilation of either a carbonate rich source in the crust or sediments could cause the higher C/3He ratios (Shaw et al., 2004). However, assimilation of crust or sediment should manifest in other trace element proxies, such as elevated Sr/Ce ratios that would potentially correlate with C/3He ratios in Galapagos glasses, neither of which is observed. The most likely explanation for the higher C/3He ratios in the Galapagos glasses is that a high 3He/4He mantle source component in the Galapagos plume is more carbonated than other oceanic basalts. This explanation is consistent with geochemical modeling that invokes a carbonatite mechanism for helium displacement in plume settings (Hofmann et al., 2011) and recent melt modeling of the Galapagos hotspot ! 119! by Villagomez et al. (2014), who invoke the production of carbonatite to explain the low velocity anomalies at greater than 150 km depth identified by seismic imaging. CONCLUSIONS The geochemical data presented in this paper define two mixing arrays: array 1 ranging between the composition of Fernandina Island and Sierra Negra/Cerro Azul volcanoes (HHe group and ITE group respectively) and array 2 between the composition of Pinta- Wolf-Darwin Islands and Genovesa Island (WD and ITD groups respectively). 4 samples from Wolf Volcano did not fall on either array, but this volcano has been shown to be anomalous geochemically in many respects, because it is isotopically depleted but has enriched incompatible trace element concentrations (Geist et al., 2005). The principle findings of this study can be summarized as follows: 1.) All of the glasses, except for the subset of samples from the ITD group that are supersaturated with CO2-H2O fluid, have extensively degassed CO2. There is minimal degassing of the other volatile species, except for a subset of samples that have degassed S in addition to CO2. 2.) The pressure of CO2-H2O vapor saturation versus the pressure of sample collection for the samples, in combination with data previously published for other oceanic basalts, reveal a fundamental difference between MORB and OIB; where either lower melt upwelling rate or faster C and H diffusion in the melt allowed for the continuous equilibration of vapor and melt in OIB compared to MORB. ! 120! 3.) The Galapagos submarine glasses range from sulfide saturated to undersaturated. 4.) Assimilation of hydrothermally altered material has affected the Cl contents of a subset of glasses from all of the groups. Cl/K ratios are correlated with H2O/Ce ratios for the HHe and ITE groups showing that H2O also might be affected by assimilation of hydrothermally altered material. Furthermore samples from the ITD group with the highest Sr/Sr* ratios have the correspondingly highest H2O/Ce, CO2/Nb, F/Nd, and S/Dy ratios. This suggests that the assimilation of plagioclase-rich cumulates is accompanied by the assimilation of volatile rich phases in the shallow lithosphere. 5.) The correlations between CO2/Nb and F/Nd and isotopes in the glasses of the WD group suggests that melts from the WD group are mixing with variably contaminated and/or less degassed samples from the ITD group at depth. 6.) The new data yields a first-order estimate of source volatile contents of the end- members defining the two geochemical arrays a. For the samples of array 1, the HHe group end-member (Fernandina) shows a H2O/Ce ratio (202 ± 26) and F/Nd ratio (20.1 ± 0.8). The ITE end-member (Sierra Negra/Cerro Azul), in contrast, shows a higher F/Nd ratio (24.9 ± 0.7) and a H2O/Ce ratio (174 ± 24) that is lower but within error of the HHe end-member, similar to that previously reported for an enriched mantle source, which may be indicative of recycled crust. S/Dy and Cl/K ratios for the two end-members of the array 1 are very similar (S/Dy = 218 ± 21, Cl/K = 0.057 ± 0.009 and S/Dy = 226 ± 20, Cl/K = 0.048 ± 0.011 for the HHe and ITE groups respectively). ! 121! b. The estimated H2O/Ce ratio for the primitive Galapagos mantle beneath Fernandina, if it has a refractory trace element composition similar to the PM, indicates that the H2O, F, Cl, and S contents for the source of the HHe samples are generally lower than what is predicted for the primitive mantle. c. Only the minimum Cl/K (0.026 ± 0.002), F/Nd ratios (17.3 ± 1.5), and S/Dy ratios (228 ± 16) of the Genovesa end-member (ITD group) from array 2 fall within range of MORB. d. The Pinta-Wolf-Darwin end-member (WD group) has high Δ7/4, Δ8/4 and Th/La ratios and low 3He/4He ratios suggesting the presence of a recycled crustal component. The volatile data for this end-member yield a similar S/Dy ratio (255 ± 40) to those of the end-members in array 1, have slightly higher Cl/K ratio (0.066 ± 0.006), an H2O/Ce ratio (230 ± 35), and a F/Nd ratio (23.5 ± 0.04) that is higher than the estimate ratio for the Fernandina mantle component, but lower than what was found for the ITE end-member. 7.) The higher H2O contents of the Fernandina HHe source will result in a mantle viscosity 5-7x lower than what we would predict for the ITD mantle component. 8.) C/3He ratios in the glasses of the Fernandina HHe end-member are, on average, an order of magnitude higher than what has been found previously for MORB and other OIB suggesting that the Galapagos plume contains a mantle component that is carbonated. ! 122! Distinctive H2O/Ce and F/Nd ratios, as a result, can differentiate the four main compositional components defined by the glasses of this study. The generally higher H2O/Ce ratios of the HHe, ITE, and WD groups suggests that the H2O contents of these sources is higher than the depleted upper mantle which will affect the physical properties of the mantle such as the variability of mantle viscosity. These effects in addition to the high C/3He ratios of the HHe group will be important to consider in future modeling plume-ridge interactions and mantle flow. ACKNOWLEDGEMENTS This work was supported by the National Science Foundation Graduate Research Fellowship (Grant No. DGE-1058262 to M.E.P.), the National Science Foundation Division of Ocean Sciences (Grant No. 0962195) and the German Ministry for Education and Research (BMBF) grant Sonne cruise SO158. Nobel gas measurements at WHOI were supported by NSF OCE. We’d like to thank F. Hauff, C. Jackson, T. Prissel, and K. Shimizu for their thoughtful discussions during the preparation of this manuscript. ! 123! References: Adam, J. & Green, T. (2011). 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Annual Review of Earth and Planetary Sciences 14, 493-571, Doi 10.1146/Annurev.Earth.14.1.493. ! 144! FIGURE CAPTIONS Figure 1. Map of the Galapagos archipelago with locations of sample dredges. Bathymetry is indicated by the color scale on the left side on the map and prominent features are labeled. Array 1 consists of samples that range between the compositions of Fernandina Island and Sierra Negra/Cerro Azul volcanoes that are further divided into HHe and ITE groups based on samples with 3He/4He ratios above and below 3He/4He = 22 Ra, respectively. Array 2 consists of samples that range between the compositions of Genovesa Island and Pinta, Wolf and Darwin Islands that have been divided into the ITD group (samples with more MORB-like compositions) and the WD group (samples characterized by high Th/La ratios and high Δ7/4, Δ8/4). Samples that have a composition similar to Wolf Volcano on Isabela are also shown separately. Map created using GeoMapApp http://www.geomapapp.org (Ryan et al., 2009). Figure 2. (A-F) Major element concentrations plotted against MgO wt% of individual glasses. FeO* indicates total Fe content. Superimposed on the glass data are the predicted liquid lines of descent (LLD) calculated using PetroLog3 (Danyushevsky & Plechov, 2011). The calculations are performed isobarically at 2 Kbars of pressure for 2 different starting compositions, which can be found in table 2. Notice the entire range in major elements cannot be reproduced by a single LLD. Blue LLD calculated from primitive melt inclusion AHA-D25C-3-43 (indicated by a grey triangle outlined in blue), collected from a submarine dredge with an enriched composition similar to those of Fernandina lavas. Green LLD calculated from a primitive melt inclusion STG06-29-06 (indicated by ! 145! a grey triangle outlined in green), collected from Santiago, with a depleted composition similar to MORB. Compositions were chosen so that LLDs would bracket the range in data. Melt inclusion compositions reported in Koleszar et al. (2009) are shown by labeled grey fields. Errors are 2σ standard error. Figure 3. Sr/Sr* versus A.) Al2O3 wt% and B.) MgO wt% of individual glasses. Sr/Sr* is quantified by SrPM/(((PrPM+NdPM)/2). Included in both figures are grey fields that represent melt inclusion data from Santiago and Fernandina reported in Koleszar et al. (2009) and Peterson et al. (2014) and LLDs as calculated in Figure 2. Notice that the data from the ITD group fall outside the area defined by the LLD. In Fig. 2A ITD Group data follow the same trend between Sr/Sr* and Al2O3 defined by the Santiago melt inclusions, which were found to undergo simple plagioclase assimilation. However in Fig. 2B the ITD Group samples define the opposite trend to the Santiago melt inclusions, with only the highest MgO content and Sr/Sr* glasses overlapping with the middle of the field defined by the melt inclusions. This indicates that the glasses with the highest MgO wt% have initially assimilated plagioclase cumulates and later began to crystallize plagioclase out, causing the opposite trend to the Santiago melt inclusions in Sr/Sr* versus MgO wt%. Errors are 2σ standard error. Figure 4. Primitive mantle normalized trace element diagram. Samples representing the different groups (HHE Group: D24A, D28C, ITE Group: D6G, D8A, WD Group: PL02- 29-3, PL02-28-11, ITD Group: DR70-1, PL02-10-8) have similar MgO contents (6.08- 7.84 wt %) and trace element contents representative of the overall characteristics of each ! 146! group. For example, notice the positive Ti anomaly in the HHe Group samples. The grey field shows the range in previously collected trace element data in whole rocks from across the Galapagos Archipelago limited to samples with MgO contents between 6.0 and 8.0 wt % (McBirney et al., 1969, Baitis & Lindstrom, 1980, Cullen et al., 1989b, Vicenzi et al., 1990, Bow & Geist, 1992, White et al., 1993, Geist et al., 1995, Reynolds & Geist, 1995, Standish et al., 1998, Geist et al., 2002, Schilling et al., 2003, Saal et al., 2007, Gibson et al., 2012). Inset shows the high MgO content ITD and WD Group glasses that have the largest Sr anomalies and have been affected by plagioclase assimilation. Normalization values from Mcdonough and Sun (1995). Figure 5. (A-C) 143Nd/144Nd and (D-F) 206Pb/204Pb versus 3He/4He, 87Sr/86Sr and 208Pb/204Pb. Notice glass samples with 3He/4He > 22Ra correspond to the Pb isotope composition indicative of Fernandina whole rocks. Notice also that the WD and ITD Groups define a compositional range from near Genovesa to near Pinta-Wolf-Darwin Islands. Some of the isotopic data is reported in Geist et al., (2008) and Harpp and White, (2001); the rest is new data. Both previously reported and new data are presented in Table 1. The grey fields are defined by previously published isotope data on whole rocks basalts from different islands and the GSC (Cullen et al., 1989b, White et al., 1993, Reynolds & Geist, 1995, Blichert-Toft & White, 2001, Schilling et al., 2003, Saal et al., 2007, Ingle et al., 2010) while the yellow and green fields show the position of unpublished isotopic data for the ITD and WD groups respectively from the SONNE SO158 cruise (Hoernle, personal communication). ! 147! Figure 6. Volatile contents versus MgO wt% and refractory trace elements with similar incompatibility. LLDs, represented by blue and green lines shown in the volatile versus MgO wt% plots, are calculated as in Figure 2. Partition coefficients for volatile elements presented in Table 2. (A-B) CO2 ppm versus MgO wt% and Nb ppm. Notice that Santiago inclusion chosen for LLD modeling is degassed for CO2. An inclusion with a higher CO2 content was not chosen for modeling because they were both more evolved and showed a larger signature of assimilation (higher Sr/Sr*). (C-D) H2O wt% versus MgO wt% and Ce ppm. (E-F) F ppm versus MgO wt% and Nd ppm. (G-H) Cl ppm versus MgO wt% and K ppm. (I-J) S ppm versus MgO wt% and Dy ppm. Black lines delineate the volatile/refractory element ratios that bracket the data (in the case of H2O versus Ce, F versus Nd, Cl versus K, and S versus Dy) or are commonly cited mantle values for ratios of volatiles and refractory elements (in the case of CO2 versus Nb). Notice that while H2O, F, Cl, and S contents increase with decreasing MgO wt%, indicating that crystal fractionation has some control over the volatile content, the CO2 content is relatively constant, indicating CO2 is controlled by degassing. Figure 7. A.) CO2 ppm versus H2O wt% concentrations of the glasses. In addition to the data are the CO2-H2O equilibrium saturation pressures (horizontally curved black lines) and equilibrium vapor phase compositions (vertically curved black lines) from the model of Dixon et al. (1995), and three degassing paths calculated using the model of Newman and Lowenstern (2002). Dashed red lines show open system degassing, calculated from the composition of two samples with the highest CO2 content from the ITD and ITE Groups (PL02-9-29, D49A respectively). Dashed and dotted red line shows the closed ! 148! system degassing path calculated from D49A (assuming 1.8 wt% vapor exsolved). Errors are 2σ standard errors. B.) S/Dy ratios versus pressure of CO2-H2O vapor saturation, calculated from the model shown in panel A (Dixon et al., 1995) for individual glasses. Melt inclusions from Koleszar et al. (2009) are also included. Notice that for P saturation values < 400 bars S/Dy ratios switch from being relatively constant to gradually decreasing, which suggests S degassing. C.) Pressure of CO2-H2O vapor saturation, calculated from the model shown in panel A (Dixon et al., 1995), versus individual pressures of collection. The error in the pressure of saturation is based on the error in the analysis of CO2. The error on the pressure of collection represents the depth change over the course of each dredge. Fields for the Galapagos spreading center (GSC), the East Pacific Rise (EPR), the Loihi seamount, Pitcain Island and from different islands in Samoa and Society archipelagos are plotted for comparison (Dixon & Clague, 2001, Schilling et al., 2003, Cushman et al., 2004, le Roux et al., 2006, Workman et al., 2006, Ingle et al., 2010, Kendrick et al., 2014, Shimizu et al., In preparation). Solid black line is a 1:1 line. Figure 8. Log values of measured sulfur concentrations versus log of calculated S concentration at sulfide saturation (SCSS) using the model of Liu et al. (2007). The calculation is based on the major element composition, pressure, temperature, and the concentration of water of a sample (see text for in depth discussion of model). Black solid line is the 1:1 line at and above which samples are considered sulfide saturated. Black dashed lines represent the 200 ppm error of the model based on the experiments of ! 149! Moune et al. (2009). Based on this model, the glasses range from sulfide saturated to undersaturated. Included error represents the average 2σ standard error of S measurement. 87 Figure 9. Volatile/refractory element ratios versus (A-F) Sr/86Sr ratios and (G-L) 143 Nd/144Nd ratios for individual glasses. Also included are grey fields representing previously reported data from the GSC (Ingle et al., 2010) and yellow and green fields showing the position of unpublished isotopic data for the ITD and WD groups respectively from the SONNE SO158 cruise (Hoernle, personal communication). Errors are 2σ standard errors. Figure 10. Cl/K ratios versus H2O/Ce ratios of the ITE and HHe Group lavas. Black solid line shows calculated York regression. The relatively good linear correlation shows that assimilation of hydrothermally altered material might be simultaneously affecting Cl and H2O contents of the glasses. Dashed box shows the expected mantle limits of a trace element enriched source (Dixon et al., 2002, Simons et al., 2002, Stroncik & Haase, 2004, Workman et al., 2006). Inset shows Cl/K ratios versus H2O/Ce ratios for all the glasses of this study. Lines show compositional effect of adding seawater (SW) (Cl = 1.935 wt%, H2O = 97.5 wt%, K = 340.2 ppm, and Ce = 0.001 ppb), a 15% NaCl brine (Cl = 9.9 wt%, H2O = 85 wt%, K = 2075 ppm, and Ce = 0.006 ppb), and a 50% NaCl brine (Cl = 30.3 wt%, H2O = 50 wt%, K = 6474 ppm, and Ce = 0.02 ppb) (Kent et al., 1999a, Kent et al., 1999b) to the composition of sample D20a. Ticks are labeled for the percent material mixed. Errors are 2σ standard errors. ! 150! Fig. 11. Volatile/refractory element ratios versus Sr/Sr* for individual glasses. Errors are 2σ standard errors. Figure 12. 3He/4He ratios versus A.) Δ7/4 and B.) F/Nd ratios. Δ7/4 calculations are based on equations found in Hart (1984) and represent positive and negative deviations from the northern hemisphere reference line. Errors are 2σ standard errors. 208 Figure 13. Volatile/refractory element ratios versus (A-E) Pb*/206Pb* and (F-J) Th/La 208 ratios for individual glasses and array end-member compositions. Pb*/206Pb* = (208Pb/204Pb-29.4761)/(206Pb/204Pb-9.3066). End-member compositions for the ITE, HHe, and WD groups are determined from samples filtered to Cl/K ratios ≤ 0.08 and S/Dy ratios ≥ 175 that are not sulfide saturated. Additionally, for ITE group, end-member components are the mean and median of the samples with the highest F/Nd ratios. These represent an end-member of the correlation between F/Nd and 3He/4He ratios defined by array 1 (Fig. 13). For the HHe group, because the in-group F/Nd ratio variation is lower, the end-member components are the average and median values of the filtered samples. The WD end-member is defined by the mean and median values of the samples with the isotopic composition closest to the Pinta component. F/Nd, Cl/K and S/Dy ratios of the ITD end-member composition are set by the samples with the lowest Cl/K ratios. CO2/Nb ratios are set by the expected CO2/Nb ratio of a MORB source (Saal et al., 2002) and H2O/Ce end-member set by the average value of Santiago normal inclusions reported in Koleszar et al. (2009). End-member compositions presented in Table 3 Dashed lines show the values generally cited for the FOZO and MORB sources (e.g. Dixon et al., 2002, ! 151! Saal et al., 2002, Simons et al., 2002). Yellow and green fields showing the position of unpublished isotopic data for the ITD and WD groups respectively from the SONNE SO158 cruise (Hoernle, personal communication). Errors bars showing the 1 standard deviation of samples used to set end-member compositions. Supplementary Figure 1. Th/La ratios versus A.) Sm/Yb ratios, B.) Ti/Gd ratios and C.) Nb/Zr ratios of individual glasses plotted with fields defining the ranges of previously measured whole rock and glass samples from the Galapagos Archipelago and the Galapagos Spreading Center (GSC) (references as in Fig. 4). Notice that glasses of this study cover almost the entire geochemical range shown by the literature data from trace element depleted to enriched signatures. Errors are 2σ standard error. 87 Supplementary Figure 2. (A-D) Sr/86Sr, (E-H) 143 Nd/144Nd, and (I-L) 208 Pb*/206Pb* versus Sm/Yb, Ti/Gd, Th/La, and Nb/Zr ratios of individual glasses with a field showing the range of data collected from the GSC and for selected islands from the Galapagos 208 Archipelago. Pb*/206Pb* = (208Pb/204Pb-29.4761)/(206Pb/204Pb-9.3066) is the ratio of the radiogenic additions to the initial terrestrial Pb. Trace element ratios are correlated with isotopic ratios indicating that the trace element variations are long-lived characteristics of the mantle sources. Notice 2 distinct arrays of data; the first array being composed of the HHe and ITE Groups and the second one being composed of the WD and ITD groups. Yellow and green fields showing the position of unpublished isotopic data for the ITD and WD groups respectively from the SONNE SO158 cruise (Hoernle, personal ! 152! communication). References for fields and isotopic data as in Figure 4. Errors are 2σ standard error. Supplementary Figure 3. 3He/4He versus Ti/Gd of individual glasses. Notice the weak trend between 3He/4He and Ti/Gd. This is often cited as a characteristic of the FOZO mantle source. References are reported in Figure 6. Errors are 2σ standard error. ! 153! FIGURES Figure 2-1 Galapagos Spreading Center (GSC) 2°N 30’ HHe Group ITE Group Array 1 2°N ITD Group Array 2 WD Group Wolf Volcano 1°N 30’ Wolf-Darwin lineament 1°N Marchena Pinta 0°N 30’ Genovesa NE Seamounts 0° Santiago Fernandina 0°S 30’ Santa Cruz Isabela San Cristobal 1°S Floreana 1°S 30’ 2°S 94°W 93°W 92°W 91°W 90°W 89°W 88°W 87°W ! 154! Figure 2-2 A. Fernandina MI D. 5 13 CaO (wt%) Fernandina MI TiO2 (wt%) 4 12 3 11 2 Santiago MI 1 10 Santiago MI B. E. 18 Na2O (wt%) Fernandina MI 3.5 17 Al2O3 (wt%) 16 3.0 Santiago MI 15 2.5 14 Santiago MI 2.0 13 Fernandina MI C. F. 14 1.25 K2O (wt%) Santiago MI FeO* (wt%) 1.00 12 0.75 10 Fernandina MI 0.50 8 0.25 Fernandina MI Santiago MI 5 6 7 8 9 10 11 5 6 7 8 9 10 11 MgO (wt%) MgO (wt%) HHe Group ITE Group ITD Group WD Group Wolf Volcano ! 155! Figure 2-3 2.25 A. B. 2.00 Santiago MI 1.75 Santiago MI 1.50 Sr/Sr* 1.25 1.00 0.75 Fernandina MI Fernandina MI 0.50 0.25 12 13 14 15 16 17 18 5 6 7 8 9 10 11 12 Al2O3 (wt%) HHe Group MgO (wt%) ITE Group ITD Group WD Group Wolf Volcano ! 156! Figure 2-4 100.0 HHe Group ITE Group ITD Group WD Group Sample/Primtive Mantle 10.0 10 1.0 1 0.1 Rb Th Nb La Pb Sr P Hf Eu Gd Dy Ho Tm Lu Ba U K Ce Pr Nd Zr Sm Ti Tb Y Er Yb 0.1 Rb Th Nb La Pb Sr P Hf Eu Gd Dy Ho Tm Lu Ba U K Ce Pr Nd Zr Sm Ti Tb Y Er Yb ! 157! Figure 2-5 A. D. 2σ 2σ 25 Fernandina Fernandina HHe Group Sierra Negra/ 20 ITE Group Cerro Azul 3He/4He ITD Group Sierra Negra/Cerro Azul WD Group Wolf Volcano Floreana 15 Floreana Wolf Vol. Wolf Vol. Pinta 10 Pinta GSC GSC B. Floreana Pinta E. 2σ 2σ Pinta Sierra Negra/ 0.7034 Cerro Azul Fernandina Wolf Vol. Floreana 87Sr/86Sr GSC GSC 0.7030 Genovesa Sierra Negra/ Cerro Azul Fernandina 0.7026 Wolf Vol. Wolf/Darwin Is. Wolf/Darwin Is. Genovesa C. Floreana 2σ 2σ F. 39.6 Sierra Negra/ Pinta Cerro Azul Wolf/Darwin Is. 39.2 208Pb/204Pb Floreana Pinta Wolf/Darwin Is. 38.8 Sierra Negra/ 38.4 Wolf Vol. GSC Cerro Azul Fernandina GSC Fernandina 38.0 Wolf Vol. Genovesa Genovesa 0.5129 0.5130 0.5131 18.4 18.6 18.8 19.0 19.2 19.4 19.6 19.8 143Nd/144Nd 206Pb/204Pb ! 158! Figure 2-6 10,000 A. B. CO2/Nb = 600 Fernandina MI CO2/Nb = CO2/Nb = 300 1,000 400 CO2 (ppm) Fernandina MI 100 10 Santiago MI Santiago MI 1 5 6 7 8 9 10 11 0.1 1 10 100 MgO (wt%) Nb C. D. H2O/Ce = 330 1.2 Fernandina MI Fernandina MI H2O (wt%) 1.0 0.8 0.6 0.4 Santiago MI H2O/Ce = 100 0.2 Santiago MI 5 6 7 8 9 10 11 10 20 30 40 50 MgO (wt%) Ce E. F. F/Nd = 27 800 700 Fernandina MI 600 Fernandina MI F (ppm) 500 HHe Group 400 ITE Group ITD Group 300 F/Nd = 15.5 WD Group Santiago MI Wolf Volcano 200 100 Santiago MI 5 6 7 8 9 10 11 10 20 30 MgO (wt%) Nd 1,000 G. H. Cl/K = Fernandina MI 0.77 100 Cl (ppm) Fernandina MI 10 Santiago MI Santiago MI Cl/K = 0.02 1 5 6 7 8 9 10 11 100 1,000 10,000 MgO (wt%) K I. J. S/Dy = 300 2,000 Fernandina MI Fernandina MI S (ppm) 1,500 1,000 S/Dy = 130 500 Santiago MI Santiago MI 5 6 7 8 9 10 11 2 4 6 8 MgO (wt%) Dy ! 159! Figure 2-7 200 A. XH2O=10% XH2O=20% XH2O= 30% XH2O= 40% 150 400 CO2(ppm) bars 100 300 bars 50 200 bars 100 bars 0 0.0 0.5 1.0 1.5 H2O (wt%) 400 B. Fernandina melt inclusions 350 continue to higher pressures 300 250 S/Dy 200 150 Degassing 100 50 0 200 400 600 800 1,000 1,200 1,400 1,600 P saturation (bars) 800 C. 700 EPR 600 P saturation (bars) Samoa 500 GSC HHe Group ITE Group ITD Group 400 WD Group Society Wolf Volcano Fernandina melt inclusions 300 Pitcairn Santiago melt inclusions Open system degassing Closed system degassing 200 100 Loihi 0 0 100 200 300 400 500 P collection (bars) ! 160! Figure 2-8 7.9 2σ Sulfide Saturation 7.7 7.5 HHe Group ITE Group ITD Group 7.3 WD Group Wolf Volcano ln(S) 7.1 6.9 6.7 6.5 7.2 7.3 7.4 7.5 7.6 7.7 7.8 ln(SCSS) Figure 8. Log values of measured sulfur concentrations versus log of calculated S concentration at sulfide saturation (SCSS) using the model of Liu et al. (2007). The calculation is based on the major element composition, pressure, temperature, and the concentration of water of a sample (see text for in depth discussion of model). Black solid line is the 1:1 line at and above which samples are considered sulfide saturated. Black dashed lines represent the 200 ppm error of the model based on the experiments of Moune et al. (2009). Based on this model, the glasses range from sulfide saturated to undersaturated. Included error represents the average 2σ standard error of S measurement. ! 161! Figure 2-9 1,000 A. HHe Group F. ITE Group GSC GSC ITD Group 100 WD Group Wolf Volcano CO2/Nb 10 1 0.1 B. G. 400 350 300 H2O/Ce 250 200 150 GSC GSC C. GSC GSC H. 24 22 F/Nd 20 18 16 1 D. I. GSC GSC Cl/K 0.1 0.01 E. J. 300 GSC GSC 250 S/Dy 200 150 0.7026 0.7028 0.7030 0.7032 0.7034 0.5129 0.5130 0.5131 87Sr/86Sr 143Nd/144Nd Figure 9. Volatile/refractory element ratios versus (A-F) 87Sr/86Sr ratios and (G-L) 143Nd/144Nd ratios for individual glasses. Also included are fields representing previously reported data from the GSC (Ingle et al., 2010). Errors are 2σ standard errors. ! 162! Figure 2-10 0.16 ITE HHe 0.14 0.8 1.6% 0.12 0.6 50%NaCl 0.10 2 r = 0.51 Cl/K 0.8% Cl/K 0.4 1.5% Expected mantle limits 0.08 15%NaCl 0.2 0.06 SW 1.3% 0.0 0.04 100 200 300 400 H2O/Ce 0.02 150 200 250 300 350 H2O/Ce Figure 10. Cl/K ratios versus H2O/Ce ratios of the ITE and HHe Group lavas. Black solid line shows calculated York regression. The relatively good linear correlation shows that assimilation of hydrothermally altered material might be simultaneously affecting Cl and H2O contents of the glasses. Dashed box shows the expected mantle limits of a trace element enriched source (Dixon et al., 2002, Simons et al., 2002, Stroncik & Haase, 2004, Workman et al., 2006). Inset shows Cl/K ratios versus H2O/Ce ratios for all the glasses of this study. Lines show compositional effect of adding seawater (SW) (Cl = 1.935 wt%, H2O = 97.5 wt%, K = 33.2 ppm, and Ce = 0.001 ppb), a 15% NaCl brine (Cl = 9.9 wt%, H2O = 85 wt%, K = 2075 ppm, and Ce = 0.006 ppb), and a 50% NaCl brine (Cl = 30.3 wt%, H2O = 50 wt%, K = 6474 ppm, and Ce = 0.02 ppb) (Kent et al., 1999a, Kent et al., 1999b) to the composition of sample D20a. Ticks are labeled for the percent material mixed. Errors are 2σ standard errors. ! 163! Figure 2-11 1,000 A. 100 CO2/Nb 10 1 0.1 B. 350 H2O/Ce 250 150 C. 24 HHe Group ITE Group F/Nd 21 ITD Group WD Group Wolf Volcano Group 18 1 D. Cl/K 0.1 0.01 E. 300 250 S/Dy 200 150 0.75 1.00 1.25 1.50 1.75 Sr/Sr* Fig. 11. Volatile/refractory element ratios versus Sr/Sr* for individual glasses. Errors are 2σ standard errors. ! 164! Figure 2-12 A. HHe Group ITE Group 4 ITD Group WD Group 2 Δ7/4 0 -2 B. 25 23 F/Nd 21 19 17 10 15 20 25 3He/4He Figure 12. 3He/4He ratios versus A.) Δ7/4 and B.) F/Nd ratios. Δ7/4 calculations are based on equations found in Hart (1984) and represent positive and negative deviations from the northern hemisphere reference line. Errors are 2σ standard errors. ! 165! Figure 2-13 1,000 A. F. MORB 100 CO2/Nb 10 HHe Group 1 ITE Group ITD Group WD Group 0.1 Wolf Volcano B. G. 400 350 300 H2O/Ce FOZO 250 200 150 MORB C. H. 25 23 FOZO F/Nd 21 19 MORB 17 1 D. I. 0.1 FOZO Cl/K MORB 0.01 0.001 E. J. 310 270 FOZO S/Dy 230 MORB 190 150 0.92 0.93 0.94 0.95 0.96 0.97 0.02 0.04 0.06 0.08 0.10 0.12 208*/206* Th/La Figure 13. Volatile/refractory element ratios versus (A-E) 208Pb*/206Pb* and (F-J) Th/La ratios for individual glasses and array end-member compositions. 208Pb*/206Pb* = (208Pb/204Pb-29.4761)/(206Pb/204Pb-9.3066). End-member compositions for the ITE, HHe, and WD groups are determined from samples filtered to Cl/K ratios ≤ 0.08 and S/Dy ratios ≥ 175 that are not sulfide saturated. Additionally, for ITE group, end-member components are the mean and median of the samples with the highest F/Nd ratios. These represent an end-member of the correlation between F/Nd and 3He/4He ratios defined by array 1 (Fig. 13). For the HHe group, because the in-group F/Nd ratio variation is lower, the end-member components are the average and median values of the filtered samples. The WD end-member is defined by the mean and median values of the samples with isotopic composition closest to the Pinta component. F/Nd, Cl/K and S/Dy ratios of the ITD end-member composition are set by the samples with the lowest Cl/K ratios. CO2/Nb ratios are set by the expected CO2/Nb ratio of a MORB source (Saal et al., 2002) and H2O/Ce end-member set by the average value of Santiago normal inclusions reported in Koleszar et al., [2009]. End-member compositions presented in Table 3 Dashed lines show the values generally cited for the FOZO and MORB sources (e.g. Dixon et al., 2002, Saal et al., 2002, Simons et al., 2002). Errors bars showing the 1 standard deviation of samples used to set end-member compositions. ! 166! Supplementary Figure 2-1 A. Wolf Vol. Fernandina 2.5 Sierra Negra/ Pinta Sm/Yb 2.0 Cerro Azul Floreana 1.5 GSC 1.0 B. Fernandina 3,000 Floreana Sierra Negra/ 2,600 Cerro Azul Ti/Gd Wolf Vol. 2,200 Pinta 1,800 GSC C. Floreana 0.25 0.20 Sierra Negra/ Cerro Azul Pinta Nb/Zr 0.15 Fernandina 0.10 Wolf Vol. 0.05 GSC 0.02 0.04 0.06 0.08 0.10 0.12 Th/La HHe Group ITE Group ITD Group WD Group Wolf Volcano Supplementary Figure 1. Th/La ratios versus A.) Sm/Yb ratios, B.) Ti/Gd ratios and C.) Nb/Zr ratios of individual glasses plotted with fields defining the ranges of previously measured whole rock and glass samples from the Galapagos Archipelago and the Galapagos ! 167! Spreading Center (GSC) (references as in Fig. 4). Notice that glasses of this study cover almost the entire geochemical range Supplementary Figure 2-2 A. Sierra Negra/Cerro Azul E. Fernandina I. Sierra Negra/Cerro Azul Fernandina Fernandina Pinta 2.5 Sierra Negra/Cerro Azul Pinta 2.0 Floreana Sm/Yb Pinta Wolf/Darwin Is. Wolf/Darwin Is. 1.5 Genovesa Genovesa Genovesa Floreana Floreana 1.0 Wolf/Darwin Is. B. Fernandina F. Fernandina HHe Group J. Pinta 0.12 Pinta ITE Group Pinta ITD Group 0.10 WD Group Fernandina Wolf Volcano Floreana 0.08 Th/La 0.06 Floreana Floreana Sierra Negra/ Sierra Negra/ 0.04 Sierra Negra/ Cerro Azul Cerro Azul GSC Cerro Azul 0.02 Fernandina C. Fernandina G. Fernandina K. 3,000 Floreana Sierra Negra/ Sierra Negra/Cerro Azul Cerro Azul Floreana 2,600 Pinta Genovesa Floreana Genovesa Ti/Gd Genovesa 2,200 Pinta Pinta 1,800 GSC GSC 1,400 D. Sierra Negra/ Floreana H. Floreana L. Floreana Cerro Azul Sierra Negra/ 0.16 Pinta Sierra Negra/Cerro Azul Cerro Azul Fernandina Fernandina 0.12 Pinta Pinta Nb/Zr Fernandina 0.08 GSC GSC GSC 0.04 0.7026 0.7030 0.7034 0.5129 0.5130 0.5131 0.92 0.94 0.96 87Sr/86Sr 143Nd/144Nd 208*/206* Supplementary Figure 2. (A-D) 87Sr/86Sr, (E-H) 143Nd/144Nd, and (I-L) 208Pb*/206Pb* versus Sm/Yb, Ti/Gd, Th/La, and Nb/Zr ratios of individual glasses with a field showing the range of data collected from the GSC and for selected islands from the Galapagos Archipelago. 208Pb*/206Pb* = (208Pb/204Pb-29.4761)/(206Pb/204Pb-9.3066) is the ratio of the radiogenic additions to the initial terrestrial Pb. Trace element ratios are correlated with isotopic ratios indicating that the trace element variations are long-lived characteristics of the mantle sources. Notice 2 distinct arrays of data; the first array being composed of the HHe and ITE Groups and the second one being composed of the WD and ITD groups. References for fields and isotopic data as in Figure 4. Errors are 2σ standard error. ! 168! Supplementary Figure 2-3 A. HHe Group ITE Group ITD Group WD Group 25 Fernandina 20 3He/4He Sierra Negra/Cerro Azul 15 Floreana 10 GSC Pinta Pinta 1,500 2,000 2,500 3,000 Ti/Gd ! 169! TABLES Table 1. Sample Information and chemical data for Galapagos glasses Group HHe Group HHe HHe HHe HHe HHe HHe HHe HHe Sample AHA18a AHA18b AHA19a AHA19b AHA20a AHA20b AHA21a AHA21b AHA22a AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- Cruise 2 2 2 2 2 2 2 2 2 Latitudea -0.17 -0.17 -0.21 -0.21 -0.24 -0.24 -0.25 -0.25 -0.19 Longitudea -91.77 -91.77 -91.81 -91.80 -91.75 -91.75 -91.75 -91.75 -91.76 Average Depth (m) 2500 2500 2800 2800 1740 1740 1670 1670 2300 Major Elements (wt%) SiO2 49.11 49.22 49.01 49.42 48.76 48.74 48.80 48.76 48.56 TiO2 2.90 3.10 2.92 3.14 3.57 3.70 3.61 3.64 3.13 Al2O3 14.04 14.09 13.92 13.92 13.88 13.78 13.82 13.81 14.64 FeO* 11.38 11.37 11.74 11.77 11.83 11.86 11.76 11.80 11.24 MnO 0.21 0.20 0.21 0.20 0.23 0.18 0.21 0.20 0.17 MgO 6.53 6.75 6.31 6.47 6.37 6.33 6.45 6.46 6.94 CaO 11.88 11.78 11.90 11.46 11.40 11.32 11.33 11.30 11.66 Na2O 2.55 2.75 2.55 2.82 3.09 3.14 3.07 3.11 2.84 K 2O 0.43 0.44 0.46 0.46 0.52 0.54 0.53 0.53 0.46 P 2O 5 0.30 0.31 0.33 0.33 0.36 0.42 0.40 0.39 0.33 Total 99.31 98.15 99.35 98.62 97.85 98.22 98.12 98.26 98.00 Volatile Content H2O (wt%) 0.593 0.683 0.635 0.732 0.834 0.779 0.738 0.761 0.683 CO2 (ppm) 114.42 98.90 122.10 105.31 77.97 67.73 68.04 72.88 94.02 F (ppm) 424.55 415.20 440.48 441.45 555.83 523.26 513.49 531.55 456.63 S (ppm) 1135.16 1260.73 1224.45 1386.20 1428.27 1395.42 1344.32 1378.54 1273.68 Cl (ppm) 139.21 183.57 213.02 284.83 255.32 227.89 230.26 240.64 206.22 P saturation (bars) 279 256 300 277 219 188 186 197 214 ! 170! Trace Elements (ppm) Rb 6.88 8.16 6.99 8.22 9.19 10.24 10.05 10.01 8.37 Ba 83.20 89.04 84.26 90.40 103.33 111.12 109.14 109.39 93.69 Th 1.21 1.19 1.27 1.20 1.60 1.49 1.46 1.46 1.25 U 0.378 0.388 0.409 0.413 0.456 0.499 0.493 0.497 0.411 Nb 20.84 20.79 21.20 21.04 25.29 27.03 26.24 26.24 21.68 K 3615.31 3452.39 3684.44 3493.89 3857.63 4278.76 4197.14 4216.56 3642.82 La 14.62 13.67 14.87 13.74 17.44 16.95 16.56 16.58 14.34 Ce 35.73 32.69 35.84 33.54 38.51 41.88 40.87 40.79 34.13 Pb 0.57 1.04 0.56 1.08 1.28 1.29 1.26 1.26 1.14 Sr 347.00 337.22 337.82 341.80 338.38 345.22 343.52 343.32 355.79 Nd 22.33 21.20 22.18 21.02 26.37 25.79 25.35 25.35 21.93 P 1373.22 1224.27 1383.01 1192.77 1482.54 1515.61 1497.55 1491.89 1356.24 Zr 172.50 168.23 170.66 163.40 229.47 207.91 202.67 202.13 174.86 Hf 4.09 4.01 4.09 3.83 5.44 4.80 4.77 4.70 4.15 Sm 5.83 5.41 5.67 5.30 6.75 6.43 6.31 6.31 5.46 Eu 1.92 1.84 1.93 1.84 2.19 2.21 2.15 2.16 1.89 Ti 17836.84 16653.09 17706.29 16537.20 19420.25 20085.96 19689.78 19696.36 16940.13 Gd 6.13 5.96 6.14 5.67 7.71 6.90 6.77 6.82 6.12 Dy 5.54 5.39 5.51 5.29 7.00 6.18 6.09 6.12 5.50 Y 28.88 28.32 28.62 27.28 36.58 32.46 31.92 31.75 28.60 Er 3.13 2.78 3.08 2.69 3.64 3.15 3.11 3.13 2.84 Yb 2.68 2.39 2.58 2.31 3.07 2.75 2.69 2.65 2.44 Pr 4.63 4.66 5.62 5.76 5.62 5.62 4.79 Tb 0.93 0.89 1.19 1.08 1.05 1.06 0.95 Ho 1.07 1.03 1.38 1.21 1.20 1.19 1.07 Tm 0.38 0.37 0.50 0.44 0.43 0.43 0.40 Lu 0.37 0.34 0.48 0.42 0.41 0.40 0.37 Isotopes 3 He/4Heb 24.79 24.68 22.62 22.20 22.79 22.55 22.59 22.61 23.07 1σ 0.14 0.12 0.13 0.08 0.14 0.12 0.26 0.10 0.10 C/3He 4.27E+09 4.26E+09 4.61E+11 4.06E+10 7.92E+09 3.39E+09 5.48E+09 6.09E+09 5.97E+09 87 Sr/86Src 0.70324 0.70323 0.70331 0.70324 0.70324 0.70329 ! 171! 143 Nd/144Ndc 0.512933 0.512944 0.512965 0.51294 εNd 5.76 5.97 6.38 5.89 206 Pb/204Pbc 19.082 19.067 19.039 19.055 19.079 19.097 207 Pb/204Pbc 15.560 15.550 15.532 15.551 15.554 15.556 208 Pb/204Pbc 38.702 38.703 38.622 38.697 38.723 38.734 Group HHe HHe HHe HHe HHe HHe HHe HHe HHe Sample AHA22b AHA23a AHA23b AHA24b AHA25a AHA25c AHA26a AHA27a AHA27c AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- Cruise 2 2 2 2 2 2 2 2 2 Latitudea -0.18 -0.26 -0.26 -0.45 -0.45 -0.45 -0.43 -0.42 -0.42 Longitudea -91.76 -91.73 -91.73 -91.78 -91.74 -91.74 -91.71 -91.67 -91.67 Average Depth (m) 2300 1395 1395 3115 2750 2750 1890 950 950 Major Elements (wt%) SiO2 48.57 49.28 49.27 48.21 48.99 49.07 49.04 49.31 49.40 TiO2 2.93 3.27 3.24 4.05 3.45 3.33 3.16 3.72 3.80 Al2O3 14.74 13.85 13.92 13.86 13.98 14.10 14.11 13.68 13.59 FeO* 11.27 11.43 11.32 12.08 11.65 11.98 11.65 12.32 12.39 MnO 0.18 0.19 0.21 0.16 0.21 0.19 0.19 0.22 0.21 MgO 6.59 6.77 6.76 6.12 6.37 6.03 6.38 5.92 5.81 CaO 11.67 11.66 11.65 11.21 11.60 11.23 11.70 10.95 10.83 Na2O 2.66 2.78 2.92 3.38 2.93 2.72 2.65 2.97 3.06 K 2O 0.47 0.42 0.41 0.55 0.45 0.44 0.45 0.52 0.53 P 2O 5 0.32 0.33 0.31 0.38 0.34 0.33 0.34 0.37 0.38 Total 99.41 98.25 98.50 97.93 98.32 99.40 99.66 98.34 98.40 Volatile Content H2O (wt%) 0.571 0.755 0.760 0.875 0.843 0.749 0.566 0.827 0.830 CO2 (ppm) 109.09 61.73 36.23 137.23 101.18 109.91 83.29 13.04 15.38 F (ppm) 444.55 439.06 425.69 561.14 458.57 475.93 437.88 524.43 527.79 S (ppm) 1096.73 1401.60 1337.36 1805.47 1404.12 1296.13 1177.93 1152.91 1306.23 Cl (ppm) 153.10 197.77 189.76 345.18 239.38 251.60 147.90 229.66 239.87 P saturation (bars) 228 187 132 294 286 289 209 93 98 ! 172! Trace Elements (ppm) Rb 7.16 7.77 7.59 9.31 8.79 7.99 7.52 10.28 10.33 Ba 86.92 84.18 82.67 99.81 91.78 86.15 82.39 107.05 107.77 Th 1.25 1.24 1.37 1.25 1.20 1.25 1.19 1.43 1.49 U 0.406 0.398 0.384 0.451 0.432 0.425 0.370 0.487 0.512 Nb 21.28 22.07 22.00 23.97 23.52 24.24 22.59 26.69 27.24 K 3814.26 3278.95 3187.74 4375.60 3661.56 3781.60 3625.93 4282.16 4294.73 La 15.16 13.69 14.43 15.84 13.90 15.60 14.98 16.33 16.72 Ce 36.66 32.41 31.85 40.03 35.16 38.12 35.94 39.97 41.24 Pb 0.51 1.00 1.01 1.44 1.18 0.50 0.58 1.31 1.35 Sr 368.06 331.05 319.62 420.13 344.38 363.97 355.44 345.56 342.83 Nd 22.58 21.20 22.15 26.00 21.67 23.94 22.64 25.15 25.80 P 1446.11 1188.44 1303.16 1600.48 1481.84 1559.53 1469.68 1629.50 1627.12 Zr 177.64 168.43 188.25 210.80 166.34 186.18 177.33 195.96 202.76 Hf 4.14 4.07 4.64 4.82 3.93 4.39 4.13 4.58 4.70 Sm 5.79 5.41 5.83 6.58 5.39 6.08 5.77 6.12 6.44 Eu 1.89 1.85 1.89 2.28 1.89 2.04 2.01 2.14 2.22 Ti 18009.30 17577.94 17782.20 22223.24 18935.84 20355.43 19296.46 20288.15 20827.68 Gd 6.17 6.08 6.77 6.92 5.89 6.51 6.35 6.64 6.91 Dy 5.52 5.55 6.23 6.17 5.34 6.01 5.79 6.07 6.29 Y 28.82 28.97 32.64 32.04 27.90 30.98 29.60 31.47 32.58 Er 3.07 2.80 3.27 3.14 2.75 3.33 3.22 3.09 3.19 Yb 2.62 2.39 2.71 2.70 2.43 2.84 2.76 2.72 2.73 Pr 4.58 4.68 5.71 4.81 5.57 5.72 Tb 0.95 1.06 1.08 0.92 1.05 1.08 Ho 1.08 1.24 1.20 1.05 1.18 1.23 Tm 0.39 0.44 0.43 0.38 0.42 0.44 Lu 0.37 0.42 0.41 0.37 0.41 0.41 Isotopes 3 He/4Heb 22.45 27.94 28.82 27.11 27.44 27.24 27.25 26.23 1σ 0.10 0.17 0.40 0.11 0.12 0.11 0.09 0.11 C/3He 6.57E+09 2.36E+10 1.53E+10 6.62E+10 6.21E+10 5.79E+09 2.08E+10 3.25E+10 87 Sr/86Src 0.70324 0.70327 0.70330 0.70319 0.70331 0.70330 0.70327 ! 173! 143 Nd/144Ndc 0.512915 0.512889 0.51294 0.512944 0.512936 0.512928 εNd 5.40 4.90 5.89 5.97 5.81 5.66 206 Pb/204Pbc 19.055 19.092 19.088 18.949 19.046 19.046 19.070 207 Pb/204Pbc 15.551 15.545 15.550 15.533 15.532 15.542 15.548 208 Pb/204Pbc 38.699 38.734 38.741 38.485 38.628 38.668 38.718 Group HHe HHe HHe HHe HHe HHe HHe HHe HHe Sample AHA29a AHA29b AHA30a AHA30b AHA32a AHA32b D11a D24a D24a-2 AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- AHA-Nemo- Cruise 2 2 2 2 2 2 DRIFT04 DRIFT04 DRIFT04 Latitudea -0.54 -0.54 -0.54 -0.54 -0.56 -0.56 -0.08 -0.33 -0.33 Longitudea -91.65 -91.65 -91.68 -91.68 -91.71 -91.71 -91.79 -91.76 -91.75 Average Depth (m) 1325 1325 1852.5 1852.5 2495 2495 2842 2758 2758 Major Elements (wt%) SiO2 49.05 48.88 49.14 49.35 49.30 48.97 49.16 49.38 49.02 TiO2 3.95 3.90 3.01 3.04 2.82 3.16 3.06 2.88 3.05 Al2O3 13.60 13.70 14.10 14.05 14.95 14.41 14.05 14.21 14.23 FeO* 12.66 12.75 11.34 11.18 10.87 11.12 11.59 11.36 11.45 MnO 0.19 0.22 0.19 0.20 0.19 0.18 0.19 0.17 0.19 MgO 5.74 5.79 6.82 6.80 6.60 6.70 6.41 6.32 6.58 CaO 10.77 10.83 11.91 11.91 11.52 11.81 11.72 11.74 11.89 Na2O 3.10 3.01 2.75 2.71 2.57 2.81 2.62 2.65 2.88 K 2O 0.53 0.53 0.43 0.43 0.44 0.48 0.39 0.41 0.40 P 2O 5 0.38 0.41 0.30 0.32 0.29 0.35 0.31 0.32 0.30 Total 98.48 98.04 97.73 98.25 99.55 97.72 99.50 99.44 97.71 Volatile Content H2O (wt%) 0.918 1.011 0.614 0.616 0.714 0.900 0.622 0.884 0.966 CO2 (ppm) 37.37 28.73 74.42 78.70 108.30 93.23 147.93 108.96 89.42 F (ppm) 547.42 549.83 404.57 396.94 420.39 456.01 429.53 406.07 404.18 S (ppm) 1501.25 1479.85 1291.11 1240.35 993.33 1214.59 1326.72 1351.46 1528.25 Cl (ppm) 243.37 244.60 179.63 172.18 207.21 326.58 206.20 280.84 372.35 P saturation (bars) 135 161 158 195 281 205 354 283 ! 174! Trace Elements (ppm) Rb 10.22 9.79 7.43 7.66 6.39 8.36 6.90 6.60 7.02 Ba 110.16 106.22 83.62 84.83 81.38 92.61 74.86 71.05 83.49 Th 1.72 1.70 1.27 1.17 1.18 1.25 1.27 1.06 1.08 U 0.499 0.485 0.381 0.389 0.367 0.403 0.379 0.344 0.369 Nb 28.11 27.17 20.36 20.42 20.74 21.96 21.26 19.90 21.64 K 4150.04 4005.53 3229.27 3336.98 3600.35 3572.23 3358.00 3194.48 3154.62 La 18.39 18.26 14.03 13.28 14.15 14.33 13.67 12.93 13.12 Ce 41.85 40.22 31.66 32.25 34.17 34.21 33.14 31.78 33.17 Pb 1.27 1.26 1.00 1.02 0.52 1.15 1.11 0.66 0.97 Sr 338.39 339.49 334.45 328.11 357.89 347.79 310.76 337.77 348.76 Nd 28.03 27.74 21.46 20.55 21.22 21.89 20.98 20.11 20.70 P 1562.19 1493.66 1153.63 1179.49 1387.56 1247.62 1474.85 1343.55 1252.86 Zr 245.44 247.45 180.61 160.17 166.73 176.83 171.18 158.20 156.50 Hf 5.70 5.82 4.42 3.83 3.91 4.11 3.98 3.69 3.79 Sm 7.13 7.19 5.58 5.17 5.39 5.54 5.21 4.91 5.21 Eu 2.34 2.29 1.87 1.81 1.94 1.91 1.79 1.73 1.83 Ti 21325.69 20703.51 16463.21 16528.05 17638.06 16882.28 18513.67 17545.56 17630.86 Gd 8.29 8.23 6.45 5.71 5.86 6.15 5.79 5.38 5.66 Dy 7.48 7.57 5.97 5.21 5.36 5.57 5.52 5.02 5.20 Y 40.31 40.76 31.28 27.35 27.47 29.18 29.13 26.70 26.37 Er 3.95 3.99 3.08 2.67 3.02 2.82 3.04 2.90 2.62 Yb 3.38 3.44 2.61 2.34 2.47 2.43 2.82 2.32 2.28 Pr 6.01 5.86 4.60 4.49 4.81 4.57 Tb 1.29 1.29 1.00 0.91 0.97 0.89 Ho 1.50 1.50 1.16 1.02 1.09 0.99 Tm 0.54 0.56 0.42 0.37 0.39 0.36 Lu 0.53 0.53 0.41 0.35 0.38 0.33 Isotopes 3 He/4Heb 25.48 26.31 22.19 22.99 23.60 23.83 25.67 25.05 25.05 1σ 0.10 0.14 0.08 0.12 0.11 0.10 0.14 0.13 0.13 C/3He 1.11E+10 5.00E+09 1.23E+10 1.08E+10 5.97E+09 4.04E+09 2.95E+10 6.66E+10 ! 175! 87 Sr/86Src 0.70325 0.70324 0.70326 0.70352 143 Nd/144Ndc 0.512963 0.512924 εNd 6.34 5.58 206 Pb/204Pbc 19.063 19.057 19.057 19.052 207 Pb/204Pbc 15.549 15.550 15.550 15.545 208 Pb/204Pbc 38.712 38.699 38.697 38.702 Group HHe HHe HHe HHe ITE Group ITE ITE ITE ITE Sample D24b D28c D9a RC5 AHA23c AHA24a D13b D14a D1a AHA-Nemo- AHA-Nemo- Cruise DRIFT04 DRIFT04 DRIFT04 DRIFT04 2 2 DRIFT04 DRIFT04 DRIFT04 Latitudea -0.33 -0.44 -0.11 -0.36 -0.26 -0.45 -0.19 -0.17 Longitudea -91.75 -91.65 -91.84 -91.79 -91.73 -91.78 -91.55 -91.70 Average Depth (m) 2758 1033.5 3087 3239 1395 3115 2053 2454 Major Elements (wt%) SiO2 49.04 48.92 49.32 48.78 49.06 48.33 48.42 48.89 47.79 TiO2 3.13 2.84 3.04 3.36 3.39 3.78 3.18 3.71 3.53 Al2O3 14.25 14.91 14.44 14.06 14.07 13.97 15.45 14.11 15.41 FeO* 11.60 10.46 11.72 12.01 11.93 12.10 11.41 12.56 11.51 MnO 0.20 0.20 0.17 0.18 0.20 0.20 0.18 0.21 0.18 MgO 6.44 7.44 5.98 6.08 6.17 5.79 5.92 5.20 5.94 CaO 11.75 12.01 11.33 10.81 11.42 11.24 10.40 10.55 10.55 Na2O 2.85 2.53 2.71 2.89 2.92 3.06 3.26 3.14 3.70 K 2O 0.41 0.38 0.44 0.51 0.47 0.56 0.70 0.65 0.88 P 2O 5 0.32 0.30 0.31 0.73 0.37 0.36 0.40 0.43 0.50 Total 97.53 97.96 99.47 99.41 97.89 99.37 99.31 99.45 98.53 Volatile Content H2O (wt%) 1.072 0.595 0.884 0.810 1.062 0.677 0.956 1.150 1.073 CO2 (ppm) 82.08 46.81 136.57 156.63 19.86 151.17 64.24 55.15 104.99 F (ppm) 419.37 375.64 463.63 530.53 487.93 539.31 621.87 685.65 715.30 S (ppm) 1559.25 884.77 1313.13 1401.56 1163.58 1470.06 1304.01 1580.81 1510.79 Cl (ppm) 453.15 208.03 284.38 187.88 203.38 246.00 239.00 240.35 875.38 ! 176! P saturation (bars) 283 133 369 369 152 294 244 Trace Elements (ppm) Rb 6.74 6.94 7.28 8.09 8.74 8.98 8.90 10.69 18.57 Ba 82.11 81.48 84.81 93.16 92.96 93.38 132.05 129.05 198.16 Th 1.14 1.07 1.26 1.34 1.32 1.26 1.73 1.82 2.08 U 0.353 0.357 0.379 0.431 0.447 0.423 0.542 0.576 0.713 Nb 21.63 20.22 21.53 24.73 23.95 23.78 26.59 31.80 36.22 K 3049.36 3067.02 3754.05 4193.80 3672.79 4636.01 5455.51 5408.79 7545.74 La 13.33 12.37 14.90 16.19 15.03 17.18 19.16 21.39 22.34 Ce 32.66 31.80 35.68 39.04 36.87 43.21 44.42 50.60 54.86 Pb 0.94 0.89 0.90 0.94 1.17 0.51 1.23 1.34 1.79 Sr 335.59 340.16 346.73 339.73 326.84 445.13 391.07 392.77 430.27 Nd 20.92 19.54 22.54 24.09 23.12 27.80 26.55 30.36 30.33 P 1181.65 1191.46 1489.46 1741.10 1404.72 1679.66 1946.07 2152.53 2189.31 Zr 161.91 145.20 180.28 202.15 180.89 223.71 226.58 255.16 237.26 Hf 3.94 3.60 4.16 4.40 4.25 5.17 4.86 5.52 5.40 Sm 5.26 4.83 5.55 5.91 5.82 6.87 6.35 7.74 7.14 Eu 1.85 1.74 1.92 1.98 1.99 2.34 2.24 2.52 2.49 Ti 17502.83 16260.21 18475.01 20273.80 18277.50 23511.46 19216.05 23356.17 20119.64 Gd 5.88 5.22 6.22 6.50 6.24 7.55 7.04 8.01 7.34 Dy 5.29 4.76 5.63 5.82 5.67 6.70 6.31 6.77 6.35 Y 27.21 23.57 29.40 31.15 29.21 33.55 32.50 36.87 32.12 Er 2.70 2.35 3.02 3.27 2.91 3.62 3.30 3.91 3.16 Yb 2.38 2.05 2.62 2.82 2.55 3.09 2.91 3.21 2.77 Pr 4.58 4.33 5.12 7.01 Tb 0.91 0.80 0.97 1.11 Ho 1.02 0.89 1.11 1.20 Tm 0.37 0.32 0.40 0.42 Lu 0.35 0.30 0.38 0.40 Isotopes 3 He/4Heb 25.71 27.71 25.94 23.47 18.42 11.84 9.50 1σ 0.11 0.09 0.18 0.12 0.08 0.04 0.10 C/3He 9.05E+09 3.42E+11 ! 177! 87 Sr/86Src 0.70323 0.70322 143 Nd/144Ndc 0.512946 0.512931 0.512944 εNd 6.01 5.72 5.97 206 Pb/204Pbc 19.089 19.103 19.066 207 Pb/204Pbc 15.550 15.555 15.554 208 Pb/204Pbc 38.723 38.767 38.714 Group ITE ITE ITE ITE ITE ITE ITE ITE ITE Sample D20a D27a D31a D34a D36 D38a D39a D40a D42a Cruise DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 Latitudea -0.27 -0.57 -0.66 -0.79 -0.76 -0.72 -0.78 -0.89 -0.99 Longitudea -91.89 -91.74 -91.64 -91.67 -91.62 -91.55 -91.58 -91.68 -91.61 Average Depth (m) 3375 3078 2845 3049 2645 2250 2218 3097 2947 Major Elements (wt%) SiO2 49.29 47.99 48.54 48.80 49.48 48.94 49.10 48.74 49.08 TiO2 3.32 3.14 4.39 2.23 3.54 2.30 2.44 3.25 2.50 Al2O3 13.79 15.53 12.92 15.26 13.32 14.67 14.50 14.25 14.57 FeO* 12.26 11.50 14.69 10.44 13.44 10.28 10.72 12.55 11.07 MnO 0.21 0.19 0.23 0.18 0.23 0.17 0.18 0.22 0.18 MgO 6.00 5.49 5.28 6.42 5.13 7.06 6.93 5.14 6.62 CaO 10.91 10.25 10.19 12.51 10.27 12.73 12.57 10.45 12.36 Na2O 2.79 3.35 2.95 2.81 3.06 2.54 2.45 3.41 2.57 K 2O 0.58 1.01 0.61 0.47 0.64 0.46 0.41 0.78 0.43 P 2O 5 0.38 0.86 0.46 0.50 0.43 0.28 0.28 0.43 0.32 Total 99.53 99.32 100.24 99.61 99.53 99.42 99.58 99.22 99.69 Volatile Content H2O (wt%) 0.498 1.049 0.572 0.622 0.727 0.504 0.498 1.015 0.713 CO2 (ppm) 183.25 117.22 149.96 160.89 111.27 125.64 108.41 137.82 137.50 F (ppm) 509.74 795.56 648.49 455.49 681.26 437.67 406.96 806.52 434.30 S (ppm) 947.55 1426.94 1432.22 1108.79 1475.36 1058.34 1038.08 1599.16 1456.49 Cl (ppm) 165.85 309.19 260.44 240.03 219.99 208.37 185.67 361.91 171.67 ! 178! P saturation (bars) 416 286 298 353 289 293 256 367 343 Trace Elements (ppm) Rb 9.33 13.21 12.29 7.66 12.27 6.12 6.51 13.90 6.87 Ba 101.14 160.23 117.19 100.33 117.55 80.81 76.83 164.43 80.13 Th 1.49 1.60 1.73 1.27 2.00 1.26 1.20 2.12 1.27 U 0.457 0.584 0.559 0.351 0.534 0.417 0.359 0.616 0.398 Nb 25.54 31.32 35.00 19.32 33.70 20.39 20.47 36.08 21.35 K 4456.29 8502.93 5219.53 4098.08 5223.79 3562.81 3302.41 6693.08 3491.39 La 16.97 22.63 22.06 13.84 21.49 13.48 14.03 25.08 14.61 Ce 40.45 52.97 52.49 31.09 49.70 30.49 32.71 55.69 33.81 Pb 1.79 0.86 1.04 0.59 1.12 2.30 0.62 0.95 0.63 Sr 332.95 522.03 364.32 344.54 304.27 300.47 296.04 389.83 311.56 Nd 24.70 31.01 32.21 18.07 28.69 17.34 19.69 30.47 20.48 P 1676.40 2340.49 2129.75 1121.90 1890.65 1143.74 1115.05 1873.33 1196.60 Zr 201.97 241.82 254.47 134.31 237.51 133.23 143.48 224.55 147.81 Hf 4.54 5.12 5.97 3.26 5.45 3.20 3.44 5.15 3.56 Sm 6.44 7.22 7.88 4.54 7.25 4.34 5.02 7.37 5.04 Eu 2.10 2.59 2.63 1.62 2.44 1.42 1.67 2.59 1.79 Ti 19890.15 18989.62 27116.31 13506.32 21618.82 13183.85 14478.65 19832.54 14956.26 Gd 6.86 7.46 8.62 4.92 7.88 4.58 5.13 7.64 5.42 Dy 6.13 6.07 7.79 4.40 7.04 4.08 4.78 6.69 4.99 Y 32.15 30.83 40.26 23.37 37.20 21.72 24.53 34.13 25.09 Er 3.44 3.22 4.35 2.42 4.00 2.20 2.63 3.59 2.67 Yb 3.00 2.63 3.68 2.13 3.18 1.81 2.24 3.09 2.27 Pr Tb Ho Tm Lu Isotopes 3 He/4Heb 10.52 13.94 12.83 14.97 18.19 13.13 17.46 1σ 0.05 0.06 0.11 0.11 0.12 0.20 0.12 C/3He ! 179! 87 Sr/86Src 0.70333 0.70336 0.70339 0.70338 143 Nd/144Ndc 0.512945 0.512933 0.512923 0.512945 εNd 5.99 5.76 5.56 5.99 206 Pb/204Pbc 19.151 19.482 19.438 19.409 19.345 207 Pb/204Pbc 15.566 15.599 15.593 15.591 15.577 208 Pb/204Pbc 38.817 39.150 39.100 39.092 38.993 Group ITE ITE ITE ITE ITE ITE ITE ITE ITE Sample D47a D48a D49a D51a D59a D5b D5g D60a D64c Cruise DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 Latitudea -1.08 -1.13 -1.29 -1.23 -1.22 0.10 0.10 -1.29 -1.47 Longitudea -91.55 -91.45 -91.19 -91.16 -91.01 -91.64 -91.64 -90.99 -90.74 Average Depth (m) 3098 3244 3504 2810 1897 3006 3006 3117 3358 Major Elements (wt%) SiO2 48.95 48.85 48.92 48.99 49.22 49.32 49.30 50.23 50.18 TiO2 2.32 2.41 3.33 3.48 3.02 2.82 3.02 2.97 2.91 Al2O3 14.75 14.52 13.52 13.47 13.77 14.74 14.64 14.08 14.65 FeO* 10.84 10.70 13.72 13.38 12.50 11.21 11.20 11.96 12.43 MnO 0.18 0.18 0.22 0.23 0.23 0.19 0.19 0.19 0.21 MgO 6.65 6.73 5.03 5.18 5.89 5.85 6.16 5.94 5.08 CaO 12.31 12.42 10.20 10.29 11.24 11.34 11.35 10.86 10.25 Na2O 2.70 2.50 3.09 3.03 2.75 2.98 3.17 2.84 2.95 K 2O 0.46 0.47 0.65 0.72 0.50 0.62 0.63 0.53 0.55 P 2O 5 0.28 0.54 0.70 0.66 0.41 0.33 0.35 0.38 0.35 Total 99.44 99.31 99.39 99.43 99.52 99.41 98.09 99.98 99.55 Volatile Content H2O (wt%) 0.767 0.645 0.848 0.787 0.549 1.040 1.153 0.526 1.086 CO2 (ppm) 160.91 164.06 188.12 118.83 84.61 127.18 115.49 134.18 85.94 F (ppm) 461.41 448.90 638.23 712.34 534.42 530.76 538.58 501.92 569.68 S (ppm) 1318.62 1328.72 1179.33 1586.92 1355.26 1344.41 1480.04 1425.46 1309.29 Cl (ppm) 261.29 276.12 459.70 267.75 183.22 346.54 421.44 198.45 433.35 ! 180! P saturation (bars) 401 369 472 314 210 379 230 314 301 Trace Elements (ppm) Rb 7.92 7.81 11.96 11.87 10.66 7.48 10.88 9.45 9.87 Ba 97.75 89.06 121.45 128.71 104.98 115.03 135.09 93.08 99.65 Th 1.35 1.37 1.75 1.99 1.46 1.44 1.43 1.58 1.79 U 0.399 0.387 0.570 0.592 0.457 0.430 0.452 0.499 0.503 Nb 20.58 21.38 30.33 33.71 26.61 22.46 23.88 25.60 26.15 K 3826.29 3759.21 5201.98 5437.47 4307.37 4716.95 4757.21 4346.69 4583.43 La 14.98 14.63 19.98 21.21 17.75 15.27 15.75 17.38 18.12 Ce 34.15 33.72 46.42 49.21 41.56 34.74 38.64 40.79 42.62 Pb 0.64 0.57 0.83 0.88 1.65 1.93 1.18 2.24 1.73 Sr 317.90 306.68 317.28 305.80 308.46 347.22 367.02 320.25 305.66 Nd 20.20 19.88 26.83 28.98 24.44 20.97 22.60 24.98 25.61 P 1158.40 1194.54 1811.95 1976.02 1453.36 1568.53 1416.03 1472.68 1510.29 Zr 138.87 148.71 214.73 230.07 183.42 176.83 172.70 199.17 201.27 Hf 3.25 3.32 4.99 5.36 4.40 4.06 4.11 4.60 4.85 Sm 4.82 4.89 6.46 6.87 6.17 5.40 5.53 6.26 6.51 Eu 1.77 1.62 2.36 2.37 2.09 1.92 1.94 2.11 2.10 Ti 14045.58 14363.08 20477.95 21148.47 18392.65 16389.17 16192.90 17922.82 17477.76 Gd 5.28 5.17 7.47 7.59 7.00 5.50 5.95 6.93 6.94 Dy 5.07 4.68 6.50 7.00 6.28 4.89 5.32 6.18 6.33 Y 25.75 24.77 35.27 36.99 31.94 26.72 26.95 31.79 33.49 Er 2.82 2.57 3.72 3.78 3.54 2.80 2.59 3.34 3.48 Yb 2.39 2.26 3.13 3.32 2.91 2.21 2.31 2.92 3.02 Pr 5.14 Tb 0.93 Ho 0.99 Tm 0.35 Lu 0.34 Isotopes 3 He/4Heb 13.99 15.81 14.67 9.60 17.60 10.94 11.42 20.46 1σ 0.06 0.05 0.25 0.17 0.09 0.07 0.04 0.09 C/3He 5.29E+09 ! 181! 87 Sr/86Src 0.70337 0.70342 0.70337 0.70342 0.70346 0.70339 143 Nd/144Ndc 0.512918 0.512924 0.512907 0.512916 0.512914 0.512927 εNd 5.46 5.58 5.25 5.42 5.38 5.64 206 Pb/204Pbc 19.457 19.404 19.464 19.464 19.254 19.252 19.295 19.376 207 Pb/204Pbc 15.594 15.597 15.599 15.597 15.594 15.595 15.581 15.590 208 Pb/204Pbc 39.127 39.086 39.148 39.137 38.977 38.983 38.953 39.041 Group ITE ITE ITE ITE ITE ITE ITE ITE ITE Sample D66f D69a D6a D6g D70c D7a D8a D8c PL02-24-14 Cruise DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 DRIFT04 PLUME02 Latitudea -1.51 -1.45 0.00 0.00 -1.44 -0.05 -0.09 -0.09 -1.29 Longitudea -90.55 -90.64 -91.69 -91.69 -90.74 -91.82 -91.76 -91.76 -90.91 Average Depth (m) 1711 2425 2543 2543 2811 2987 2654 2654 2719 Major Elements (wt%) SiO2 49.28 49.65 47.90 47.98 49.17 46.95 46.87 46.76 49.46 TiO2 3.21 2.73 3.20 3.39 3.18 3.16 3.21 3.60 2.99 Al2O3 13.45 14.17 15.83 15.66 13.52 16.41 16.52 16.23 13.91 FeO* 13.24 11.70 11.06 11.17 13.08 12.25 12.15 12.48 12.17 MnO 0.22 0.18 0.18 0.17 0.23 0.20 0.19 0.20 0.21 MgO 5.46 6.29 6.10 6.33 5.34 6.22 6.08 5.84 6.04 CaO 10.70 11.05 10.69 10.58 10.67 9.85 9.43 9.60 11.13 Na2O 2.93 2.75 3.30 3.46 2.85 3.29 3.47 3.81 2.74 K 2O 0.56 0.51 0.76 0.80 0.56 0.90 0.96 0.98 0.49 P 2O 5 0.45 0.39 0.43 0.45 0.50 0.44 0.43 0.49 0.32 Total 99.50 99.41 99.44 98.16 99.10 99.67 99.30 98.75 99.45 Volatile Content H2O (wt%) 0.592 0.630 0.758 0.908 0.668 0.711 0.702 1.030 0.525 CO2 (ppm) 43.94 80.32 124.37 107.52 90.14 151.14 124.78 121.37 129.76 F (ppm) 567.09 514.68 607.96 642.58 584.30 614.01 611.09 699.35 476.06 S (ppm) 1457.48 1191.84 890.18 1026.51 1554.38 1038.59 979.87 1252.03 1183.67 Cl (ppm) 164.14 231.73 267.52 352.94 188.18 307.15 301.59 447.66 139.17 ! 182! P saturation (bars) 127 210 237 380 236 222 188 311 304 Trace Elements (ppm) Rb 10.70 9.41 10.53 12.93 11.06 14.78 14.86 18.44 9.67 Ba 103.55 93.60 148.94 166.79 98.27 184.74 187.14 208.56 104.75 Th 1.59 1.70 1.77 2.06 1.65 2.02 2.07 2.36 1.50 U 0.471 0.461 0.520 0.548 0.490 0.570 0.604 0.637 0.441 Nb 27.11 24.17 29.17 31.17 27.13 32.81 33.29 37.91 23.45 K 4545.03 4274.15 6213.40 5582.29 4515.26 7684.09 8056.29 7412.03 4116.54 La 18.45 16.63 19.89 20.97 18.24 23.17 23.38 24.31 16.69 Ce 43.12 38.78 45.54 45.81 43.26 52.95 53.81 52.80 39.20 Pb 2.08 1.75 1.69 1.36 2.03 0.41 0.51 1.63 0.28 Sr 306.77 290.69 428.98 431.82 298.42 514.04 521.44 489.77 313.88 Nd 25.97 23.42 26.81 28.83 26.39 30.09 30.87 32.39 23.82 P 1552.60 1384.11 1934.43 1604.41 1561.72 2005.62 2087.33 1929.09 1370.40 Zr 205.00 188.84 222.48 237.14 211.32 240.44 247.77 273.45 175.90 Hf 4.94 4.31 5.00 5.52 5.02 5.23 5.48 6.26 4.55 Sm 6.49 6.04 6.50 6.92 6.72 7.14 7.54 7.77 6.09 Eu 2.26 2.01 2.36 2.33 2.23 2.75 2.70 2.61 1.84 Ti 19726.31 16499.00 18880.15 18587.69 19599.45 19407.38 19873.34 20897.84 18515.70 Gd 7.26 6.47 6.94 7.49 7.29 7.63 7.79 8.10 6.61 Dy 6.81 6.19 5.78 6.54 6.85 6.29 6.59 7.07 6.01 Y 35.44 31.32 30.30 32.85 35.42 31.67 32.30 35.08 29.49 Er 3.86 3.40 3.22 3.23 3.87 3.36 3.54 3.40 3.51 Yb 3.22 2.89 2.66 2.81 3.32 2.81 2.84 2.89 2.86 Pr 6.39 7.26 Tb 1.13 1.24 Ho 1.24 1.29 Tm 0.44 0.46 Lu 0.41 0.42 Isotopes 3 He/4Heb 9.95 9.76 8.24 8.51 8.37 16.84 1σ 0.04 0.08 0.05 0.06 0.07 0.11 C/3He 3.99E+10 1.12E+10 1.06E+10 ! 183! 87 Sr/86Src 0.70337 0.70344 0.70320 0.70341 0.70322 0.70321 143 Nd/144Ndc 0.512925 0.512924 0.51292 0.512933 0.512934 εNd 5.60 5.58 5.50 5.76 5.77 206 Pb/204Pbc 19.371 19.377 19.271 19.402 19.432 19.407 207 Pb/204Pbc 15.592 15.593 15.594 15.595 15.604 15.604 208 Pb/204Pbc 39.033 39.048 38.985 39.065 39.181 39.159 Group ITE ITE ITD Group ITD ITD ITD ITD ITD ITD Sample PL02-24-32 PL02-24-37 D10a D22a D23a PL02-10-8 PL02-11-5 PL02-12-1 PL02-13-24 Cruise PLUME02 PLUME02 DRIFT04 DRIFT04 DRIFT04 PLUME02 PLUME02 PLUME02 PLUME02 Latitudea -1.29 -1.29 -0.08 0.37 0.49 -0.55 -0.19 -0.21 0.07 Longitudea -90.91 -90.91 -91.95 -89.85 -89.83 -88.53 -88.67 -88.65 -89.04 Average Depth (m) 2719 2719 3100 819 2126 1074 1006 1251 1050 Major Elements (wt%) SiO2 49.46 49.26 50.22 50.39 50.45 49.98 49.26 49.49 49.25 TiO2 3.05 3.00 1.16 1.17 1.24 1.50 2.08 1.60 1.03 Al2O3 13.76 14.20 17.02 14.92 14.49 14.74 14.99 14.50 16.09 FeO* 12.29 11.83 8.50 10.16 10.62 10.38 11.32 11.54 9.28 MnO 0.20 0.20 0.20 0.17 0.21 0.23 0.20 0.24 0.17 MgO 5.98 6.24 8.63 7.89 7.44 7.41 6.17 7.05 8.44 CaO 10.95 11.15 12.03 13.13 12.94 12.87 11.48 12.51 12.44 Na2O 2.77 2.78 2.48 2.58 2.65 2.68 3.47 2.76 2.21 K 2O 0.51 0.51 0.14 0.06 0.05 0.11 0.38 0.07 0.04 P 2O 5 0.46 0.31 0.13 0.09 0.08 0.14 0.29 0.18 0.04 Total 99.43 99.47 100.51 100.55 100.17 100.05 99.64 99.95 98.99 Volatile Content H2O (wt%) 0.640 0.646 0.252 0.148 0.165 0.299 0.533 0.290 0.140 CO2 (ppm) 139.07 108.65 153.80 70.64 129.50 35.79 35.15 72.91 22.61 F (ppm) 486.13 505.26 161.64 125.05 124.48 194.57 376.27 185.01 91.28 S (ppm) 1149.12 1253.89 823.55 1015.30 1130.29 1049.24 1090.64 1187.18 938.56 Cl (ppm) 134.07 155.46 40.66 52.10 148.02 27.50 133.74 14.47 81.30 ! 184! P saturation (bars) 337 273 337 155 282 85 101 165 51 Trace Elements (ppm) Rb 8.70 9.66 1.53 9.17 2.44 1.34 5.95 0.79 0.38 Ba 83.91 119.29 11.63 5.32 5.36 14.94 60.42 8.39 4.46 Th 1.45 1.61 0.18 0.06 0.07 0.19 0.56 0.15 0.05 U 0.421 0.470 0.051 0.024 0.029 0.057 0.199 0.038 0.018 Nb 23.35 24.75 2.49 1.02 1.08 2.80 8.59 1.68 0.70 K 4008.54 4448.10 852.20 329.57 360.31 873.08 3049.16 552.99 232.27 La 15.74 16.85 2.98 1.64 1.71 2.98 7.50 2.69 1.06 Ce 37.04 40.43 8.95 5.99 6.03 8.95 19.94 9.33 4.05 Pb 0.33 0.26 0.39 0.31 0.30 0.34 0.53 0.34 0.19 Sr 301.22 303.38 156.86 110.27 101.93 140.15 205.56 120.01 80.97 Nd 22.47 24.42 8.00 7.08 7.01 8.99 15.62 9.67 5.55 P 1413.42 1420.12 542.80 379.69 362.30 594.15 1135.83 549.51 261.48 Zr 182.12 182.00 76.20 62.89 63.18 88.98 158.25 98.48 44.63 Hf 4.28 4.65 1.91 2.00 2.11 2.48 3.68 2.74 1.92 Sm 5.68 6.12 2.91 2.54 2.92 3.30 4.62 3.75 2.34 Eu 2.05 1.95 1.09 1.06 1.14 1.24 1.59 1.39 0.94 Ti 18380.36 18267.52 6772.97 6906.85 7314.44 8555.03 12190.16 9297.61 6087.44 Gd 6.20 6.79 3.92 3.70 4.08 4.57 5.41 5.98 3.37 Dy 5.77 6.17 4.38 4.46 4.99 4.85 6.26 7.28 4.09 Y 30.26 30.52 21.92 26.19 28.14 27.86 37.41 36.45 22.86 Er 3.05 3.56 2.63 2.75 3.04 2.89 3.91 4.71 2.58 Yb 2.65 2.91 2.56 2.59 2.86 2.71 3.64 4.54 2.41 Pr 1.62 1.21 1.19 1.90 1.90 0.83 Tb 0.68 0.69 0.70 0.87 1.11 0.61 Ho 0.95 0.99 1.01 1.18 1.67 0.90 Tm 0.41 0.43 0.44 0.50 0.72 0.39 Lu 0.39 0.41 0.41 0.47 0.74 0.37 Isotopes 3 He/4Heb 16.56 17.21 8.36 8.79 8.12 7.72 8.04 1σ 0.19 0.13 0.06 0.06 0.04 0.42 0.07 C/3He 5.88E+10 ! 185! 87 Sr/86Src 0.70337 0.70343 0.70265 0.70260 0.70270 0.70276 0.70264 143 Nd/144Ndc 0.512945 0.512912 0.513095 0.513101 0.513111 0.513126 εNd 5.99 5.35 8.92 9.03 9.23 9.52 206 Pb/204Pbc 19.329 19.307 18.471 18.418 18.514 18.758 18.766 18.790 207 Pb/204Pbc 15.586 15.569 15.492 15.507 15.512 15.584 15.541 15.629 208 Pb/204Pbc 38.948 38.908 37.971 37.932 38.013 38.554 38.290 38.692 Group ITD ITD ITD ITD ITD ITD ITD ITD ITD Sample PL02-13-25 PL02-13-27 PL02-14-3 PL02-14-8 PL02-27-1 PL02-27-17 PL02-9-20 PL02-9-29 PL02-9-33 Cruise PLUME02 PLUME02 PLUME02 PLUME02 PLUME02 PLUME02 PLUME02 PLUME02 PLUME02 Latitudea 0.07 0.07 0.10 0.07 0.88 0.88 -0.47 -0.47 -0.47 Longitudea -89.04 -89.04 -89.09 -89.12 -91.32 -91.32 -88.53 -88.53 -88.53 Average Depth (m) 1050 1050 1365 1365 1557 1557 1180 1180 1180 Major Elements (wt%) SiO2 50.02 49.57 48.11 48.37 48.76 48.61 47.34 47.47 47.41 TiO2 1.03 1.10 1.17 1.19 1.83 1.77 0.99 1.01 1.02 Al2O3 16.00 15.58 17.10 17.17 15.53 15.64 17.77 17.90 17.75 FeO* 9.68 9.87 9.15 9.03 11.17 11.11 9.07 9.00 9.02 MnO 0.17 0.19 0.17 0.18 0.22 0.22 0.16 0.17 0.18 MgO 8.47 8.21 8.86 8.80 7.27 7.26 9.54 9.59 9.62 CaO 12.42 12.83 12.73 12.66 11.93 11.82 12.17 11.93 12.49 Na2O 2.36 2.35 2.55 2.52 3.12 3.02 2.34 2.30 2.36 K 2O 0.04 0.03 0.03 0.02 0.08 0.10 0.02 0.02 0.01 P 2O 5 0.19 0.07 0.19 0.06 0.14 0.33 0.11 0.07 0.14 Total 100.39 99.80 100.05 100.00 100.05 99.87 99.51 99.45 99.15 Volatile Content H2O (wt%) 0.133 0.156 0.117 0.107 0.346 0.327 0.166 0.176 0.156 CO2 (ppm) 78.18 23.18 102.27 85.45 118.84 92.44 16.99 184.84 147.55 F (ppm) 82.78 97.19 124.20 123.65 182.10 178.89 120.53 120.27 119.95 S (ppm) 876.18 978.52 950.44 969.62 1032.48 1033.03 886.88 891.12 901.15 Cl (ppm) 156.92 93.00 3.80 4.19 53.61 77.62 9.38 10.37 10.36 ! 186! P saturation (bars) 171 52 160 146 242 179 24 240 249 Trace Elements (ppm) Rb 0.36 0.43 0.19 0.22 1.10 1.08 0.11 0.13 0.13 Ba 3.52 4.55 2.37 2.71 12.49 10.40 1.87 1.99 2.34 Th 0.04 0.05 0.03 0.03 0.16 0.17 0.02 0.03 0.03 U 0.016 0.020 0.010 0.010 0.050 0.054 0.010 0.011 0.011 Nb 0.65 0.81 0.43 0.44 2.18 2.27 0.42 0.42 0.48 K 203.19 243.64 158.79 166.10 648.53 644.26 146.63 150.97 153.71 La 0.92 1.13 1.07 1.21 3.10 3.01 1.11 1.22 1.27 Ce 3.29 4.25 4.71 5.39 10.24 10.24 4.25 5.25 5.49 Pb 0.15 0.21 0.22 0.26 0.40 0.36 0.22 0.29 0.31 Sr 79.22 79.91 106.17 113.25 136.02 139.01 117.70 128.60 132.86 Nd 4.77 5.72 6.52 7.59 11.23 11.05 5.24 6.34 6.62 P 255.30 272.38 310.57 310.37 652.76 625.00 319.82 286.97 296.77 Zr 41.31 47.49 60.28 61.83 97.12 101.49 57.26 57.32 59.28 Hf 1.44 1.93 1.91 2.43 3.18 2.99 1.51 1.81 1.92 Sm 2.04 2.29 2.54 2.94 4.09 4.16 1.89 2.19 2.39 Eu 0.85 0.94 1.08 1.16 1.55 1.56 0.80 0.91 0.95 Ti 5896.57 6383.48 6681.36 7009.40 10787.35 10587.87 5435.53 5796.76 5950.51 Gd 2.83 3.37 3.64 4.13 5.60 5.79 2.50 3.26 3.26 Dy 3.63 4.06 4.03 4.63 6.13 6.34 2.93 3.71 3.65 Y 21.84 23.82 25.14 25.85 34.39 35.65 18.29 20.02 20.95 Er 2.28 2.58 2.59 2.89 3.85 3.85 1.80 2.37 2.41 Yb 2.05 2.42 2.39 2.64 3.40 3.48 1.72 2.07 2.13 Pr 0.77 0.90 1.19 1.20 2.01 1.95 1.14 1.13 1.12 Tb 0.55 0.60 0.67 0.74 1.01 1.03 0.54 0.59 0.54 Ho 0.77 0.86 0.95 1.06 1.39 1.40 0.76 0.87 0.79 Tm 0.32 0.37 0.40 0.46 0.58 0.57 0.33 0.37 0.34 Lu 0.31 0.34 0.39 0.43 0.53 0.55 0.32 0.36 0.32 Isotopes 3 He/4Heb 7.27 8.78 8.78 8.51 8.53 8.19 1σ 0.04 0.05 0.05 0.05 0.05 0.05 C/3He ! 187! 87 Sr/86Src 0.70261 0.70257 0.70257 0.70261 0.70257 0.70265 143 Nd/144Ndc 0.513116 0.513153 0.513153 0.513086 0.513044 0.51313 εNd 9.32 10.05 10.05 8.74 7.92 9.60 206 Pb/204Pbc 18.760 18.396 18.396 18.641 18.617 18.339 207 Pb/204Pbc 15.603 15.490 15.490 15.516 15.482 15.484 208 Pb/204Pbc 38.624 37.854 37.854 38.076 37.975 37.775 Group ITD ITD ITD ITD ITD ITD ITD ITD ITD Sample PL02-9-47 PL02-9-51 DR69-7 DR70-2 DR71-A-1 DR72-3 DR72-A-1 DR77-9 DR78-1 Cruise PLUME02 PLUME02 SO 158 SO 158 SO 158 SO 158 SO 158 SO 158 SO 158 Latitudea -0.47 -0.47 0.58 0.41 0.45 0.11 0.09 0.19 0.29 Longitudea -88.53 -88.53 -90.33 -89.87 -89.63 -89.29 -89.25 -88.88 -88.77 Average Depth (m) 1180 1180 1193 1483 1450 1982 1441 1617 1871 Major Elements (wt%) SiO2 47.76 47.88 50.26 50.04 48.51 48.71 47.87 50.78 50.78 TiO2 1.01 1.01 0.97 1.30 1.10 0.85 0.86 0.97 1.06 Al2O3 17.84 17.77 15.55 14.55 16.08 17.00 18.38 14.56 14.32 FeO* 9.07 9.15 9.63 10.76 10.12 9.87 7.87 9.84 10.46 MnO 0.15 0.17 0.20 0.18 0.16 0.19 0.15 0.19 0.16 MgO 9.53 9.41 8.44 7.41 8.42 8.81 9.85 8.06 7.79 CaO 12.04 12.35 13.36 12.78 12.89 12.23 12.78 13.08 12.80 Na2O 2.35 2.38 2.13 2.65 2.39 2.38 2.30 2.14 2.19 K 2O 0.02 0.03 0.05 0.08 0.05 0.03 0.01 0.02 0.04 P 2O 5 0.11 0.07 0.06 0.11 0.07 0.04 0.05 0.05 0.06 Total 99.88 100.21 100.65 99.87 99.79 100.10 100.11 99.70 99.66 Volatile Content H2O (wt%) 0.156 0.174 0.125 0.193 0.192 0.111 0.100 0.114 0.141 CO2 (ppm) 171.19 161.96 77.18 91.99 94.73 114.94 103.17 89.44 181.97 F (ppm) 121.02 121.67 90.98 138.82 123.07 73.78 89.97 85.44 99.56 S (ppm) 903.23 897.98 902.63 1138.23 1005.04 880.23 759.58 1005.87 1109.80 Cl (ppm) 9.68 10.21 37.05 20.00 41.05 11.21 4.05 82.89 25.40 ! 188! P saturation (bars) 241 236 168 202 171 221 150 195 394 Trace Elements (ppm) Rb 0.11 0.13 0.88 0.69 0.47 0.1136 0.18 0.36 0.53 Ba 2.02 2.21 10.63 8.88 5.82 2.8114 1.83 4.70 6.70 Th 0.03 0.03 0.12 0.10 0.08 0.0233 0.02 0.06 0.07 U 0.010 0.012 0.037 0.030 0.024 0.0092 0.009 0.016 0.019 Nb 0.43 0.45 1.83 1.37 1.06 0.5427 0.26 0.69 1.00 K 160.03 158.94 436.56 499.36 386.62 124.007 143.28 212.27 278.98 La 1.06 1.32 1.76 2.24 1.80 0.8426 0.87 0.97 1.27 Ce 4.44 5.40 5.05 7.76 6.13 3.1035 4.51 3.62 4.39 Pb 0.22 0.30 0.18 0.32 0.25 0.1194 0.21 0.14 0.24 Sr 120.42 132.11 77.51 123.93 114.31 78.8355 120.67 62.35 62.66 Nd 5.05 6.56 5.45 8.23 6.54 4.0398 5.44 4.70 5.35 P 315.18 287.70 282.71 425.04 340.43 188.3518 235.07 264.90 289.20 Zr 57.90 58.70 40.67 69.96 56.90 33.2839 39.39 37.48 44.01 Hf 1.53 1.77 1.82 2.58 2.01 1.2754 1.51 1.64 1.98 Sm 1.98 2.27 2.06 3.16 2.32 1.4307 2.08 1.82 2.14 Eu 0.80 0.95 0.85 1.23 0.94 0.693 0.91 0.83 0.83 Ti 5556.45 5931.13 5640.78 7789.73 6348.72 4652.3108 5593.19 5609.10 6214.21 Gd 2.77 3.16 2.99 4.41 3.43 2.2304 2.82 3.32 3.32 Dy 3.03 3.59 3.88 4.94 3.93 2.9287 3.19 4.18 4.11 Y 19.00 20.23 22.19 27.36 22.37 18.5673 16.98 21.93 24.29 Er 1.88 2.28 2.37 3.15 2.54 1.9272 1.94 2.54 2.76 Yb 1.81 2.08 2.27 2.95 2.38 1.8625 1.89 2.47 2.49 Pr 1.13 1.15 0.93 1.47 1.29 0.80 0.94 0.86 0.97 Tb 0.58 0.57 0.56 0.76 0.69 0.49 0.50 0.67 0.70 Ho 0.82 0.83 0.83 1.09 0.98 0.75 0.69 1.02 1.08 Tm 0.35 0.36 0.36 0.46 0.43 0.34 0.29 0.43 0.46 Lu 0.35 0.34 0.35 0.44 0.42 0.33 0.28 0.42 0.46 Isotopes 3 He/4Heb 1σ C/3He ! 189! 87 Sr/86Src 143 Nd/144Ndc εNd 206 Pb/204Pbc 207 Pb/204Pbc 208 Pb/204Pbc Group ITD ITD ITD WD Group WD WD WD WD WD Sample DR79-1 DR80-2 DR83-6 PL02-26-16 PL02-26-21 PL02-26-25 PL02-26-7 PL02-28-1 PL02-28-11 Cruise SO 158 SO 158 SO 158 PLUME02 PLUME02 PLUME02 PLUME02 PLUME02 PLUME02 Latitudea 0.42 -0.02 -0.31 0.83 0.83 0.83 0.84 1.02 1.02 Longitudea -88.69 -88.51 -88.39 -91.29 -91.29 -91.29 -91.29 -91.55 -91.55 Average Depth (m) 2377 1606 1321 676 676 676 676 1183 1183 Major Elements (wt%) SiO2 47.87 50.56 47.89 49.13 49.47 48.64 49.32 48.69 48.49 TiO2 0.95 1.27 0.96 1.87 1.87 1.79 1.89 1.76 1.78 Al2O3 17.08 14.72 17.51 15.69 15.78 15.64 15.67 16.82 16.59 FeO* 9.82 10.04 8.97 10.17 10.10 11.11 10.08 9.74 9.80 MnO 0.16 0.20 0.16 0.17 0.20 0.18 0.21 0.16 0.16 MgO 9.65 7.58 9.29 7.45 7.57 7.03 7.46 7.98 7.84 CaO 11.96 12.71 12.62 11.89 11.98 11.80 11.97 11.68 11.74 Na2O 2.42 2.53 2.38 2.96 2.92 3.05 2.95 2.90 2.85 K 2O 0.05 0.05 0.03 0.26 0.25 0.11 0.25 0.20 0.18 P 2O 5 0.06 0.10 0.07 0.20 0.20 0.36 0.18 0.19 0.18 Total 100.02 99.77 99.89 99.80 100.34 99.72 99.99 100.13 99.62 Volatile Content H2O (wt%) 0.185 0.200 0.141 0.444 0.401 0.407 0.461 0.314 0.322 CO2 (ppm) 147.36 87.12 179.13 41.14 23.17 86.83 40.7 52.14 64.1 F (ppm) 120.81 144.06 105.93 270.79 262.58 181.46 270.36 234.07 231.4 S (ppm) 917.63 1048.25 880.49 780.52 715.78 1026.42 787.06 759.39 814.09 Cl (ppm) 43.45 14.95 28.29 69.78 70.61 88.85 70.05 53.29 49.63 ! 190! P saturation (bars) 215 192 261 106 64 177 107 108 122 Trace Elements (ppm) Rb 0.52 0.58 0.2185 4.3099 4.3832 1.07 4.3881 3.6079 3.4015 Ba 8.40 7.06 2.4501 49.7096 49.9265 10.89 48.222 38.4638 33.3533 Th 0.09 0.08 0.0404 0.7122 0.7072 0.18 0.7242 0.5462 0.4975 U 0.026 0.024 0.0156 0.2155 0.2358 0.052 0.203 0.1777 0.1714 Nb 1.25 1.13 0.3938 9.0863 8.9781 2.28 8.7721 6.9053 6.2984 K 344.67 362.23 165.5492 2167.4007 2070.7699 663.66 2095.6714 1640.82 1467.5311 La 1.72 1.78 1.4002 7.5789 7.9264 3.10 7.5265 6.3928 5.8035 Ce 5.41 6.15 5.8641 18.9457 19.9004 10.07 19.0325 17.0915 15.5883 Pb 0.21 0.23 0.2387 0.1542 0.1931 0.34 0.1614 0.1599 0.1529 Sr 106.76 96.32 117.9786 205.5402 212.0721 137.55 204.6942 187.1786 175.7391 Nd 5.69 7.21 6.6622 14.1379 14.9552 11.05 14.2343 14.2505 13.1395 P 285.72 418.73 292.4683 827.2338 862.4216 636.26 841.2189 825.4759 788.8159 Zr 50.60 61.15 52.3018 112.5106 116.5399 104.00 114.2235 114.0026 109.4997 Hf 1.63 2.28 1.7579 3.2915 3.4859 3.06 3.1139 3.6078 3.2884 Sm 1.98 2.67 2.4588 4.0461 4.4965 3.97 4.413 4.4087 4.2039 Eu 0.80 1.06 1.0246 1.2688 1.3099 1.49 1.2404 1.26 1.219 Ti 5542.24 7271.43 5905.574 11254.6069 11334.5612 10831.36 11215.4798 10887.4561 10661.6578 Gd 2.88 3.75 3.4962 4.5072 4.845 5.68 4.5813 4.7833 4.5212 Dy 3.46 4.58 4.1353 5.1057 5.4858 6.47 5.1861 5.7824 5.5199 Y 19.67 25.90 18.9853 27.5403 28.7964 35.91 27.4486 30.2221 29.6103 Er 2.23 2.94 2.6251 3.2456 3.5035 3.90 3.16 3.5737 3.6198 Yb 2.00 2.66 2.4582 2.6021 2.8197 3.57 2.6355 3.1181 2.9872 Pr 1.25 1.53 1.19 1.98 Tb 0.62 0.82 0.63 1.07 Ho 0.93 1.19 0.93 1.45 Tm 0.39 0.50 0.40 0.60 Lu 0.38 0.48 0.39 0.57 Isotopes 3 He/4Heb 8.79 8.80 8.51 8.81 8.42 8.45 1σ 0.05 0.05 0.05 0.05 0.05 0.05 C/3He ! 191! 87 Sr/86Src 0.70284 0.70283 0.70287 0.70279 0.70276 0.70277 143 Nd/144Ndc 0.513003 0.513011 0.512994 0.512997 0.513038 0.513037 εNd 7.12 7.28 6.94 7.00 7.80 7.78 206 Pb/204Pbc 19.121 19.152 19.161 19.127 19.067 19.028 207 Pb/204Pbc 15.584 15.592 15.582 15.575 15.588 15.564 208 Pb/204Pbc 38.778 38.844 38.766 38.791 38.673 38.602 Group WD WD WD WD WD WD WD WD WD Sample PL02-29-3 PL02-30-1 DR40-1 DR40-3 DR41-A-1 DR46-A-1 DR47-A-1 DR50-1 DR58-B-2 Cruise PLUME02 PLUME02 SO 158 SO 158 SO 158 SO 158 SO 158 SO 158 SO 158 Latitudea 1.90 1.59 0.98 0.98 1.24 1.71 1.89 1.87 1.93 Longitudea -92.16 -91.92 -90.84 -90.84 -91.12 -91.19 -91.28 -91.70 -92.08 Average Depth (m) 1723 2256 1560 1560 445 2235 1603 2214 2304 Major Elements (wt%) SiO2 50.25 48.34 49.15 49.44 50.17 49.01 50.22 50.21 48.93 TiO2 1.38 1.52 1.21 1.18 1.72 2.11 1.25 1.82 1.24 Al2O3 14.51 16.38 16.42 16.32 15.41 15.92 15.04 14.57 16.12 FeO* 10.41 9.52 10.67 10.58 10.38 10.60 10.03 11.64 10.30 MnO 0.19 0.19 0.17 0.19 0.18 0.19 0.20 0.22 0.21 MgO 7.32 7.82 7.93 8.01 7.28 7.14 7.98 6.90 8.21 CaO 12.48 12.14 11.54 11.46 12.12 11.00 12.90 11.08 12.10 Na2O 2.43 2.58 2.70 2.66 2.82 2.78 2.35 2.53 2.34 K 2O 0.18 0.36 0.20 0.21 0.24 0.43 0.11 0.27 0.09 P 2O 5 0.37 0.32 0.12 0.15 0.16 0.33 0.13 0.26 0.15 Total 99.52 99.18 100.12 100.19 100.48 99.51 100.21 99.51 99.69 Volatile Content H2O (wt%) 0.249 0.472 0.385 0.345 0.360 0.702 0.236 0.445 0.232 CO2 (ppm) 114.9 120.75 100.78 94.44 10.73 95.68 193.72 116.94 159.74 F (ppm) 200.19 325.69 217.55 216.52 265.00 395.15 155.88 303.49 151.04 S (ppm) 1127.04 964.90 1078.48 1093.09 855.38 1228.66 999.99 1279.18 1045.01 Cl (ppm) 123.93 181.75 115.98 120.36 126.42 278.27 110.98 272.66 102.55 ! 192! P saturation (bars) 253 222 230 214 34 252 421 270 349 Trace Elements (ppm) Rb 3.8423 6.8494 4.9368 5.12 4.67 6.77 1.78 6.19 1.54 Ba 34.9704 87.8552 59.4427 62.54 56.53 97.67 24.09 58.63 19.06 Th 0.5294 1.204 0.8013 0.86 0.83 1.30 0.33 0.90 0.26 U 0.1539 0.3114 0.1922 0.211 0.249 0.350 0.094 0.235 0.075 Nb 6.1775 15.868 8.5886 9.00 8.95 15.87 3.79 9.61 3.30 K 1338.0145 3057.23 1641.6904 1740.38 2174.95 3623.12 931.32 2094.54 713.95 La 4.8853 10.98 6.5166 6.75 7.67 12.10 3.42 7.19 3.14 Ce 11.9439 24.28 14.6067 14.82 18.70 28.41 9.44 17.83 8.62 Pb 0.245 0.23 0.1716 0.19 0.19 0.21 0.31 0.17 0.30 Sr 111.6015 220.97 144.668 148.58 163.53 239.72 97.78 113.61 89.48 Nd 9.463 14.07 9.0857 9.30 13.46 18.03 7.88 13.32 7.24 P 560.2203 897.10 519.7734 516.92 806.37 1237.91 472.95 906.50 440.69 Zr 73.464 109.11 64.6133 65.85 105.54 144.27 64.28 109.26 61.79 Hf 2.2616 2.60 2.1215 2.32 3.35 3.79 2.28 3.50 2.21 Sm 3.0594 3.81 2.5899 2.62 4.26 5.01 2.62 4.02 2.50 Eu 1.0151 1.35 0.9007 0.98 1.30 1.55 0.99 1.20 0.97 Ti 8005.1267 8890.25 7157.9846 7254.28 10531.18 12825.14 7496.73 10984.16 7232.57 Gd 3.9043 4.14 3.0823 3.13 4.57 5.20 3.92 4.72 3.68 Dy 5.1269 4.61 3.8605 3.96 5.62 5.65 4.79 6.10 4.52 Y 28.4253 25.44 21.3247 21.51 30.11 30.36 25.19 35.76 26.58 Er 3.0279 2.69 2.6033 2.64 3.76 3.67 2.98 4.18 2.93 Yb 2.846 2.50 2.1235 2.19 3.15 3.05 2.91 3.69 2.87 Pr 1.57 1.51 Tb 0.76 0.70 Ho 1.11 1.05 Tm 0.47 0.46 Lu 0.46 0.46 Isotopes 3 He/4Heb 7.81 6.87 1σ 0.05 0.08 C/3He ! 193! 87 Sr/86Src 0.70295 0.70301 143 Nd/144Ndc 0.513008 0.512964 εNd 7.22 6.36 206 Pb/204Pbc 19.001 19.366 207 Pb/204Pbc 15.599 15.620 208 Pb/204Pbc 38.895 39.282 Wolf Wolf Wolf Wolf Group WD WD WD WD WD Volcano Volcano Volcano Volcano Sample DR60-1 DR62-2 DR65-1 DR68-3 DR68-4 D3a D4e D3a-2 D4a Cruise SO 158 SO 158 SO 158 SO 158 SO 158 DRIFT04 DRIFT04 DRIFT04 DRIFT04 Latitudea 1.72 1.19 0.89 0.86 0.86 0.27 0.19 0.27 0.19 Longitudea -91.72 -91.69 -91.42 -90.84 -90.84 -91.43 -91.41 -91.43 -91.41 Average Depth (m) 1848 2038 1493 1561 1561 2595 2015 2595 2015 Major Elements (wt%) SiO2 49.81 48.75 49.05 50.19 50.35 48.91 49.07 48.85 48.95 TiO2 1.25 0.68 2.10 2.53 2.55 3.02 3.50 3.23 4.30 Al2O3 15.50 17.53 15.15 13.80 13.83 14.59 13.86 14.48 13.28 FeO* 10.24 9.36 10.85 12.72 12.64 11.32 12.29 11.34 13.81 MnO 0.22 0.17 0.18 0.20 0.21 0.18 0.19 0.19 0.23 MgO 8.00 9.38 6.85 5.47 5.46 6.08 5.38 6.28 4.90 CaO 12.66 12.63 11.71 10.43 10.51 11.35 11.31 11.50 9.59 Na2O 2.50 2.05 3.12 3.22 3.21 3.10 3.32 3.22 3.67 K 2O 0.09 0.03 0.27 0.52 0.52 0.52 0.63 0.52 0.74 P 2O 5 0.15 0.09 0.24 0.32 0.30 0.36 0.46 0.38 0.53 Total 100.42 100.68 99.52 99.38 99.59 99.45 100.00 97.53 98.04 Volatile Content H2O (wt%) 0.206 0.098 0.479 0.662 0.512 0.867 0.839 1.033 1.2317 CO2 (ppm) 106.02 108.15 52.68 90.59 90.44 99.15 66.18 84.25 43.5488 ! 194! F (ppm) 141.67 61.40 315.35 535.08 512.15 502.78 647.46 526.06 840.6638 S (ppm) 1072.59 887.88 951.72 1582.37 1549.65 1168.1 1510.5 1318.57 2185.0281 Cl (ppm) 140.95 14.48 83.35 493.31 482.14 230.65 207.91 297.46 303.9469 P saturation (bars) 233 212 134 236 219 285 210 266 230 Trace Elements (ppm) Rb 1.43 0.35 4.80 9.55 9.72 5.6781 7.0773 6.76 10.1489 Ba 17.20 4.90 55.00 121.04 119.66 55.0504 65.3289 66.48 85.8537 Th 0.25 0.06 0.86 1.78 1.65 0.9846 1.2619 1.01 1.5163 U 0.066 0.022 0.244 0.472 0.454 0.3581 0.4539 0.382 0.6116 Nb 2.60 0.82 9.97 23.86 23.47 16.5052 20.6707 18.34 26.9698 K 667.41 191.39 2281.94 4293.91 4150.27 3996.7644 5014.1729 4120.73 6155.6629 La 2.86 1.02 8.53 16.52 16.73 14.3924 17.7609 15.18 21.9666 Ce 8.45 3.26 21.53 38.08 38.38 36.1482 44.6224 40.94 59.3694 Pb 0.27 0.13 0.20 0.28 0.30 1.785 1.8253 1.37 1.9228 Sr 97.57 68.53 211.90 235.27 236.68 381.852 364.4298 391.62 361.9929 Nd 8.14 3.58 16.19 22.51 23.10 24.3693 29.1668 26.65 37.1913 P 453.54 181.92 953.19 1339.56 1359.47 1670.7093 2070.1348 1583.62 2287.9665 Zr 66.07 28.80 131.14 160.24 163.79 228.4943 290.2098 232.21 348.0472 Hf 2.27 1.26 3.80 4.46 4.44 5.0699 6.2698 5.36 7.5852 Sm 2.74 1.54 4.82 5.77 5.77 6.4207 7.283 6.80 9.2496 Eu 1.05 0.65 1.41 1.82 1.87 1.9359 2.46 2.39 3.1334 Ti 7467.33 4009.74 12581.17 15536.64 15511.46 17993.7745 20957.4181 18612.71 24759.5253 Gd 4.19 2.34 5.09 6.35 6.37 6.829 7.9125 7.28 9.8653 Dy 4.99 3.16 5.99 6.43 6.37 5.9788 7.2362 6.53 8.7207 Y 27.67 18.34 31.13 34.20 34.90 31.9168 37.6285 32.59 44.3709 Er 3.23 2.06 3.69 4.03 4.08 3.2925 3.8514 3.16 4.273 Yb 3.26 2.10 3.00 3.35 3.25 2.7326 3.2357 2.77 3.7715 Pr 1.49 0.62 5.78 8.22 Tb 0.83 0.43 1.13 1.53 Ho 1.25 0.68 1.22 1.63 Tm 0.55 0.32 0.43 0.58 Lu 0.55 0.33 0.39 0.54 ! 195! Isotopes 3 He/4Heb 1σ C/3He 87 Sr/86Src 0.70275 143 Nd/144Ndc 0.513024 εNd 7.53 206 Pb/204Pbc 18.943 18.961 18.943 207 Pb/204Pbc 15.550 15.547 15.550 208 Pb/204Pbc 38.441 38.426 38.441 VG-2B VG-2B measured accepted Major Elements (wt%) SiO2 51.23 50.81 TiO2 1.86 1.85 Al2O3 13.98 14.06 FeO* 11.94 11.84 MnO 0.22 0.22 MgO 6.89 6.71 CaO 11.35 11.12 Na2O 2.59 2.62 K 2O 0.20 0.19 P 2O 5 0.21 0.20 Total 100.47 99.62 ! a reported as decimal degrees b He isotope analyses previously published by Graham et al., (1993) and Kurz et al., (2009) are in plain text while unpublished data are in bold c Sr-Nd-Pb isotope analyses previously published by Harpp and White (2001) and Geist et al., (2008) are in plain text while unpublished data are in bold ! 196! Table 2. Fractional Crystallization model Starting Composition Enriched Depleted Compositiona Compositionb Major Elements (wt%) SiO2 47.48 46.46 TiO2 2.29 1.34 Al2O3 16.38 14.36 FeO* 9.16 11.77 MnO 0.12 0.20 MgO 9.38 11.86 CaO 11.58 11.50 Na2O 2.25 2.14 K 2O 0.33 0.08 P 2O 5 0.26 0.07 H2O (wt%) 0.68 0.12 Volatile Elements (ppm) CO2 279.76 2.74 F 293.36 176.48 S 923.26 420.41 Cl 118.42 34.54 Starting Parametersc Volatile Kdsd Ol CPX PLAG CO2 5.00E-06 3.00E-04 0 F 0.00259 0.0464 0.084 S 0.0027 0.44 0.019 Cl 5.00E-06 3.00E-04 0 Pressure 2kbars Temperature 1200 Oxygen Fugacity QFM a Composition of melt inclusion AHA D25C-3-43 reported in Koleszar et al., (2009) b Composition of melt inclusion STG06-29-06 reported in Koleszar et al., (2009) c Models used in PETROLOG: Ol, CPX, PLAG- Danyushevskey, 2001, Melt density- Lange and Carmichael, 1987, Melt viscosity- Bottinga & Weill 1972. Oxygen fugacity set at QFM and then run closed for oxygen d CO2, S, and Cl partition coefficients (Kd) for olivine (OL) and clinopyroxene (CPX) based on the partition coefficients of Ba and Dy from Halliday et al. [1995]. Kd's for OL and CPX for fluorine based on values of Hauri et al., [2006]. Plagioclase Kd's for F and S based on Nd and Dy values from Bedard (2006) ! 197! Table 3. End-member compositions CO2/Nb Groupa † 1σ H2O/Ce 1σ F/Nd 1σ Cl/K 1σ S/Dy 1σ HHe group mean 4 2 202 26 20.1 0.8 0.057 0.009 218 21 median 4 2 200 26 19.9 0.8 0.057 0.009 216 21 ITE groupb mean 5 2 174 24 24.9 0.7 0.049 0.010 237 20 median 4 2 165 24 25.2 0.7 0.049 0.010 235 20 WD groupc mean 10 2 230 35 23.5 0.4 0.066 0.006 255 40 median 10 2 233 35 23.3 0.4 0.069 0.006 276 40 ITD groupd mean 300* 117° 17.3 1.5 0.026 0.002 228 16 median 300* 117° 16.6 1.5 0.025 0.002 236 16 † values are mean and median of extensively degassed samples and therefore do not represent mantle compositions * set by expected value for MORB mantle (Saal et al., 2002) ° set by average value of Santiago melt inclusions (Koleszar et al., 2009) a End-member compositions for HHe, ITE, and WD group determined from samples that have been filtered to Cl/K < 0.08, S/Dy > 175, Sr/Sr* ≤ 1, and are not within error of being sulfide saturated. Further conditions exist where indicated. Further conditions exist where indicated. 1σ standard deviation between samples used to define end-member are indicated. b Determined from a subset of samples that represent the end-member composition determined from F/Nd ratios: D36, D51a, D34a, D38a, D27a c Determined from a subset of samples with an isotopic composition closest to the Pinta component d Determined from a subset of samples with the lowest Cl/K ratios: PL02-14-3, PL02-14-8, DR72-A-1 ! 198! Table 4. Calculation of mantle viscosity: Wet diffusion/dislocation creep with constant [OH] Rheological Parametersa Diffusion creep Dislocation creep regime regime A 1.00E+06 90 n 1 3.5 p 3 0 r 1 1.2 E* (kJ/mol) 335 480 V* (m^3/mol) 4.00E-06 0.000011 P (Pa) 7.00E+09 7.00E+09 R 8.314 8.314 T (°C) 1400 1400 d (μm) 10000 10000 σ (MPa) 0.3 0.3 Calculation Strain Rate Diffusion creep regime H2O (ppm) C(OH) = H/10^6Sib (!!) Viscosity (Pa*s) PM 338.35 5532.0225 7.69E-15 3.90E+19 DMM 64.35 1052.1225 1.46E-15 2.05E+20 Dislocation creep regime PM 338.35 5532.0225 1.68E-13 1.79E+18 DMM 64.35 1052.1225 2.29E-14 1.31E+19 a Parameters found in Table 1 of Hirth and Kohlstedt (2003) b Calculated for Ol (Fo90)L 1ppm H2O = 16.35H/10^6 Si (Paterson, 1982) ! 199! CHAPTER 3 The oxidation state of Fe in glasses from the Galapagos Archipelago: variation in oxygen fugacity as a function of mantle source. M. E. Peterson1*, K. Kelley2, L. Cottrell3, A. Saal1, M. D. Kurz4 1 Department of Earth, Environmental and Planetary Sciences, Brown University, Providence, Rhode Island, 02912, United States 2 Graduate School of Oceanography, University of Rhode Island, Narragansett, Rhode Island, 02882, United States 3 Department of Mineral Sciences, Smithsonian Institution, Washington DC 20013, United States 4 Department of Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, 02543, United States To be submitted to Earth and Planetary Science Letters ! 200! ABSTRACT The oxidation state of the Earth’s mantle plays an intrinsic role in the magmatic evolution of the Earth. Heterogeneity in magmatic oxidation has been recognized in lavas from mid-ocean ridges with variable isotopic enrichment. This suggests that isotopically diverse ocean island basalts may also preserve a large variability in their magmatic oxidation state. Here we present new μ-XANES measurements of Fe3+/ΣFe (a proxy for system ƒO2), using a geochemically diverse suite of submarine glasses from the Galapagos Archipelago. Utilizing previously presented major elements, trace elements, dissolved volatiles and isotopic data we show that Fe3+/ΣFe ratios vary with the influence of shallow level processes. Fe3+/ΣFe ratios generally increase with magmatic differentiation, but still show a large variation at a given MgO, which suggests other shallow level processes are affecting Fe3+/ΣFe of the melt. Progressive degassing of sulfur is shown to accompany decreasing Fe3+/ΣFe ratios, while assimilation of hydrothermally altered crust is shown to increase Fe3+/ΣFe ratios. Even after taking these processes into account there is still variability in the Fe3+/ΣFe ratios of the isotopically distinct sample suites studied, yielding a magmatic ƒO2 that ranges from ΔQFM = +0.16 to +0.74 and showing that oxidation state varies as a function of mantle source composition in the Galapagos hotspot system. After correcting back to a common MgO content = 8.0 wt%, the trace element depleted group similar to MORB (ITD), and the trace element enriched group similar to Pinta (WD = high Th/La, Δ7/4, Δ8/4 ratios) show Fe3+/ΣFe ratios within range of MORB (ITD = 0.162 ± 0.003 and WD = 0.164 ± 0.006). The other trace element enriched group similar to Sierra Negra and Cerro Azul (ITE) shows evidence of mixing between an oxidized and reduced source with the reduced end- ! 201! member showing a radiogenic isotope signature similar to Pinta (maximum ITE = 0.177 ± 0.008). This suggests that mantle sources in the Galapagos that are thought to contain recycled components (WD and ITE) are different in terms of their oxidation state. The high 3He/4He Fernandina samples are shown to have the most oxidized glasses (HHe = 0.175 ± 0.006) with high C/3He ratios indicating the primitive mantle may be a more carbonated and oxidized source than previously thought. ! 202! INTRODUCTION Oxygen fugacity (ƒO2) is an intensive thermodynamic property that provides a measure of a system’s redox potential at equilibrium. ƒO2 controls the speciation and partitioning of multi-valent elements and affects igneous processes such as magmatic phase relations during melting and crystallization, depth of melt initiation and the composition of volcanic gases (e.g. Ballhaus, 1993, Holloway & Blank, 1994, Canil, 1999). Therefore, understanding ƒO2 variability in space and time can potentially constrain the evolution of the mantle, as well as present-day mantle melting conditions . The preservation of magmatic and mantle source ƒO2 remains a topic of considerable debate (Ballhaus, 1993, Canil, 2002, Lee et al., 2005, Cottrell et al., 2009, Kelley & Cottrell, 2009, Lee et al., 2010, Cottrell & Kelley, 2011, Kelley & Cottrell, 2012, Cottrell & Kelley, 2013). Using Fe3+/ΣFe ratios as a proxy for ƒO2, a large variation is found in measured Fe3+/ΣFe ratios in mid-ocean ridge basalts (MORB) and arc basalts (Bezos & Humler, 2005, Kelley & Cottrell, 2009, Cottrell & Kelley, 2011, Kelley & Cottrell, 2012, Cottrell & Kelley, 2013, Brounce et al., 2014, Le Voyer et al., 2015). Shallow level processes can account for some of this variation and can be tracked by variations in major, trace and volatile elements (H2O, CO2, F, S, and Cl). Magmatic differentiation can change Fe3+/ΣFe ratios due to early crystallizing phases, such as olivine, partitioning Fe2+ over Fe3+ and thereby concentrating Fe3+ in the melt, which causes an increase in Fe3+/ΣFe ratios as magma evolves. Degassing has the potential to affect Fe3+/ΣFe ratios when there is an oxidation state difference between volatile species dissolved in the melt and the major degassing species of volatile (Anderson & Wright, 1972, Kelley & Cottrell, 2012, Brounce et al., 2014). Assimilation of wall rock by the ! 203! melt can also affect Fe3+/ΣFe ratios as melt interacts with oxidized crust. Shallow level processes, however, are shown to inadequately recreate the total variation between measured Fe3+/ΣFe ratios of MORB and arc basalts, which suggests the mantle source ƒO2 of these settings is fundamentally different (Kelley & Cottrell, 2012, Brounce et al., 2014). Recently, Fe3+/ΣFe ratios were shown to vary systematically as a function of source composition in MORB with more enriched MORB (EMORB) mantle domains with higher Pb isotope ratios being more reduced than depleted MORB domains (samples filtered to only non-plume influenced MORB by Cottrell and Kelley (2013)). Ocean island basalts (OIB) are derived from several global geochemically distinct mantle reservoirs and display extensive regional geochemical heterogeneity (e.g. Hofmann, 1997, Stracke et al., 2005, White, 2010). If the mantle can indeed preserve long-lived heterogeneities in Fe3+/ΣFe ratios, OIBs have the potential to record large variations in and the possible evolution of mantle source oxidation state. However, whether OIB melt from a source that contains more primitive mantle that houses high 3He/4He and solar-like neon isotopes and/or recycled crustal components preserve measurably different oxidation state from either depleted upper mantle or arc mantle has yet to be extensively tested. The goal of this study is to evaluate the oxidation state of the Galapagos mantle using μ-XANES to measure Fe3+/ΣFe ratios of submarine basaltic glass collected from across the archipelago. By using major, trace and volatile elements in addition to radiogenic isotope ratios, we show that some of the variation in Fe3+/ΣFe ratios that we measure is due to shallow level processes such as magmatic differentiation, degassing, ! 204! and assimilation of hydrothermally altered material. However, shallow level processes cannot explain all of the variation in Fe3+/ΣFe ratios. There is evidence for variable oxidation states between the isotopically unique end-members in the Galapagos mantle source. The trace element depleted end-member in the Galapagos lavas has Fe3+/ΣFe ratios comparable to MORB values at similar MgO contents (Cottrell & Kelley, 2011, Le Voyer et al., 2015). The high 3He/4He end-member and the trace element and isotope enriched end-member similar to Sierra Negra and Cerro Azul have the highest Fe3+/ΣFe ratios of this study, while the trace element enriched end-member similar to Pinta (high Th/La and Δ7/4, Δ8/4) has intermediate Fe3+/ΣFe ratios to these and the depleted end- member at similar MgO contents. These results place new constraints on the ƒO2 relevant for OIB magmas. We will explore the implications of these findings in terms of mechanisms controlling ƒO2 of the OIB mantle during plume melting. GEOLOGIC BACKGROUND AND PREVIOUS WORK In the Galapagos Archipelago, located on the equatorial Pacific, 16 of the 21 emergent volcanoes have erupted since the Holocene, making the Galapagos one of the most active volcanic regions in the world. The islands and seamounts sit on a shallow submarine platform 100-200 km to the south of the Galapagos Spreading Center (GSC), on the Nazca plate which is moving eastward with respect to a hotspot reference frame (Gripp & Gordon, 2002). Seismic studies have identified both a low-velocity anomaly and a thin transition zone in the mantle below the Galapagos Archipelago, characteristics attributed to a hot mantle plume column extending to at least as deep as the transition zone (Hooft et al., 2003, Villagomez et al., 2007, Villagomez et al., 2011, Villagomez et al., 2014). ! 205! The plume column tilts towards the GSC, against the direction of plate motion (Villagomez et al., 2014). This in conjunction with geochemical studies conducted along the ridge show the strong influence the Galapagos plume has on the GSC (Detrick et al., 2002, Schilling et al., 2003, Cushman et al., 2004, Ingle et al., 2010, Gibson et al., 2015). There are four main compositional components that are generally distinguishable geographically from each other in the Galapagos Archipelago. The western component, found primarily near Fernandina Island, is characterized by high 3He/4He ratios reaching up to 29 Ra (3He/4He ratio normalized to the atmospheric ratios, 1.4e-6) in addition to intermediate 87Sr/86Sr, 143Nd/144Nd, 206Pb/204Pb, 207Pb/204Pb, 208Pb/204Pb and 176Hf/177Hf ratios (Graham et al., 1993, White et al., 1993, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Kurz et al., 2009) analogous to a minimally degassed, primitive mantle component (e.g. Kurz et al., 1982, Farley et al., 1992, Hart et al., 1992, Hanan & Graham, 1996). The northern component, characterized by the lavas of Pinta 207 Island and the Wolf/Darwin lineaments, is discernible by elevated Pb/204Pb and 208 Pb/204Pb at a given 206 Pb/204Pb, resulting in high Δ7/4 and Δ8/4 as defined by Hart 143 (1984), and the lowest Nd/144Nd and 176 Hf/177Hf isotopic ratios yet measured in the Galapagos (Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Saal et al., 2007). The southern component, characterized by lavas from Floreana Island with the highest 87Sr/86Sr, 206 Pb/204Pb, 207 Pb/204Pb and 208 Pb/204Pb isotopic ratios found in the Galapagos with intermediate 3He/4He ratios (up to 14 Ra), 143 Nd/144Nd isotopic ratios 176 and unusually elevated Hf/177Hf ratios for the measured Nd isotopes (Bow & Geist, 1992, White et al., 1993, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Kurz et al., 2009, Harpp, 2014). While the Pb isotopes of this component ! 206! 87 are typical of a HIMU mantle component, the high Sr/86Sr and 176 Hf/177Hf ratios are unusual in comparison to other HIMU localities (Bow & Geist, 1992, Harpp & White, 2001, Harpp et al., 2014). There is also a strong presence of a depleted component similar to MORB throughout the Galapagos, which is concentrated in the central and northeastern region of the archipelago. In general, lavas from Genovesa Island characterize this component, which is hypothesized to be either the product of depleted upper mantle, or a depleted component intrinsic to the plume. The submarine basaltic glass measured for Fe3+/ΣFe ratios for this study have been previously characterized for major, trace, and volatile elements, as well as isotopic ratios (Harpp & White, 2001, Geist et al., 2006, Kurz et al., 2009, Peterson et al., submitted to Journal of Petrology). Glass chips were taken from dredges collected during the PLUME02 cruise in 1990 (Christie et al., 1992), AHA-Nemo-2 cruise using the R/V Melville in 2000 (Geist et al., 2006), the DRIFT04 cruise using the R/V Revelle in 2001 (Kurz et al., 2001), and the R/V SONNE cruise, SO 158 in 2001 (Werner, 2002) from depths ranging from 445-3504 m below sea level. The samples are collected from across the archipelago (Fig. 1) and generally fall on 2 compositional trends. These trends cover almost the entire range reported in the Galapagos, with the exception of the Floreana component, which was not sampled. The first array (Array 1) ranges between the compositions of Fernandina Islands and Sierra Negra/Cerro Azul Volcanoes. Lavas from Sierra Negra/Cerro Azul, located on the southern half of Isabela Island, are geochemically intermediate between the compositions of Fernandina, Floreana and Pinta (White et al., 1993, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Saal et al., 2007, Harpp, 2014). The ! 207! compositions of Sierra Negra and Cerro Azul are thus either the result of a mixture of melts from several end-members or the product of a unique mantle source with an intermediate composition. Though it is possible that a unique composition exists (Vidito et al., 2013), samples from Sierra Negra and Cerro Azul fall over a large compositional space making it more likely they are the result of mixing of melts from several sources. Array 1 is sub-divided into the high 3He/4He ratio (HHe) group and the incompatible trace element enriched (ITE) group using 3He/4He ratios above and below 22 Ra respectively. Array 1 is characterized by samples that are enriched in incompatible trace elements (high Th/La ratios) and show characteristics consistent with generation in the garnet stability field (high Sm/Yb ratios) (Fig. 2). The HHe group also possesses a higher average Ti/Gd ratio than the ITE group, consistent with proposed high Ti characteristic of the a deep primitive mantle source (e.g. Kurz & Geist, 1999, Saal et al., 2007, Jackson et al., 2008). The samples of array 1, in terms of volatile/refractory element variations, are mainly distinguishable in terms of F/Nd and H2O/Ce ratios with the HHe group falling at higher H2O/Ce (202 ± 26) and lower F/Nd (20.1 ± 0.8) than the ITE group (H2O/Ce = 174 ± 24, F/Nd = 24.9 ± 0.7), while showing similar S/Dy and Cl/K ratios (HHe: S/Dy = 218 ± 21, Cl/K = 0.057 ± 0.009, ITE: S/Dy = 226 ± 20, Cl/K = 0.048 ± 0.011) (Peterson et al., submitted to Journal of Petrology). Volatile elements are often reported as ratios between a given volatile element and a refractory element of similar incompatibility (Nb, Ce, Nd, K, and Dy with CO2, H2O, F, Cl, and S, respectively) to overcome the variation in concentrations during crystal fractionation and melting (Schilling et al., 1980, Michael, 1995, Michael & Cornell, 1998, Danyushevsky et al., 2000, Saal et al., 2002, Hauri et al., 2006, Workman et al., 2006). The samples of array 1 have extensively degassed CO2, but ! 208! C/3He ratios calculated for samples from the HHe group are, on average, an order of magnitude higher than MORB and among the highest yet reported for OIB samples (Marty & Jambon, 1987, Trull et al., 1993, Marty & Tolstikhin, 1998, Marty & Zimmermann, 1999, Shaw et al., 2004, Barry et al., 2014, Peterson et al., submitted to Journal of Petrology). C and 3He have similar solubilities in basaltic melt (Lux, 1987, Pan et al., 1991, Dixon et al., 1995, Jendrzejewski et al., 1997) and, therefore, will not be greatly fractionated during degassing. The high C/3He ratios, as a result, imply that the HHe end-member of array 1 comes from a carbonated mantle source. The samples of array 1 with 3He/4He ratios above 22Ra have lower Pb and Sr at similar Nd isotope ratios than samples with 3He/4He ratios less than 22 Ra (ITE group), consistent with previously measured isotopes for Fernandina and Sierra Negra/Cerro Azul whole rock samples (Fig. 3) (White et al., 1993, Kurz & Geist, 1999, Saal et al., 2007). The second array (Array 2) ranges between the compositions of Genovesa and Pinta Is./Wolf, Darwin lineaments. The array is subdivided into the incompatible trace element depleted (ITD) group similar to MORB and an enriched composition (different from the ITE group) that displays the characteristics of the Pinta-Wolf-Darwin mantle end-member (WD group). The samples of array 2 range from trace element depleted to trace element enriched with the WD group having higher Th/La ratios than either the ITE or HHe groups (Fig. 2; high Th/La ratios are a characteristic of Pinta; Saal et al., 2007). The WD group has a F/Nd (23.5 ± 0.4), Cl/K (0.066 ± 0.006) and S/Dy ratio (255 ± 40) similar to those of Sierra Negra/Cerro Azul but with a significantly higher H2O/Ce ratio (230 ± 35). In comparison, secondary processes have significantly affected the volatile content of the ITD glasses with only few samples showing volatile/refractory element ! 209! ratios close to those previously reported for MORB (Peterson et al., submitted to Journal of Petrology). The samples of the ITD group have isotope ratios similar to MORB while the WD group display elevated Δ7/4 and Δ8/4 and low 3He/4He ratios (Fig. 3), consistent with the Pinta-Wolf-Darwin mantle end-member in the Galapagos and other plumes that contain an EM component (e.g. Stracke et al., 2003, Workman et al., 2006, Kendrick et al., 2012). SAMPLE PREPARATION AND METHODS For microbeam analysis, individual samples were handpicked under a microscope to select glass that is free of alteration, mounted in epoxy and doubly polished to a thickness that created a large pool of glass (thickness variation = 30-100 μm) to ensure the presence of a crystal free path for the beam. Photographs were taken of wafered glass in reflected, plane-polarzied and cross-polarized light. Fe K-edge XANES spectra were collected at station X26A (bending magnet) at the National Synchrotron Light Source (NSLS), Brookhaven National Lab (BNL) in two sessions. NSLS operates at 2.8 GeV and 300 mA and we achieved energy selection using a Si (311) channel-cut monochromator. XANES spectra were collected in fluorescence mode using a 9x5 μm spot size. A minimum of 3 spectra were taken on independent spots per Galapagos glass sample to allow for a quantitative assessment of reproducibility and precision. Each spectra was inspected for evidence of crystal interference, which, if identified, resulted in rejection of that spectra from further analysis. Centroid precisions are determined by simultaneously fitting the baseline with a linear function, constrained to have a positive slope, and a damped harmonic oscillator (DHO) function and fitting the ! 210! pre-edge peaks with two Gaussian functions over a 7110-7118 eV sub-sample of the pre- edge region. To correct for monochromator energy drift over the course of each session, multiple spectra were taken of reference glass LW_0. We define centroid energy for LW_0 as 7112.3 and model energy drift as a linear function between each analysis. Use of the drift monitor allows for the direct comparison of centroid and Fe3+/ΣFe determinations between beam sessions. Further details of spectral collection and fitting are given in Cottrell et al. (2009). We are able to determine Fe3+/ΣFe ratios of Galapagos glasses by referencing the drift-corrected centroid (area-weighted energy of the pre-edge 1s ! 3d multiplet) of individual analyses to a calibration curve constructed from experimental glasses. Experimental basaltic glasses were equilibrated at 1 atm at multiple oxygen fugacities from 2 log units below to 4.5 log units above the Quartz-Fayalite-Magnetite (QFM) buffer (Kress & Carmichael, 1991). 1σ uncertainty of Fe3+/ΣFe ratios of Galapagos glasses averaged ± 0.0037. RESULTS The 27 glass chips from the Galapagos prepared for analysis range in Fe3+/ΣFe ratios from 0.153-0.194 (Fig. 4 and presented in table 1). The six ITD glasses with MORB-like compositions have an average Fe3+/ΣFe ratio of 0.162 ± 0.003, which is comparable to ratios previously reported for MORB glass (average Fe3+/ΣFe = 0.16) (Bezos & Humler, 2005, Kelley & Cottrell, 2009, Cottrell & Kelley, 2011, Le Voyer et al., 2015). Of the remaining 21 samples, the 6 samples from the WD group (characterized by high Th/La ratios, Δ7/4 and Δ8/4 and low 3He/4He ratios) have an average Fe3+/ΣFe = 0.161 ± 0.008, ! 211! similar to MORB, while the one sample from Wolf volcano has the highest measured Fe3+/ΣFe ratio = 0.194. In comparison, the 9 samples from the ITE group (characterized by an enriched trace element and isotopic composition similar to Sierra Negra/Cerro Azul) and the 5 samples from the HHe group (characterized by 3He/4He ratios above 22Ra) are both slightly more oxidized than MORB at comparable MgO contents (Fe3+/ΣFe = 0.178 ± 0.007, and Fe3+/ΣFe = 0.182 ± 0.005 respectively). These values are most similar to the Fe3+/ΣFe ratios of back-arc basin basalts and Mariana trough basalts (0.171 ± 0.011) while being less oxidized than values reported for volcanic front basalts (0.236 ± 0.021) (Kelley & Cottrell, 2009, Kelley & Cottrell, 2012, Brounce et al., 2014). Using the empirical expression of Sack et al. (1980) and the regression coefficients of Kress and Carmichael (1991), Fe3+/ΣFe ratios of a basalt can be related to ƒO2 for a specified temperature and composition. We calculated magmatic ƒO2 values for the glasses of this study at a pressure of 2 kbars (determined from the average pressure of entrapment of high MgO melt inclusions from the Galapagos measured by Koleszar et al. (2009)) and temperature of the melt calculated using an equation provided in Wallace and Carmichael (1992), that uses the predictable variation in Al2O3 and MgO content of a melt undergoing crystal fractionation, with an included correction for the effect of H2O content of the liquidus temperature (Medard & Grove, 2008). Resulting ƒO2 values varied from ΔQFM = +0.02 to +0.74 (ΔQFM referring to the log unit difference in ƒO2 of a sample from the quartz-fayalite-magnetite buffer). Similar to Fe3+/ΣFe ratios, this is slightly more oxidized than the range found for MORB (-0.25 - +0.4; Christie et al., 1986, Bezos & Humler, 2005, Cottrell & Kelley, 2011, Cottrell & Kelley, 2013, Le Voyer et al., ! 212! 2015), but less oxidized than the range stated for arc mantle (+1 - +1.8; Kelley & Cottrell, 2009, Kelley & Cottrell, 2012, Brounce et al., 2014). There is a large variation in Fe3+/ΣFe ratios within each compositional group. Due to the correlation between Fe3+/ΣFe ratios and MgO, some of this variation can be attributed to crystal fractionation. Some of this variation, however, occurs at the same or over a small range of MgO contents. This implies that the Fe3+/ΣFe ratios of the glasses of this study are varying in terms of degree of melting, sampling other shallow level processes, such as degassing or assimilation of shallow lithospheric material, or there is an inherent heterogeneity in mantle source redox. DISCUSSION Magmatic Oxidation during differentiation and partial melting Magmatic differentiation can cause changes in Fe3+/ΣFe ratios because of the differences in partitioning behavior of Fe2+ and Fe3+ for different crystallizing phases (e.g. Sugawara, 2001, Lundgaard & Tegner, 2004, Mallmann & O'Neill, 2009). There is a clear increase in Fe3+/ΣFe ratios with decreasing MgO contents for the glasses of this study (Fig. 4). This is consistent with partitioning Fe2+ over Fe3+ in early crystallizing phases such as olivine and clinopyroxene and therefore concentrating Fe3+ in the melt. At a given MgO content, however, there is a large range in Fe3+/ΣFe ratios between and within the different groups of glasses. We have calculated possible liquid lines of descent (LLD) using PetroLog3 (Danyushevsky & Plechov, 2011) in order to assess the extent to which fractional crystallization can change Fe3+/ΣFe ratios. ! 213! For this modeling we assume isobaric fractional crystallization occurring at 2 kbars of pressure (Koleszar et al., 2009) (varying the pressure between 0.5-4 Kbars does not change our interpretation of the results). We also use the compositions of both an olivine-hosted melt inclusion taken from Fernandina and one from Santiago (composition presented in Table 2 and reported in Koleszar et al. (2009)) in order to generate the LLD of an enriched melt and a depleted melt respectively. The water content and minor/trace element enrichment of the Fernandina mantle component is higher than the depleted mantle component in the Galapagos (e.g. White et al., 1993, Koleszar et al., 2009, Peterson et al., submitted to Journal of Petrology). This will change the overall trajectory of the LLD, especially because water can have a large effect on plagioclase and clinopyroxene crystallization. We used 2 different compositions to better bracket the data. We ran the model as a system closed to oxygen. Implicit in the assertion that fractional crystallization can change Fe3+/ΣFe ratios is the assumption that the system has remained closed to oxygen exchange with the environment. This assumption is supported by a global study of MORB glass with well-defined LLDs within single suites of data, that are well reproduced by fractional crystallization of olivine ± plagioclase ± clinopyroxene in a system closed to oxygen (Cottrell & Kelley, 2011, Cottrell, personal communication). Finally, we assume D(olivine)Fe3+ = 0, D(plagioclase)Fe3+ = 0.3, D(clinopyroxene)Fe3+ = 0.45, based on current experimentally determined partition coefficients for Fe3+ (Sugawara, 2001, Lundgaard & Tegner, 2004, Mallmann & O'Neill, 2009). The results of the modeling show that despite the differences in the LLD of an enriched versus a depleted melt, fractional crystallization over a decrease in MgO content of ~9.3-5.0 wt% is only predicted to increase Fe3+/ΣFe ratios by 0.03. For the ITE and WD groups (the ! 214! groups with the largest variation in Fe3+/ΣFe ratios), the samples of each group have MgO contents within 2 wt. % and yet cover the same range in Fe3+/ΣFe ratios of ~0.03 and ~0.02 respectively. This variation in Fe3+/ΣFe ratios therefore, cannot be explained by crystal fractionation of olivine ± plagioclase ± clinopyroxene alone. Crystallization of Fe-oxides such as magnetite (Fe3O4) can have a large effect on Fe3+/ΣFe ratios. Although magnetite has a mixed valence composition, it is dominantly ferric (2Fe3+: 1Fe2+); therefore, crystallization of magnetite in a system closed to oxygen would lead to a decrease in Fe3+/ΣFe ratios of the melt. For the glasses of this study, however, magnetite is not predicted to saturate before MgO ~6 wt%, which is only applicable to the ITE and HHe groups. These groups do not show the other predicted effects of magnetite fractionation, such as a strong decrease in TiO2 or FeO* (e.g. Berndt et al., 2005, Jenner et al., 2010, Zimmer et al., 2010). Therefore we conclude that magnetite fractionation cannot explain the range in Fe3+/ΣFe ratios that we see for samples at similar MgO contents. The samples of this study have experienced different levels of crystal fractionation, so to account for these effects we recalculate Fe3+/ΣFe ratios to an arbitrary MgO content = 8 wt%, with Fe3+/ΣFe(8) being analogous to Na2O(8) in the style of Klein and Langmuir (1987). Because the samples of this study do not define a well-constrained LLD for Fe3+/ΣFe ratios, we use a linear regression of the data presented here in addition to Fe3+/ΣFe ratios of MORB samples measured with the XANES techniques described in Cottrell et al. (2009) to define the slope of the correction for all of the samples (Cottrell et al., 2009, Cottrell & Kelley, 2011, Le Voyer et al., 2015; Cottrell and Kelley, unpublished data). The slope is approximately equivalent to that defined by olivine only ! 215! fractionation. MORB data can be used to define the regression because they are well modeled by the sequence olivine ± plagioclase ± clinopyroxene, similar to MORB and EMORB samples, although the isotopic and trace element enriched groups in the Galapagos do show a slightly different trajectory for the LLD from MORB samples (as explored in Peterson et al. (submitted to Journal of Petrology)). Therefore a regression solely defined by Galapagos plume samples should have a similar slope to that defined by the MORB data from the literature. We also use a single slope for all the groups of this study because even though the predicted LLD of an enriched versus a depleted melt defines two slightly differently sloped trends, the partition coefficients for Fe3+ are defined by a limited number of studies and there is a large uncertainty in the predicted LLDs. Also, the difference in Fe3+/ΣFe(8) ratios calculated using the 1 slope defined by the data of this study in combination with MORB data reported in the literature and the 2 slopes defined by predicted LLDs is generally less than 2.5%. This ultimately results in a primary fO2 difference of 0.18 log units, which is within the 0.5 log unit uncertainty expected using the method of Kress and Carmichael (1991) to predict oxygen fugacity. Fe3+/ΣFe ratios have not been found to vary with degree of melting in other settings despite the fact that Fe3+ is more incompatible than Fe2+ during melting (e.g. Bezos & Humler, 2005, Cottrell & Kelley, 2011, Dauphas et al., 2014). Fe3+/ΣFe ratios within the groups of this study do not vary with La/Yb ratios, implying that Fe3+/ΣFe ratios are not varying with degree of melting (not shown). La/Yb ratios are sensitive to the degree of melting due to the large differences in the incompatibility of La relative to Yb. The one sample collected just north of Isabela with the composition of Wolf Volcano shows the highest Fe3+/ΣFe ratio (0.194) at a given MgO content. This sample is ! 216! characterized by low Th/La ratios and more depleted Sr, Nd and Pb isotopic compositions than the compositions of the other volcanoes of Isabela, but with higher Nb/Zr, Sm/Yb and Ti/Gd ratios than MORB. The trace elements and isotopic ratios are consistent with a low degree of melting in the garnet stability field of a depleted mantle (Geist et al., 2005, Saal et al., 2007, Peterson et al., submitted to Journal of Petrology). The high Fe3+/ΣFe ratio of the one sample with the composition of Wolf Volcano may be evidence that low degrees of melting can have a measurable effect on the Fe3+/ΣFe ratios that has not been previously recognized. Alternatively, the high Fe3+/ΣFe ratio of Wolf Volcano may be evidence for the presence of an unusual depleted source in the plume that is more oxidized than depleted upper mantle or this sample may have been altered after emplacement. However, it is difficult to constrain the origin of the high Fe3+/ΣFe ratio because only one sample with this composition has been measured, and more measurements will thus be required to explain the ƒO2 of the mantle source contributing to Wolf Volcano. The role of volatiles in magmatic oxidation Volatiles, such as sulfur, have the potential to play a role in determining magmatic oxidation. For example, the large electronic difference between oxidized sulfate (S+6) and reduced sulfide (S2-) means that, despite its low abundance, sulfur has the potential to drive changes in redox through the Fe-S redox couple: 8Fe2+ + S+6 " 8Fe3+ + S2- (1) ! 217! The speciation of S in the melt will determine whether sulfide fractionation or degassing can change the Fe3+/ΣFe ratios. For the estimated range of oxygen fugacity in ocean island basalts (ΔQFM = 0 - +2; Ballhaus (1993)), S is dissolved in basaltic melt as a mixture of S2- and SO42- (e.g. Metrich et al., 2009, Jugo et al., 2010). Magnetite fractionation has been shown to shift equation 1 to the right and cause the initiation of sulfide fractionation, which subsequently produces Fe3+ (Jenner et al., 2010). In their analysis of the volatile contents of the glasses of this study, Peterson et al. (submitted to Journal of Petrology), determined that the melt was just saturated with a sulfide liquid to sulfide undersaturated (using the model of Liu et al. (2007) and an Fe3+/ΣFe ratio = 0.17 to determine sulfur content at sulfide saturation (SCSS)). Recalculation of the SCSS using the Fe3+/ΣFe ratios measured in this study maintains this result. Only glasses that are sulfide undersaturated were chosen for XANES analysis. Sulfur degassing also has the potential to affect the oxidation state of the melt. The 2 main species of S that degas are SO2 (S4+) and H2S (S2-), though for the glasses of this study at the pressure at which the melt is in equilibrium with a CO2-H2O vapor phase (23-472 bars calculated using the model of Dixon et al. (1995) and discussed in Peterson et al. (submitted to Journal of Petrology)) we would expect the gas phase to be mainly composed of SO2 (Carroll & Webster, 1994, Symonds et al., 1994, Moretti et al., 2003, Oppenheimer, 2003, Wallace, 2005). Degassing, as a result, will cause a redistribution of electrons in the melt and change the Fe3+/ΣFe ratios. Peterson et al. (submitted to Journal of Petrology), when comparing S/Dy ratios to the H2O-CO2 pressure of saturation, found that S begins to degas at approximately 400 bars of pressure. When S/Dy ratios are ! 218! compared to Fe3+/ΣFe(8) ratios there is a distinct reduction trend with decreasing S contents (Fig. 5). If S is mainly degassing as SO2, this implies that S is mainly dissolved in the melt as S2- consistent with the parameterization of Jugo et al. (2010) for the magmatic ƒO2’s calculated for the Fe3+/ΣFe ratios of this study. Degassing of other volatiles (H-C-O-Cl species) could also potentially affect Fe3+/ΣFe ratios. Degassing of H2 could drive the oxidation of a basaltic melt through the redox couple: 3Fe2+O + H2O(melt) " H2(fluid) +Fe2+O# Fe23+O3(Magnetite) (2) This reaction will cause the melt to oxidize as H2 escapes the system. For the vapor phase to have a significant proportion of H2, the melt must be much more reduced (< ΔQFM -1) than the magmatic ƒO2 calculated for the glasses of this study and, in general, predicted for OIB and MORB (Carmichael, 1991, Cottrell et al., 2009, Zhang & Duan, 2009). The process is also self-limiting as increasing the oxidation of the melt will drive the proportion of H2 in the fluid down. Finally, this reaction requires the precipitation of magnetite, for which there is no evidence in the glasses of this study. As a result, it is unlikely that H2 degassing is affecting the Fe3+/ΣFe ratios. It is also unlikely that CO2 degassing is affecting the melt Fe3+/ΣFe ratios. At QFM > -1, graphite is not stable in the melt, therefore C will be speciated as carbonate (CO32-) (e.g. Stagno et al., 2013). As a result, CO2 degassing will be redox neutral. This is consistent with the fact that CO2 degassing is not correlated with changes in oxidation in the glasses of this study (such as ! 219! variation in CO2/Nb with Fe3+/ΣFe ratios, not shown). Cl degassing also has the potential to oxidize the melt through the redox couple (Bell & Simon, 2011): Fe2+O + HCl(fluid) " H2O(melt) +FeCl(fluid) (3) However, there is no evidence of Cl degassing from the melts of this study such as a correlation between Cl/K ratios and CO2-H2O saturation pressure (not shown). Therefore, it is unlikely that Cl degassing is affecting Fe3+/ΣFe ratios. Magmatic Oxidation during assimilation of hydrothermally altered material Reactions with seawater have been shown to oxidize the crust (e.g. Alt, 1995), therefore assimilation of hydrothermally altered crust and/or seawater has the potential to oxidize melt. There is a clear correlation between increasing values of Sr/Sr* >1 (Sr anomaly on a primitive mantle normalized diagram = Srpm/[(Ndpm + Cepm)/2] (Mcdonough & Sun, 1995)) and increasing Fe3+/ΣFe(8) ratios (Fig. 6). Sr/Sr* can be used as a measure for plagioclase fractionation (<1) and assimilation of plagioclase cumulates (>1). Sr/Sr* is shown to correlate with increasing volatile/refractory element ratios and thus is a good indicator of the assimilation of hydrothermally altered crust in the Galapagos (Peterson et al., submitted to Journal of Petrology). The correlation between Fe3+/ΣFe(8) ratios and Sr/Sr* is therefore consistent with assimilation of hydrothermally altered crust. The sample with the highest Cl/K ratio (PL02-13-25; Cl/K = 0.77) is also indicated in Figure 6. This sample has most probably assimilated brine or seawater in addition to hydrothermally altered crust (Kent et al., 1999, Stroncik & Haase, 2004, ! 220! Peterson et al., submitted to Journal of Petrology). This sample is not significantly more oxidized than the other samples (whose Cl/K ratios range from 0.025 - 0.069). In addition, there is no evidence for a correlation between Fe3+/ΣFe(8) ratios and Cl/K ratios in the glasses of this study. Therefore, it is more likely that assimilation of hydrous components in the crust are causing the increase in Fe3+/ΣFe ratios as opposed to assimilation of seawater or brine. Mantle source variation in Fe3+/ΣFe ratios After accounting for the major processes that can affect Fe3+/ΣFe ratios in the melt, the different groups of glasses still show variation outside of analytical uncertainty (Fig. 7). The ITD and WD groups have similar Fe3+/ ΣFe(8) ratios to each other (ITD = 0.162 ± 0.003 and WD = 0.164 ± 0.006) and MORB values (Bezos & Humler, 2005, Cottrell & Kelley, 2011, Le Voyer et al., 2015), while the HHe group (Fernandina source) are more oxidized (HHe = 0.175 ± 0.006, maximum value = 0.179) than MORB, but less oxidized than reported for arc basalts (Kelley & Cottrell, 2009, Kelley & Cottrell, 2012, Brounce et al., 2014). The ITE group (Sierra Negra/Cerro Azul) have the largest range in Fe3+/ ΣFe(8) ratios (average = 0.169 ± 0.008, maximum = 0.177 ), which ranges to more oxidized than the ITD and WD groups but more reduced relative to the HHe group. This could suggest mixing with melts from the WD mantle source as Fe3+/ΣFe(8) ratios broadly correlate with Δ7/4 values. Both the Pinta (WD group) and Sierra Negra/Cerro Azul mantle end-members (ITE group) are hypothesized to contain recycled components (e.g. White et al., 1993, Harpp & White, 2001, Schilling et al., 2003, Harpp et al., 2014). The Pinta component ! 221! has an isotopic signature similar to an EM mantle end-member. The Sierra Negra/Cerro Azul end-member is either the product of a mixture, between the Floreana, Fernandina and Pinta end-members, or alternatively a unique component with an intermediate isotopic composition. The differences in Fe3+/ ΣFe(8) ratios and the correlation between Fe3+/ ΣFe(8) ratios and Pb isotopes for the ITE group between an oxidized end-member and a reduced end-member similar to the WD group suggests that the source material of these two end-members is intrinsically different and that there are long-lived, heterogeneous redox domains in the mantle, consistent with the conclusions of Cottrell and Kelley (2013). The HHe group is the most oxidized group within the glasses of this study. The characteristics of the source material for the HHe group (Fernandina mantle end-member) are consistent with a mantle component that is generally attributed to less-degassed, primitive lower mantle material (e.g. Allègre et al., 1983, Farley et al., 1992, Hart et al., 1992, Hanan & Graham, 1996). Peterson et al. (submitted to Journal of Petrology) also showed that the high 3He/4He end-member in the Galapagos is more carbonated than MORB and other oceanic islands using the high C/3He ratios of this end-member. The more oxidized nature of the HHe group may be due to the higher carbon content of the mantle source. Stagno et al. (2013) demonstrates with experimental and natural data that carbon hosted as graphite or diamond in the upwelling mantle will oxidize to produce carbonate melt through the reduction of Fe3+. As a result, the depth of carbonate melt formation is dependent on the Fe3+/ΣFe ratio of the bulk rock. In the Galapagos, lavas with the highest 3He/4He ratios are located above a seismic low velocity zone that is at a depth of 200-250km, which is estimated to be the result of the initiation of carbonate ! 222! melting (Villagomez et al., 2014). At this depth, a mantle containing 30 ppm C (estimated C content of the MORB mantle source) would require a Fe3+/ΣFe ratio of at least 6-7% to stabilize a carbonated melt phase (Stagno et al., 2013). Subsequent oxidation of the carbon would then cause a ~1% reduction in the Fe3+/ΣFe ratio of the mantle source. Increasing the carbon content of a source without changing the Fe3+/ΣFe ratio would therefore result in a more reduced melt. If the Fe3+/ΣFe(8) ratios of the glasses of this study are representative of the mantle, then to end with an Fe3+/ΣFe ratio more oxidized than MORB for a source that contains an order of magnitude more carbon (as observed in the HHe group) would require an even more oxidized Fe3+/ΣFe ratio than the 6-7% hypothesized for MORB by Stagno et al. (2013) and much more than the 0.036 estimated for the primitive mantle from peridotite and pyroxenite xenoliths (Canil et al., 1994). CONCLUSIONS Galapagos glass compositions show a range in Fe3+/ΣFe ratios from MORB-like to more oxidized than MORB, within the range of back-arc basin basalts and Mariana trough basalts, but less oxidized than volcanic front basalts. Some of this variability can be attributed to crystal fractionation, but not variable degrees of melting, consistent with previous studies. There is also evidence that S degassing causes magmatic reduction, while assimilation of hydrothermally altered crust is causing magmatic oxidation. After accounting for these shallow level processes, the different isotopic groups still vary from each other in their Fe3+/ΣFe ratios. The Fe3+/ΣFe ratios of this study result in a magmatic ƒO2 for the Galapagos Archipelago that varies from ΔQFM = +0.16 - +0.74, intermediate ! 223! to the MORB and arc systems. The ITD group has a Fe3+/ΣFe ratio within range of what has been found for MORB. The WD group was also found to posses a Fe3+/ΣFe ratio within the range of MORB while the ITE group is generally more oxidized. The WD and ITE groups have both been hypothesized to contain recycled components in their mantle sources, but the difference in Fe3+/ΣFe ratios of these groups and the correlation between Fe3+/ΣFe ratios and Pb isotopes suggest that their mantle source compositions are intrinsically different. The HHe group is the most oxidized group of this study with the highest Fe3+/ΣFe ratios. This in combination with the more carbonated nature of the source could be evidence that less degassed, primitive lower mantle (the source of the high 3He/4He ratios of the HHe group) is inherently more oxidized than depleted upper mantle. The new data suggest that the observed Fe3+/ΣFe ratios reflect mantle source variations instead of simply lithospheric processes. This has important implications for the preservation of long-lived mantle heterogeneities in oxidation state and simultaneously implies that the mantle is neither a well-mixed homogenous source nor a well-buffered system. ACKNOWLEDGMENTS This work was supported by the National Science Foundation Graduate Research Fellowship (Grant No. DGE-1058262 to M.E.P.) and the National Science Foundation Division of Ocean Sciences (Grant No. 0962195).We would like to thank K. Shimizu, E. ! 224! Chin and M. 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Mantle flow and multistage melting beneath the Galapagos hotspot revealed by seismic imaging. Nature Geoscience 7, 151-156, Doi 10.1038/Ngeo2062. ! 239! Villagomez, D. R., Toomey, D. R., Hooft, E. E. E. & Solomon, S. C. (2007). Upper mantle structure beneath the Galapagos Archipelago from surface wave tomography. Journal of Geophysical Research-Solid Earth 112, Artn B07303, Doi 10.1029/2006jb004672. Villagomez, D. R., Toomey, D. R., Hooft, E. E. E. & Solomon, S. C. (2011). Crustal structure beneath the Galapagos Archipelago from ambient noise tomography and its implications for plume-lithosphere interactions. Journal of Geophysical Research-Solid Earth 116, Artn B04310 Doi 10.1029/2010jb007764. Wallace, P. & Carmichael, I. S. E. (1992). Sulfur in Basaltic Magmas. Geochimica Et Cosmochimica Acta 56, 1863-1874, Doi 10.1016/0016-7037(92)90316-B. Wallace, P. J. (2005). Volatiles in subduction zone magmas: concentrations and fluxes based on melt inclusion and volcanic gas data. Journal of Volcanology and Geothermal Research 140, 217-240, Doi 10.1016/J.Jvolgeores.2004.07.023. Werner, R. (2002). Cruise report SO158 (MEGAPRINT): Multidisciplinary Examination of Galapagos Plume Ridge Interaction. GEOMAR-Rep 104, 1-53. White, W. M. (2010). Oceanic Island Basalts and Mantle Plumes: The Geochemical Perspective. Annual Review of Earth and Planetary Sciences, Vol 38 38, 133-160, Doi 10.1146/Annurev-Earth-040809-152450. White, W. M., Mcbirney, A. R. & Duncan, R. A. (1993). Petrology and Geochemistry of the Galapagos-Islands - Portrait of a Pathological Mantle Plume. Journal of Geophysical Research-Solid Earth 98, 19533-19563. Workman, R. K., Hauri, E., Hart, S. R., Wang, J. & Blusztajn, J. (2006). Volatile and trace elements in basaltic glasses from Samoa: Implications for water distribution ! 240! in the mantle. Earth and Planetary Science Letters 241, 932-951, 10.1016/j.epsl.2005.10.028. Zhang, C. & Duan, Z. H. (2009). A model for C-O-H fluid in the Earth's mantle. Geochimica Et Cosmochimica Acta 73, 2089-2102, Doi 10.1016/J.Gca.2009.01.021. Zimmer, M. M., Plank, T., Hauri, E. H., Yogodzinski, G. M., Stelling, P., Larsen, J., Singer, B., Jicha, B., Mandeville, C. & Nye, C. J. (2010). The Role of Water in Generating the Calc-alkaline Trend: New Volatile Data for Aleutian Magmas and a New Tholeiitic Index. Journal of Petrology 51, 2411-2444, Doi 10.1093/Petrology/Egq062 ! 241! FIGURE CAPTIONS Figure 1. Map of the Galapagos archipelago with locations of sample dredges and prominent features labeled. Scale bar and bathymetry color scale are included the left side on the map. Array 1 consists of samples that range between the compositions of Fernandina Island and Sierra Negra/Cerro Azul volcanoes that are further divided into HHe and ITE groups based on samples with 3He/4He ratios above and below 3He/4He = 22 Ra, respectively. Array 2 consists of samples that range between the compositions of Genovesa Island and Pinta, Wolf and Darwin Islands that have been divided into the ITD group (samples with more MORB-like compositions) and the WD group (samples characterized by high Th/La ratios and high Δ7/4, Δ8/4). Samples that have a composition similar to Wolf Volcano on Isabela are shown separately. Map created using GeoMapApp http://www.geomapapp.org (Ryan et al., 2009). Figure 2. Th/La ratios versus A.) Ti/Gd ratios and B.) Sm/Yb ratios of individual glasses plotted over fields that represent ranges of previously measured whole rock and glass samples from the Galapagos Archipelago (Cullen et al., 1989, Vicenzi et al., 1990, Bow & Geist, 1992, White et al., 1993, Geist et al., 2002, Saal et al., 2007, Gibson et al., 2012). Errors shown are 2σ standard error for individual glasses. Inset shows larger submarine glass data set analyzed in Peterson et al., (submitted to Journal of Petrology). 143 Figure 3. (A-B) Nd/144Nd and (C-E) Δ7/4 versus 87Sr/86Sr and 3He/4He. Notice glass samples with 3He/4He > 22Ra correspond to the isotope composition indicative of ! 242! Fernandina whole rocks. Notice also that the WD and ITD Groups define a compositional range from near Genovesa to near Pinta-Wolf-Darwin Islands (isotopic data for individual glasses reported in Harpp & White, 2001, Schilling et al., 2003, Geist et al., 2008, Peterson et al., submitted to Journal of Petrology). The grey fields are defined by previously published isotope data on whole rocks basalts from different islands (Cullen et al., 1989, White et al., 1993, Reynolds & Geist, 1995, Blichert-Toft & White, 2001, Schilling et al., 2003, Saal et al., 2007, Ingle et al., 2010). Average 2σ standard error indicated by the grey plus sign in each panel. Inset shows larger submarine glass data set analyzed in Peterson et al., (submitted to Journal of Petrology). Figure 4. MgO wt% versus A.) TiO2 wt %, B.) Fe+3/ΣFe ratios, C.) Al2O3 wt% and D.) K2O wt% for individual glasses from this study. Superimposed on the data are predicted liquid lines of descent (LLD) for olivine (olv) - olv + plagioclase (plag) – olv + plag + clinopyroxene (cpx) (solid red and green lines, where black dot on the lines indicates the appearance of a new crystal phase in the fractionating mineral assemblage). Calculations were performed in PetroLog3 (Danyushevsky & Plechov, 2011) isobarically at 2 Kbars of pressure using 2 different starting compositions, which can be found in Table 2. Red LLD is calculated from an enriched Fernandina composition, while green LLD is calculated from a depleted Santiago composition. LLDs for Fe+3/ΣFe ratios calculated using mineral/melt partition coefficients Dolv = 0, Dplag = 0.35, Dcpx =0.45 (Sugawara, 2001, Lundgaard & Tegner, 2004, Mallmann & O'Neill, 2009) in a system closed to oxygen. Notice in panel B that at a given MgO content there is a large range in Fe+3/ΣFe ratios between and within different groups showing that the variation in Fe+3/ΣFe ratios ! 243! cannot be explained by differentiation alone. Grey field covers the range in Fe+3/ΣFe ratios reported for MORB (Cottrell & Kelley, 2011, Le Voyer et al., 2015). Errors are 2σ standard error for TiO2, Al2O3, and K2O and 1σ for Fe+3/ΣFe ratios. Figure 5. S/Dy ratios versus A.) Pressure of CO2-H2O vapor saturation and B.) Fe+3/ΣFe(8) ratios showing that S degassing correlates with magmatic reduction. Pressure of saturation calculated using the model of Dixon et al. (1995). Glass data shown in panel A include all the data reported in Peterson et al., (submitted to Journal of Petrology) filtered for samples affected by assimilation of hydrothermally altered material, which will independently affect S/Dy ratios. Also shown are melt inclusion data from Koleszar et al. (2009). Error for S/Dy ratios are 2σ standard error, error for Fe+3/ΣFe(8) ratios are 1σ standard error. Error for H2O – CO2 P saturation calculated from the 2σ standard error of CO2 measurements. Figure 6. Fe+3/ΣFe(8) ratios versus Sr/Sr* for individual glasses. Notice the clear trend between increasing Sr/Sr* (proxy for plagioclase assimilation) and magmatic oxidation. Also labeled is the ITD sample with the highest Cl/K ratio (PL02-13-25 collected from the eastern seamounts). Cl/K is used as an indicator of assimilation of hydrothermally altered material. While this sample is oxidized, it is not unusually oxidized in comparison to the other affected samples in the ITD group. This is consistent with assimilation of shallow, hydrothermally altered, lithospheric material causing magmatic oxidation as opposed to simply assimilating seawater. ! 244! Figure 7. Fe+3/ΣFe(8) ratios versus A.) 87Sr/86Sr, B.) 143Nd/144Nd, C.) 3He/4He, and D.) Δ7/4 for individual glasses. Small symbols indicate samples whose Fe+3/ΣFe(8) ratio has been identified as being affected by shallow processes. Notice the broad correlation between isotope signature and Fe+3/ΣFe(8) ratios. ! 245! FIGURES Figure 3-1 HHe Group 5°N ITE Group Array 1 ITD Group Array 2 4°N Galapagos Spreading Center (GSC) WD Group Wolf Volcano 3°N 2°N Wolf-Darwin Latitude lineament Marchena 1°N Pinta Genovesa NE Seamounts Santiago 0°N Fernandina Isabela San Cristobal Santa Cruz 1°S Floreana 2°S 98°W 97°W 96°W 95°W 94°W 93°W 92°W 91°W 90°W 89°W 88°W 87°W 86°W 85°W 84°W Longitude ! 246! Figure 3-2 A. Sierra Negra/ Fernandina Cerro Azul 3,200 3,200 Floreana 2,800 2,800 Ti/Gd 2,400 Ti/Gd 2,400 2,000 Wolf Vol. 1,600 2,000 GSC 1,600 2.4 B. 2.0 Wolf Vol. Fernandina Sm/Yb 2.6 1.6 2.4 Pinta Sierra Negra/ 2.2 Cerro Azul 1.2 2.0 0.8 1.8 Sm/Yb Floreana 1.6 0.02 0.06 0.10 1.4 Th/La 1.2 HHe Group ITE Group 1.0 GSC ITD Group 0.8 WD Group Wolf Volcano 0.02 0.04 0.06 0.08 0.10 0.12 Th/La ! 247! Figure 3-3 A. Floreana 2σ 2σ C. Floreana Sierra Negra/ 0.7036 Cerro Azul Sierra Negra/ 0.7034 Cerro Azul Pinta 0.7032 Fernandina Pinta Wolf/Darwin Is. 87Sr/86Sr 0.7030 Fernandina GSC 0.7028 Wolf/Darwin Is. Genovesa Wolf Vol. 0.7026 Wolf Vol. GSC Genovesa 0.7024 B. 2σ 2σ D. Fernandina Fernandina 28 24 20 3He/4He Sierra Negra/ Sierra Negra/ Cerro Azul Cerro Azul 16 Floreana 12 Floreana Wolf Vol. Wolf Vol. 8 Pinta GSC GSC Pinta 0.51290 0.51300 0.51310 -2 -1 0 1 2 3 4 143Nd/144Nd Δ7/4 2σ 2σ HHe Group 0.7036 ITE Group ITD Group 0.7034 WD Group Wolf Volcano 0.7032 87Sr/86Sr 0.7030 0.7028 0.7026 0.7024 2σ 2σ 28 24 3He/4He 20 16 12 8 0.51290 0.51300 0.51310 -2 -1 0 1 2 3 4 143Nd/144Nd Δ7/4 ! 248! Figure 3-4 A. C. 4.0 18 3.5 17 3.0 Al2O3 (wt%) TiO2 (wt%) 16 2.5 15 2.0 14 1.5 1.0 13 olv olv+plag olv+plag+cpx B. D. 0.19 HHe Group 1.0 ITE Group 0.18 ITD Group K2O (wt%) WD Group 0.8 Wolf Volcano 0.17 Fe+3/ΣFe MORB 0.6 0.16 0.4 0.15 0.14 0.2 6 7 8 9 10 5 6 7 8 9 10 MgO (wt%) MgO (wt%) ! 249! Figure 3-5 A. B. 300 250 S/Dy 200 Degassing Degassing 150 HHe Group ITE Group ITD Group WD Group 100 Wolf Volcano Fernandina melt inclusions Santiago melt inclusions 200 400 600 800 1,000 1,200 1,400 0.150 0.160 0.170 0.180 H2O - CO2 Pressure of saturation (bars) Fe+3/FeT(8) ! 250! Figure 3-6 0.19 0.18 Cl/K = 0.77 0.17 Fe+3/ΣFe(8) Assimilation 0.16 HHe Group ITE Group ITD Group 0.15 WD Group Wolf Volcano 0.14 0.4 0.6 0.8 1.0 1.2 1.4 1.6 1.8 2.0 Sr/Sr* ! 251! Figure 3-7 A. 0.7034 0.7032 87Sr/86Sr 0.7030 0.7028 0.7026 B. 0.51315 0.51310 143Nd/144Nd 0.51305 0.51300 0.51295 0.51290 HHe Group ITE Group ITD Group WD Group C. Wolf Volcano 25 20 3He/4He 15 10 D. 4 3 2 Δ7/4 1 0 -1 0.15 0.16 0.17 0.18 Fe+3/ΣFe(8) ! 252! TABLES Table 1. Fe3+/ΣFe ratios as measured by XANES and calculated deviations in ƒO2 from the Quartz-Fayalite-Magnetite Buffer (ΔQFM) for the samples of tis study Temperature Sample ID* Latitude Longitude Group Fe3+/ΣFe 1σ ΣdiXia (°C)b ΔQFMc(magmatic) PL02-12-1 -0.21 -88.65 ITD 0.163 0.001 0.228 1164 0.19 PL02-13-25 0.07 -89.04 ITD 0.166 0.005 0.208 1208 0.29 PL02-14-8 0.07 -89.12 ITD 0.160 0.003 0.221 1226 0.17 PL02-9-29 -0.47 -88.53 ITD 0.158 0.007 0.172 1245 0.25 DR70-2 0.41 -89.87 ITD 0.163 0.005 0.243 1176 0.16 DR72-A-1 0.09 -89.25 ITD 0.161 0.008 0.212 1254 0.21 D3a 0.27 -91.43 Wolf 0.194 0.002 0.227 1123 0.66 PL02-26-21 0.83 -91.29 WD 0.153 0.002 0.230 1178 0.03 PL02-28-11 1.02 -91.55 WD 0.160 0.003 0.210 1195 0.18 PL02-29-3 1.90 -92.16 WD 0.173 0.003 0.230 1172 0.35 PL02-30-1 1.59 -91.92 WD 0.166 0.007 0.223 1187 0.26 DR40-1 0.98 -90.84 WD 0.159 0.007 0.182 1191 0.23 DR41-A-1 1.24 -91.12 WD 0.153 0.001 0.229 1171 0.02 D13b -0.19 -91.55 ITE 0.165 0.005 0.197 1123 0.29 D38a -0.72 -91.55 ITE 0.179 0.006 0.256 1158 0.38 D39a -0.78 -91.58 ITE 0.180 0.001 0.238 1154 0.43 D48a -1.13 -91.45 ITE 0.182 0.002 0.238 1145 0.46 D51a -1.23 -91.16 ITE 0.192 0.002 0.178 1102 0.74 D5g 0.10 -91.64 ITE 0.178 0.003 0.233 1112 0.41 D60a -1.29 -90.99 ITE 0.179 0.003 0.193 1130 0.52 D8a -0.09 -91.76 ITE 0.175 0.004 0.162 1146 0.52 PL02-24-32 -1.29 -90.91 ITE 0.174 0.002 0.190 1125 0.44 ! 252! Temperature Latitude Longitude Group Fe3+/ΣFe 1σ ΣdiXia (°C)b ΔQFMc(magmatic) AHA20b -0.24 -91.75 HHe 0.187 0.002 0.230 1125 0.56 AHA22b -0.18 -91.76 HHe 0.180 0.002 0.210 1147 0.49 AHA23a -0.26 -91.73 HHe 0.185 0.009 0.221 1134 0.54 AHA23b -0.26 -91.73 HHe 0.186 0.005 0.229 1134 0.54 AHA26a -0.43 -91.71 HHe 0.175 0.004 0.212 1138 0.41 * Sample ID’s have been shortened to not include cruise information; IDs beginning with PL02 are from the PLUME02 cruise, DR are from the R/V SONNE cruise, D are from the DRIFT04 cruise, and AHA are from the AHA-Nemo-2 cruise. a ΣdiXi parameterizes the basalt composition: sum of coefficients (di) multiplied by the oxide mole fractions (Xi), as defined in Kress and Carmichael (1991) b Temperature of melt calculated from calculation of Wallace and Carmichael (1992) with a correction for the effect of water content on the liquidus from Medard and Grove (2008) c ΔQFM calculated using the regression coefficients of Kress and Carmichael (1991) at a pressure of 2 Kbars and calculated temperature of the melt ! 253! Table 2. Fractional crystallization model Starting Composition Enriched Depleted Composition Compositionb a Major Elements (wt%) SiO2 47.48 46.46 TiO2 2.29 1.34 Al2O3 16.38 14.36 FeO* 9.16 11.77 MnO 0.12 0.20 MgO 9.38 11.86 CaO 11.58 11.50 Na2O 2.25 2.14 K 2O 0.33 0.08 P 2O 5 0.26 0.07 H2O (wt%) 0.68 0.12 Fe3+/ΣFec 0.165 0.141 Pressured 2kbars Temperature 1200 a Composition of melt inclusion AHA D25C-3-43 reported in Koleszar et al., (2009) b Composition of melt inclusion STG06-29-06 reported in Koleszar et al., (2009) c derived by taking an average (for the enriched composition) of the HHe, ITE, and WD group Fe3+/ΣFe ratios and (for the depleted composition) taking an average of the ITD group glasses and correcting this back to the MgO content of the inclusions using the slope of a linear regression through the data of this study in addition to Fe3+/ΣFe ratios reported for MORB in Cottrell and Kelley (2011) d Models used in PETROLOG: Ol, CPX, PLAG- Danyushevskey, 2001, Melt density- Lange and Carmichael, 1987, Melt viscosity- Bottinga & Weill 1972. Oxygen fugacity set at QFM and then run closed for oxygen ! 254! Chapter 4 Oxygen isotopes in lavas from the Galapagos Archipelago: evaluating contributions from ancient and present-day crustal sources M. E. Peterson1, Z. Wang2, A. E. Saal1, M. D. Kurz3, J. M. Eiler4 1 Department of Earth, Environmental and Planetary Sciences, Brown University, Providence, Rhode Island, 02912, United States 2 Department of Earth and Atmospheric Science, The City College of New York, New York, New York, 10031, United States 3 Department of Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, 02543, United States 4 Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, California, 91125, United States ! 255! ABSTRACT New measurements of oxygen isotopes in olivine phenocrysts collected across the Galapagos Archipelago show a larger range in δ18O values than previously reported. Fernandina, Floreana, and Pinta, which represent the main isotopic end-members of the Galapagos hotspot, and Cerro Azul and Santiago, which have more intermediate compositions, have δ18O values = 5.02 ± 0.08‰, within the accepted range of the oxygen isotope value of upper mantle olivine. Therefore, δ18O values do not correlate with radiogenic isotopes, but there is an overall trend of increasing Sr/Sr* and Ba/Nb ratios with δ18O values for Sierra Negra, Darwin Volcano, Santa Cruz and San Cristobal. This suggests that the high δ18O values of San Cristobal (5.34 ± 0.05‰) and the large ranges in δ18O values for Sierra Negra and Darwin volcanoes and Isla Santa Cruz are the result of the assimilation of plagioclase cumulates. Olivines from Sierra Negra, Alcedo, Darwin, Wolf and Ecuador volcanoes with δ18O values lower than upper mantle values (4.85 ± 0.1‰), in comparison, do not show unique trace element or isotopic signatures making it difficult to unambiguously identify the occurrence of assimilation for these melts. Assimilation of a small amount of seawater (δ18O = 0‰) and/or hydrothermally altered gabbro without a cumulate component (δ18O = +2.5/4.0‰), however, can produce the low values for δ18O without a large change in the trace element or isotopic composition of the melt. These results show that while the three end member components of the Galapagos mantle have a generally homogenous δ18O value indistinguishable from “normal” upper mantle, there is a more widespread affect of lithospheric contamination in melts than previously thought. ! 256! INTRODUCTION One of the main concerns when inferring mantle compositions from lavas is distinguishing between signatures of melt/rock interaction in the shallow lithosphere from the signature of recycled material in the mantle source. Oxygen isotopes are uniquely suited to address this because many crustal materials have oxygen isotope ratios that are measurably different from those that are characteristic of the mantle due to low temperature weathering and water-rock interaction. As a result, deviations in δ18O (δ18O = 1000[18O/16Osample/18O/16OVSMOW - 1], where VSMOW is the Vienna standard mean ocean water at an 18O/16O = 0.0020052) from the mantle source value (average upper mantle olivine = 5.0-5.4‰, Mattey et al., 1994, Eiler, 2001) are inferred to indicate the presence of material that at one time resided at or near the surface. Ocean island basalts (OIB) display a wider range in δ18O values than mid-ocean ridge basalts (2.9-7.5‰ versus 5.7 ± 0.2‰; Harmon & Hoefs, 1995, Eiler, 2001). The large range in δ18O values have been explained both by meteoric water/seawater interaction with magma and interaction with lower (18O-depleted) and/or upper (18O- enriched) hydrothermally altered oceanic crust as either a magmatic contaminant sampled during a melt’s passage through the lithosphere or as a recycled component in the mantle source (e.g. Taylor & Forester, 1979, Eiler, 2001, Kent et al., 2004, Bindeman et al., 2006, Wang & Eiler, 2008, Workman et al., 2008). Correlations between oxygen isotopes and radiogenic isotopes and trace element ratios are essential in distinguishing between variable δ18O values of the mantle and shallow level contamination. Differentiating between mantle and lithospheric sources of oxygen can thus constrain the composition of ! 257! mantle plumes and also reveal details of the magmatic plumbing system transporting magma to the surface. The Galapagos Archipelago is the eruption site of a mantle plume that has 4 isotopically distinct mantle end-members (e.g. White et al., 1993, Blichert-Toft & White, 2001, Harpp & White, 2001, Saal et al., 2007). These components are generally isolated to the northern, western, southern, and central/eastern regions of the archipelago. The western component is characterized by the lavas of Fernandina and is distinguishable by high 3He/4He isotopic ratios (Graham et al., 1993, Kurz & Geist, 1999, Kurz et al., 2009). This component is generally compared to a relatively volatile-rich primitive mantle component (Kurz et al., 1982, Farley et al., 1992, Hart et al., 1992, Hanan & Graham, 1996). There is a MORB-like component that is concentrated in the central and eastern region and anchored by the lavas of Genovesa, though it can be found across the archipelago (Harpp et al., 2002, Harpp et al., 2003, Geist et al., 2005). The northern and southern components (illustrated by the lavas of Pinta and Floreana respectively) display enrichments in trace elements in addition to isotopic signatures generally associated with ancient recycled material in the source. Floreana is most commonly compared to a HIMU 238 component (high-μ = high U/204Pb) due to enrichments in Pb isotopes, despite radiogenic Sr and Hf isotopes which are atypical of a HIMU source (e.g. Harpp et al., 207 2014). Pinta, in comparison, has low Nd and Hf isotopes with elevated Pb/204Pb and 208 Pb/204Pb at a given 206Pb/204Pb ratio, which are characteristics generally associated with an Enriched Mantle (EM) component. Lavas collected from the other islands and the submarine platform in the Galapagos have intermediate compositions suggesting widespread mixing of these 4 end-members throughout the archipelago (e.g. Harpp & ! 258! White, 2001, Gibson et al., 2012). However, shallow lithosphere can also influence the geochemistry of erupted basalts (Peterson et al., 2014, Peterson et al., In prep, Peterson et al., submitted to Journal of Petrology). The limited previous work on Galapagos Archipelago oxygen isotopes has revealed very little regional variation in δ18O values, despite variable radiogenic isotopic signatures. This has lead to the conclusion that magmas in the Galapagos are derived from a source with homogenous δ18O values, within error of MORB, and that significant assimilation of hydrothermally altered crustal material is not occurring (Geist et al., 1995, Geist et al., 1998). Only rhyolites from Alcedo with evidence for magnetite crystallization (which will increase δ18O of the melt) and hydrothermally altered xenoliths from Cerro Azul and the plume influenced Galapagos Spreading Center (GSC) have δ18O values above (in the case of the rhyolites and Cerro Azul) and below (in the case of the GSC) what has been measured for MORB (Geist et al., 1995, Geist et al., 1998). We present new δ18O values of olivine phenocrysts taken from subaerial eruptions sampled across the Galapagos Archipelago. Results are combined with the previously published geochemistry and petrogenesis of the Galapagos lavas from which the olivine grains derive (Kurz & Geist, 1999, Saal et al., 2007). With this expanded dataset we show that there is a larger range in δ18O values in the Galapagos and evidence for a more widespread affect of contamination of melt with lithospheric material than previously suspected. METHODS AND PREVIOUS RESULTS ! 259! Samples from 12 volcanoes in the Galapagos Archipelago were selected for this analysis (Fig. 1, Islands listed in Table 1). Lavas from Fernandina, Floreana and Pinta represent the main isotopic end-members both within the 12 volcanoes sampled and in the Galapagos Archipelago (as described above, Fig. 2). Cerro Azul, Sierra Negra, Ecuador and Alcedo (volcanoes located on Isabela) are intermediate in their isotopic composition to Fernandina, Floreana, and Pinta, (White et al., 1993, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Saal et al., 2007, Kurz et al., 2009). Lavas from Darwin Volcano and Wolf Volcano on Isabela, and Santiago, which is in the central part of the archipelago, have lower Sr isotope and higher Nd isotope ratios than the other volcanoes on Isabela, with Pb isotopes that range from MORB-like to similar to Wolf and Darwin Islands and 3He/4He ratios that are slightly elevated in comparison to MORB (White et al., 1993, Kurz & Geist, 1999, Blichert-Toft & White, 2001, Harpp & White, 2001, Saal et al., 2007, Gibson et al., 2012). Santa Cruz, located in the central part of the archipelago east of Santiago, has similar isotopic characteristics to Santiago, but on average is more MORB-like (White et al., 1993, Blichert-Toft & White, 2001, Harpp & White, 2001, Saal et al., 2007). San Cristobal is an island in the southeastern end of the archipelago (Fig. 1). Lavas from San Cristobal fall over a range in geochemical compositions from similar to MORB to more intermediate, comparable to the range of 143 Santa Cruz, but with higher Sr isotopes at a given value for Nd/144Nd (Fig. 2; White et al., 1993, Saal et al., 2007). Sample ID’s refer to the lava from which olivines were picked. Geochemistry and petrogenesis of the lavas have been previously reported in the literature (Kurz & Geist, 1999, Saal et al., 2007). ! 260! Coarse olivine grains (≥ 1 mm in size) free of inclusions and alteration were hand- picked under a binocular microscope. Oxygen isotope analyses were conducted at the California Institute of Technology using previously described CO2-laser fluorination techniques (Sharp, 1990, Valley et al., 1995, Eiler et al., 2000a, Eiler et al., 2000b). Measurements of working standards UWG garnet (2-4 times per day) and SCO olivine (2-6 times per day) were interspersed with measurements of unknown samples. Results for these standards yielded δ18O values of 5.78 ± 0.09‰ (UWG, n = 22) and 5.29 ± 0.07‰ (SCO, n=27), which are within the accepted values of UWG (5.80‰) and SCO (5.25‰) (Valley et al., 1995, Eiler et al., 2000a, Eiler et al., 2000b). All standard deviations are 1σ and refer to long-term reproducibility of raw measurements. All δ18O values for unknown samples on a given day were corrected by the difference between the measured and accepted values for standards on that day. Duplicated measurements (2-3) were performed on each unknown sample with 1σ standard deviations typically less than 0.1‰ (standard deviations for each sample reported in Table 1). RESULTS Oxygen isotope data from olivine grains are presented in Table 1. The δ18O values for the total data set (4.74-5.40‰) fall over a greater range than previously reported for phenocryst δ18O values in the Galapagos (4.64-5.04‰; range reported in Geist et al. (1998) for plagioclase cumulates from several locations in the Galapagos converted to the range in equilibrium olivine values using a Δplagioclase-olivine = 0.7‰ (Eiler, 2001 and references therein)). This range also exceeds the range of pristine, “normal” mantle olivine (5.2 ± 0.2‰; Mattey et al., 1994). ! 261! In general, olivine from Fernandina, Pinta and Floreana (which represent the non- MORB mantle end-members in the Galapagos) show a narrow range of δ18O values (4.97-5.11‰; Fig. 2). Olivine from Cerro Azul and Santiago also fall within to lower than this range. The highest δ18O values can be found in the 3 olivine samples from San Cristobal (maximum δ18O = 5.40‰). The lowest δ18O values can be found in the one sample from Ecuador Volcano and one of the samples from Sierra Negra (4.74‰), (White et al., 1993, Reynolds & Geist, 1995, Blichert-Toft & White, 2001, Harpp & White, 2001, Saal et al., 2007). However, Sierra Negra also displays one of the largest ranges in δ18O values for a given island in this study (4.74-4.95‰), along with Darwin Volcano (4.78-5.04‰), Cerro Azul Volcano (4.86-5.13) and Santa Cruz (4.75-5.22‰). In the following discussion we will examine δ18O values within individual islands as well as across the archipelago as a region. DISCUSSION Fractional Crystallization and Partial Melting Magmatic differentiation tends to cause only small changes in δ18O values because common crystallizing phases have both positive, in the case of plagioclase, and negative, in the case of olivine and clinopyroxene, isotopic mineral-magma fractionation factors (e.g. Chiba et al., 1989, Harris et al., 2000, Eiler, 2001, Cooper et al., 2004). To test the expected effects of fractional crystallization on δ18O values, the liquid line of descent (LLD) is calculated for both a trace element depleted and enriched composition using PetroLog3 (Fig. 3; Danyushevsky & Plechov, 2011). The isobaric LLD is calculated at 2 kbar in a system closed to oxygen. For the modeling, we use the trace element depleted ! 262! composition of a high MgO olivine-hosted melt inclusion from Santiago Island and the trace element enriched composition of a high MgO inclusion from Fernandina (reported in table 2 and Koleszar et al. (2009)). While these compositions do not represent the entire range in major element data found in the Galapagos, these compositions include measurements for H2O contents, which can have a great effect on crystallization trends, and generally bracket the major element data thus providing end-member cases for the expected δ18O fractionation (Fig. 3). The pressure is based on an average pressure of entrapment of melt inclusions collected from Fernandina and Santiago, though changing the pressure between 0.5 and 4 kbars does not change the interpretation. The variation in δ18O for Rayleigh distillation is calculated using the fractionation factors presented in table 2 for crystallizing minerals, recalculating the bulk fractionation factor (α) at each step based on the modal abundances of crystallizing phases. These values are assumed to be constant over the course of fractional crystallization to simplify the calculation. For tholeiitic basalts and a temperature change < 200°C at magmatic temperatures this has found to be a valid simplification (Chiba et al., 1989, Harris et al., 2000, Eiler, 2001, Wang & Eiler, 2008). The primary δ18O values of the crystallizing melt are fixed so that resulting trends pass through the measured value for Santiago and Fernandina, though changing this value does not effect the total change in δ18O over the course of the LLD. The results of the calculation show that the δ18O of olivine in equilibrium with the crystallizing melt can only be expected to change by a maximum of 0.15‰, which inadequately addresses the range we measure in δ18O values. Variations in the degree of partial melting are also not predicted to appreciably increase the δ18O values of a melt by more than ~0.1‰ (Eiler, 2001). Large increases in ! 263! δ18O values due to partial melting are the result of an increasing albitic component in the melt and only become a concern at very small degrees of melt. For this study, the δ18O values within a given island do not vary with La/Yb or Sm/Yb values (trace element ratios sensitive to degree and depth of melting, Fig. 4). This includes volcanoes that show large variations in δ18O, such as Santa Cruz, with very little variation in Sm/Yb or La/Yb, and volcanoes that show large variations in Sm/Yb and La/Yb, but change very little in δ18O, such as is the case for Floreana. Lavas from Wolf Volcano tend to have values in Sm/Yb and La/Yb equivalent to lavas from trace element enriched components such as Fernandina, but are isotopically more similar to Genovesa. These characteristics are attributed to low degree melts of a depleted mantle source (Geist et al., 2005). The δ18O values from the Wolf Volcano olivines are slightly lower than normal upper mantle values (4.89 ± 0.04‰), opposite to what would be predicted if a low degree of melting were greatly affecting δ18O values. Samples from Santa Cruz (which show the largest range in δ18O values) also show this trend, with the two samples with the slightly higher La/Yb ratios (and presumably lower degree of melting) also having lower δ18O values. Fractional crystallization and partial melting, as a result, are not the major factors controlling δ18O variation in the Galapagos Archipelago. δ18O variation of the main isotopic end-members in the Galapagos Samples from the main isotopic mantle end-members in the Galapagos (Fernandina, Floreana, and Pinta) cover a large range in radiogenic isotopes, but only show a narrow range in δ18O values (4.97-5.11‰), within error of the range found for upper mantle olivine (Fig. 2; Mattey et al., 1994, Eiler, 2001). This range is at the lower end of the ! 264! upper mantle range, similar to high 3He/4He olivine from Baffin Island (Kent et al., 2004). This suggests that the Galapagos mantle has a homogenous value for δ18O. This value is at the lower end of the range predicted for normal upper mantle olivine, and is consistent with previous studies of oxygen isotopes in the Galapagos (Geist et al., 1995, Geist et al., 1998). Assimilation δ18O values greater than normal upper mantle Oxygen isotope fractionation between minerals and water is temperature dependent. As a result, low temperature (< 200°C) hydrothermal alteration leads to 18O enrichment while high temperature alteration (> 250°C) leads to 18O depletion of crust relative to unaltered basalts (e.g. Muehlenbachs & Clayton, 1976, Muehlenbachs, 1998, Alt & Bach, 2006). Recrystallization of cumulates can also lead to oxygen exchange with seawater and has been shown in the case of plagioclase cumulates to increase δ18O values up to 7.4‰ (Alt & Bach, 2006, Gao et al., 2006). Interaction with different parts of the crust, as a result, can potentially lead to different variations in δ18O values (Garcia et al., 1998, Wang et al., 2003, Gaffney et al., 2005, Wang & Eiler, 2008, Genske et al., 2013). The island of San Cristobal, which has the highest δ18O values measured here, is identified by Saal et al. (2007) as having interacted with plagioclase-rich cumulates during melt transport through the oceanic lithosphere. This is indicated by positive Sr, Ba and Eu anomalies on a primitive mantle normalized diagram (elements that are difficult to fractionate from other trace elements with similar incompatibility without the presence of plagioclase). Using Sr/Sr* ≥ 1 (SrPM/[(CePM +NdPM )/2]), and Ba/Nb ratios as indicators ! 265! of assimilation of plagioclase cumulates, the data of this study show a positive correlation between increasing δ18O values and increasing amounts of assimilation (Fig. 5). This includes the islands with large δ18O variability, such as Sierra Negra and Santa Cruz. There is also a weak correlation between increasing δ18O values and Δ7/4 ratios (measure of the deviation of 207Pb/204Pb from the Northern Hemisphere Reference Line and defined in Hart (1984); Fig. 2c, 2f). The Δ7/4 ratios, however, are in range of the values of Genovesa, which is MORB-like, rather than Pinta, which is characterized by high Δ7/4. Sr isotopes show this as well, with San Cristobal falling at low values of 87Sr/86Sr, closer to the Genovesa end of the array. If the high δ18O values were due to ancient recycled material in the mantle source, a stronger correlation of δ18O values with more radiogenic isotopes would be expected. Samples from Floreana, which also show high values of Sr/Sr* and Ba/Nb, do not show correspondingly high δ18O values. Therefore the higher δ18O values of San Cristobal and the volcanoes that display large ranges in δ18O values are more consistent with the assimilation of plagioclase cumulates with the isotopic composition of Genovesa in the present-day lithosphere. δ18O values lower than normal upper mantle Samples with the lowest δ18O values come from Sierra Negra, Ecuador, Alcedo, Darwin, Wolf and Santa Cruz volcanoes which produce lavas with intermediate isotopic compositions to Fernandina, Floreana, and Pinta (Fig. 2). δ18O values that are outside the accepted upper mantle range that are explained by the presence of recycled hydrothermally altered material in the mantle source are often accompanied by very radiogenic isotopic signatures (e.g. Eiler, 2001, Skovgaard et al., 2001, Cooper et al., ! 266! 2004, Workman et al., 2008, Day et al., 2009, Day et al., 2010). Therefore the lowest δ18O values in the Galapagos would be predicted to reside in the Floreana and/or Pinta mantle sources that are thought to contain recycled material and have some of the most extreme radiogenic isotopes in the Galapagos. However, what is observed is that the lowest δ18O values are found in lavas with intermediate isotopic compositions. Therefore, the low δ18O values are either being produced by a low δ18O component with intermediate isotopes, in contrast to what is observed elsewhere, or are produced by interaction with a low δ18O component in the shallow lithosphere. The lighter δ18O olivine from Sierra Negra, Ecuador, Alcedo, Darwin, Wolf and Santa Cruz volcanoes come from lavas that are more evolved and have experienced fractional crystallization (Fig. 3). These lavas range in terms of enrichment with Darwin and Wolf Volcanoes anchoring the more depleted, MORB-like end of the variability. Assimilation will often occur concurrently with fractional crystallization (e.g. Depaolo, 1981). The lavas from these volcanoes, however, do not show unusual trace element or radiogenic isotope characteristics, making it difficult to unambiguously prove that assimilation is occurring in these samples. Meteoric water, seawater and gabbro that has interacted with seawater at high temperatures will have δ18O values lower than upper mantle olivine and can potentially be sources of low δ18O in the Galapagos (Gao et al., 2006, Thirlwall et al., 2006), but this interaction must not create a significant change in the trace element or radiogenic isotope characteristics of the melt. Meteoric water has been shown to impart a δ18O value distinct from that of seawater in areas where precipitation leads to an active hydrothermal system. This is illustrated in Hawaii where there is a correlation between areas of greater precipitation ! 267! and lower δ18O values in olivine (Wang & Eiler, 2008). In the Galapagos, however, there does not appear to be a correlation with precipitation patterns. In Santa Cruz, for example, two of the three samples were collected in the dry lowlands (area of low precipitation) while one sample was collected in the humid highlands (area of greatest precipitation) (Kurz & Geist, 1999, Saal et al., 2007, Trueman & d’Ozouville, 2010). The sample collected in the humid highlands shows the highest value for δ18O in contrast to what would be predicted if the source of low δ18O in the Galapagos were the result of the meteoric hydrothermal system. Seawater and gabbro that is hydrothermally altered at high temperatures can impart a low δ18O value to melt (Kent et al., 2004, Gao et al., 2006, Thirlwall et al., 2006, Bindeman et al., 2008). To test the effects of each of these assimilants on Galapagos melt, we model a mixture of the composition of GI92-19 (Sierra Azul) with seawater and 1 hydrothermally altered gabbro composition with 2 different δ18O values. GI92-19 is set to have an initial δ18O value equal to the melt composition in equilibrium with the average olivine values measured for Fernandina, Floreana and Pinta (5.52‰). For seawater the δ18O = 0‰ while the gabbro composition (composition Bp; Hart et al., 1999), uses both a minimum δ18O value (+2.5‰) found for hydrothermally altered gabbro and the global average δ18O value (+4.0‰) (compositions of assimilants presented in Table 3). We chose a gabbro composition without a strong plagioclase cumulate trace element signature so as to produce the smallest change in the composition of the melt. The gabbro composition has an assigned 87Sr/86Sr ratio equivalent to Genovesa to model a gabbro created at the GSC. To reproduce the δ18O value of GI92-19 requires 18% mixing with the gabbro with the global average δ18O, 10% for the gabbro with the minimum value for ! 268! δ18O, and only 2.5% mixing with seawater (Fig. 6). This is close to the amount of seawater that would be needed to produce the full range in Cl/K ratios of the submarine glass with the composition of Sierra Negra/Cerro Azul reported in Peterson et al. (submitted to Journal of Petrology). Mixing with each of these three components does not greatly affect the radiogenic isotope or trace element composition of the melt, which is consistent with the samples not having unusual compositions outside of their oxygen isotope signatures. Therefore interaction with seawater and hydrothermally altered gabbro could potentially be the source of low δ18O values in the Galapagos Archipelago. Alternative explanation for light δ18O in the Galapagos The presence of pyroxenite in the mantle source of a given hotspot has been invoked to explain variations in δ18O values both above and below the upper mantle range (e.g. Eiler, 2001, Skovgaard et al., 2001, Cooper et al., 2004, Thirlwall et al., 2006, Day et al., 2009, Day et al., 2010). The low δ18O values of these studies, however, are correlated with highly radiogenic isotopic signatures. In the Galapagos, the low δ18O values are associated with lavas of intermediate isotopic compositions to Fernandina, Floreana and Pinta (that have δ18O values within the accepted range of pristine upper mantle). This suggests, as we have argued, that these values are indicative of interaction with the present-day shallow level crust as opposed to being a characteristic of the mantle source. The study of Vidito et al. (2013), however, using the minor element composition of olivine collected across the archipelago argues that source lithology and the presence of pyroxenite in the Galapagos is not evidenced by the radiogenic isotope signatures of lavas. Pyroxenite in the source of OIBs is thought to be the product of eclogite melt ! 269! reacting with peridotite in the mantle to form a pyroxene-rich source. Olivine formed from partial melts of this reacted source record the hybridization by retaining high Ni and Si contents, but low Mn and Ca contents. Because clinopyroxene has a negative mineral- magma fractionation factor (Chiba et al., 1989, Eiler, 2001) it is conceivable that a pyroxene-rich source would produce a light δ18O melt. Vidito et al. (2013) used measurements of Ni, Si, Mn and Ca in olivine collected from across the Galapagos Archipelago to create a map of proposed areas with high concentrations of pyroxenite in the mantle. They identified Sierra Negra, Ecuador, Darwin, Alcedo, Santiago and Roca Redonda as having > 40% pyroxenite in the source. With the exception of Santiago, these cover almost all of the locations that exhibit low δ18O values in this study. Though the study of Vidito et al. (2013) only reports on one sample from Santiago (which produces lavas with a large range in compositions), it has a similar composition to the sample from Santiago analyzed here. Santiago, however has δ18O values within range of normal upper mantle, in contrast to the hypothesis that pyroxenite in the mantle source is the cause low δ18O. Wolf volcano, which also has olivine with low δ18O values, is identified as a purely peridotite source and would therefore require a different source of low δ18O. If the low δ18O values of the samples of this study are the result of the presence of pyroxenite in the source that has an intermediate isotopic composition, the olivines with low δ18O should also show enrichments in Ni and Si but depletions in Mn and Ca. These measurements will therefore be the focus of future measurements. CONCLUSIONS ! 270! New oxygen isotope data measured in olivine grains taken from 12 different volcanoes/islands in the Galapagos Archipelago show a larger range in δ18O values than previously measured. The main non-MORB isotopic mantle end-members in the Galapagos (Fernandina, Floreana, and Pinta) show δ18O values (5.02 ± 0.08‰,) within range of olivine from pristine upper mantle. Generally, δ18O values do not correlate with radiogenic isotopes, which indicates that variation above and below normal mantle values is more likely due to interaction with shallow, hydrothermally altered material than the presence of low δ18O components in the mantle source. This is corroborated for δ18O values above the accepted mantle limit by the trend of increasing Sr/Sr* and Ba/Nb ratios, which are indicators of the assimilation of plagioclase cumulates, with increasing δ18O values. Therefore the high δ18O values of San Cristobal (maximum = 5.40‰) and the large range in δ18O values of Sierra Negra Santa Cruz and Darwin Volcano are the result of the assimilation of plagioclase cumulates. The δ18O values of Sierra Negra, Alcedo, Darwin, Wolf and Ecuador volcanoes that fall below the mantle limit (4.85 ± 0.1‰), do not exhibit geochemical characteristics indicative of assimilation which makes it difficult to unambiguously identify this process as the cause of the low δ18O values in the Galapagos. However, a small amount of assimilation of seawater (δ18O = 0‰) and/or hydrothermally altered gabbro without a cumulate component (δ18O = +2.5/4.0‰) can produce the low values for δ18O without a large change to the trace element or radiogenic isotope composition of the melt and assimilation of hydrothermally altered material and/or seawater is consistent with volatile/refractory element variations previously measured in the Galapagos. These results show that interaction with the hydrothermally ! 271! altered crust is a widespread phenomenon in the Galapagos Archipelago and should be considered when inferring mantle geochemistry of lavas erupted at the surface. ACKNOWLEDGEMENTS This work was supported by the National Science Foundation Graduate Research Fellowship (Grant No. DGE-1058262 to M.E.P.), the National Science Foundation Division of Ocean Sciences (Grant No. 0962195). We would like to thank T. Prissel, B. Parks and E. Chin for their thoughtful discussions during the preparation of this manuscript. ! 272! References Alt, J. C. & Bach, W. (2006). Oxygen isotope composition of a section of lower oceanic crust, ODP Hole 735B. Geochemistry Geophysics Geosystems 7, Artn Q12008 Doi 10.1029/2006gc001385. Bindeman, I. N., Fu, B., Kita, N. T. & Valley, J. W. (2008). Origin and evolution of silicic magmatism at Yellowstone based on ion microprobe analysis of isotopically zoned zircons. Journal of Petrology 49, 163-193, Doi 10.1093/Petrology/Egm075. Bindeman, I. N., Sigmarsson, O. & Eiler, J. (2006). 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J. & Taylor, R. N. (2006). Low delta O-18 in the Icelandic mantle and its origins: Evidence from Reyjanes Ridge and Icelandic lavas. Geochimica Et Cosmochimica Acta 70, 993- 1019, Doi 10.1016/J.Gca.2005.09.008. Trueman, M. & d’Ozouville, N. (2010). Characterizing the Galapagos terrestrial climate in the face of global climate change. Galapagos Research 67, 26-37. Valley, J. W., Kitchen, N., Kohn, M. J., Niendorf, C. R. & Spicuzza, M. J. (1995). UWG- 2, a garnet standard for oxygen isotope ratios: Strategies for high precision and ! 280! accuracy with laser heating. Geochimica Et Cosmochimica Acta 59, 5223-5231, Doi 10.1016/0016-7037(95)00386-X. Vidito, C., Herzberg, C., Gazel, E., Geist, D. & Harpp, K. (2013). Lithological structure of the Galapagos Plume. Geochemistry Geophysics Geosystems 14, 4214-4240, Doi 10.1002/Ggge.20270. Wang, Z. G., Kitchen, N. E. & Eiler, J. M. (2003). Oxygen isotope geochemistry of the second HSDP core. Geochemistry Geophysics Geosystems 4, Artn 8712 Doi 10.1029/2002gc000406. Wang, Z. R. & Eiler, J. A. (2008). Insights into the origin of low-delta O-18 basaltic magmas in Hawaii revealed from in situ measurements of oxygen isotope compositions of olivines. Earth and Planetary Science Letters 269, 376-386, Doi 10.1016/J.Epsl.2008.02.018. White, W. M., Mcbirney, A. R. & Duncan, R. A. (1993). Petrology and Geochemistry of the Galapagos-Islands - Portrait of a Pathological Mantle Plume. Journal of Geophysical Research-Solid Earth 98, 19533-19563. Workman, R. K., Eiler, J. M., Hart, S. R. & Jackson, M. G. (2008). Oxygen isotopes in Samoan lavas: Confirmation of continent recycling. Geology 36, 551-554, Doi 10.1130/G24558a.1. ! 281! FIGURE CAPTIONS Figure 1. Map of the Galapagos archipelago with prominent features labeled. Scale bar and bathymetry color scale are included the left side on the map. Volcanoes where olivine grains were collected are indicated by symbols used for geochemical plotting. Map created using GeoMapApp http://www.geomapapp.org (Ryan et al., 2009). Figure 2. 87Sr/86Sr, 3He/4He, and Δ7/4 versus (A-C) 143Nd/144Nd and (D-F) δ18O values of olivine phenocrysts. Isotope ratios for lava associated with the δ18O of olivine phenocrysts were reported Saal et al., (2007) and Kurz and Geist (1999). Grey field in panels D-F show the range of typical upper mantle olivine δ18O (Mattey et al., 1994, Eiler, 2001). Notice that samples from Floreana, Fernandina, and Pinta (representing the extreme enriched isotopic end-members in the Galapagos) are within the upper mantle field. Errors shown for δ18O values are 1σ. Figure 3. MgO wt% versus A.) TiO2 wt %, and B.) δ18O values for individual samples. Major element data for lava associated with the δ18O of olivine phenocrysts were reported in (Saal et al., 2007)). Superimposed on the data are predicted liquid lines of descent (LLD) for olivine (olv) - olv + plagioclase (plag) – olv + plag + clinopyroxene (cpx) (solid black and purple lines, where black circle indicates the appearance of a new fractionating crystal phase). Calculations were performed in PetroLog3 (Danyushevsky & Plechov, 2011) isobarically at 2 Kbars of pressure using 2 different starting compositions, which can be found in Table 2. Black LLD is calculated from an enriched Fernandina ! 282! composition, while purple LLD is calculated from a depleted Santiago composition. LLDs for δ18O values was calculated using mineral/melt fractionation factors presented in Table 2 and described in the text. Notice in panel B that at a given MgO content there is a large within island variation in δ18O values. Errors shown for δ18O values are 1σ. Figure 4. δ18O values versus A.) Sm/Yb and B.) La/Yb for individual olivines. Trace element data for lava associated with the δ18O of olivine phenocrysts were reported in Saal et al. (2007) Notice there is very little within island variation in trace element ratios with δ18O values. Errors shown for δ18O values are 1σ. Figure 5. δ18O values versus A.) Sr/Sr* (SrPM / [(CePM + NdPM) / 2]) and B.) Ba/Nb ratios for individual samples. Also shown is the composition of plagioclase-rich gabbro Fp and Hp reported in Hart et al. (1999) with the maximum value for δ18O measured for hydrothermally altered plagioclase cumulates (Alt & Bach, 2006, Gao et al., 2006). Samples Fp and Hp are unaltered, lower crustal protolith gabbros from the 500 m gabbroic section drilled at site 735 during Leg 118 (Ocean Drilling Project) in the SE Indian Ridge. Inset at left shows a zoomed in version of panels A and B. Notice the trend of increasing Sr/Sr* and Ba/Nb ratios towards the values of the gabbro with increasing δ18O. This suggests δ18O values are being increased by the assimilation of plagioclase cumulates. Errors shown for δ18O values are 1σ. Figure 6. δ18O values versus A.) 87Sr/86Sr and B.) Ba/Nb ratios for individual samples. Also included are simple mixing lines between lava composition of GI92-8 (δ18O value ! 283! set to melt in equilibrium with average olivine value from Fernandina, Floreana, and Pinta) and both seawater and an unaltered protolith lower crustal gabbro composition that does not show trace element signature of plagioclase cumulate (Bp; Hart et al., 1999) with 2 different values for δ18O (2.5 and 4.0 ‰ explained in text). Each point along mixing lines represent 1% mixing. Notice that the trends are hyperbolic, which allows the δ18O value of the melt to change without greatly affecting the isotopic or trace element composition. Inset shows zoomed out view of panels A and B. Errors shown for δ18O values are 1σ. ! 284! FIGURES Figure 4-1 4°N Galapagos Spreading Center (GSC) Fernandina Meters Floreana 2000 Pinta 0 Sierra Negra 3°N -2000 Cerro Azul -4000 Alcedo Ecuador -6000 Darwin Vn Wolf Vn 2°N Santiago Wolf-Darwin San Cristobal lineament Marchena Santa Cruz 1°N Pinta Genovesa NE Seamounts 0°N Fernandina Santiago San Cristobal Isabela Santa Cruz 1°S Floreana 2°S 94°W 93°W 92°W 91°W 90°W 89°W 88°W 87°W 86°W 85°W ! 285! Figure 4-2 A. D. 0.7036 0.7034 87Sr/86Sr 0.7032 0.7030 0.7028 0.7026 B. E. 30 Mantle value 28 26 24 22 3He/4He 20 18 16 14 12 10 8 C. F. 5 4 3 2 1 Δ7/4 0 -1 -2 0.51290 0.51300 0.51310 4.8 4.9 5.0 5.1 5.2 5.3 5.4 0.51295 0.51305 143Nd/144Nd Fernandina δ18Oolivine (‰) Floreana Pinta Sierra Negra Cerro Azul Alcedo Ecuador Darwin Vn Wolf Vn Santiago San Cristobal Santa Cruz Genovesa ! 286! Figure 4-3 A. olv olv+plag olv+plag+cpx 5 4 TiO2 (wt%) 3 2 Fernandina Floreana 1 Pinta Sierra Negra Cerro Azul Alcedo B. Ecuador 5.4 Darwin Vn Wolf Vn 5.3 Santiago San Cristobal δ18Oolivine (‰) 5.2 Santa Cruz 5.1 5.0 4.9 4.8 4.7 5 6 7 8 9 10 11 12 13 MgO (wt%) ! 287! Figure 4-4 A. B. 5.4 5.3 δ18Oolivine (‰) 5.2 5.1 5.0 4.9 4.8 4.7 1.0 1.5 2.0 2.5 1 2 3 4 5 6 7 8 9 Sm/Yb La/Yb Fernandina Floreana Pinta Sierra Negra Cerro Azul Alcedo Ecuador Darwin Vn Wolf Vn Santiago San Cristobal Santa Cruz ! 288! Figure 4-5 10 A. 1.8 1.6 Sr/Sr* 1 1.4 Sr/Sr* 1.2 1.0 0.8 0.1 B. 18 18 16 16 14 14 12 Ba/Nb 12 Ba/Nb 10 10 8 8 6 6 4 4 4.7 4.8 4.9 5.0 5.1 5.2 5.3 5.4 δ18Oolivine (‰) 5.0 5.4 5.8 6.2 6.6 7.0 7.4 Fernandina δ18Oolivine (‰) Floreana Pinta Sierra Negra Cerro Azul Alcedo Ecuador Darwin Vn Wolf Vn Santiago San Cristobal Santa Cruz Plagioclase cumulate ! 289! Figure 4-6 A. A. 0.7034 0.709 0.708 0.7032 0.707 87Sr/86Sr 87Sr/86Sr 0.706 0.7030 0.705 0.704 0.7028 0.703 B. B. 15 15 Ba/Nb Ba/Nb 10 10 5 5 0 1 2 3 4 5 4.60 4.80 5.00 5.20 5.40 Fernandina 18Oolivine (‰) 18Oolivine (‰) Floreana Pinta Sierra Negra Cerro Azul Alcedo Ecuador Darwin Vn Wolf Vn Santiago San Cristobal Santa Cruz Gabbro (2.5 ‰) Gabbro (4.0 ‰) Seawater 290! TABLES Table 1. Results of in situ analysis for oxygen isotope compositions of individual olivine grains from subaerial lavas collected across the Galapagos Archipelago Sample ID Island δ18O (‰) 1σ (‰) GI92-16 Alcedo 4.86 0.05 CA90-19-1 Cerro Azul 4.86 0.20 SG 93-12 Cerro Azul 5.05 0.06 SG 93-5 Cerro Azul 5.13 0.09 GI92-17 Darwin Vn 4.78 0.01 GI92-18 Darwin Vn 5.04 0.17 GI92-19 Ecuador 4.74 0.04 NSK97-215 Fernandina 5.00 0.05 SG93-19 Floreana 5.00 0.10 SG93-22 Floreana 4.97 0.02 SG93-23 Floreana 5.11 0.12 SKV98-108 Pinta 5.01 0.02 NSK97-126 San Cristobal 5.40 0.12 NSK97-252 San Cristobal 5.32 NSK97-254 San Cristobal 5.30 0.10 GI92-04 Santa Cruz 5.22 0.18 GI92-1 Santa Cruz 4.75 0.05 GI92-2A Santa Cruz 4.84 0.09 NSK97-238 Santiago 5.05 0.05 GI92-5 Sierra Negra 4.95 0.09 GI92-8 Sierra Negra 4.74 W95-74 Wolf Vn 4.92 0.01 W95-75-R Wolf Vn 4.86 0.13 Refer to Saal et al., (2007) for detailed description of geochemistry ! 291! Table 2. Fractional crystallization model Starting Composition Enriched Depleted Compositiona Compositionb Major Elements (wt%) SiO2 47.11 46.46 TiO2 2.23 1.34 Al2O3 15.97 14.36 FeO* 9.21 11.77 MnO 0.12 0.2 MgO 10.34 11.86 CaO 11.29 11.5 Na2O 2.19 2.14 K 2O 0.32 0.08 P 2O 5 0.25 0.07 H2O (wt%) 0.68 0.12 Pressurec 2kbars δ18Odo-melt 5.5 5.62 Equilibrium fractionation factorse Δ α plagioclase 0.2 1.0002000 clinopyroxene -0.2 0.9998000 olivine -0.5 0.9995001 a Composition of melt inclusion AHA D25C-3-43 reported in Koleszar et al., (2009), but modified to an MgO = 10.3 wt% (value of Fernandina lava with olivine grain measured for oxygen isotopes). Modification was done by adding equilibrium olivine back into the composition in 0.5% increments. Correction required 2.5% olivine addition. b Composition of melt inclusion STG06-29-06 reported in Koleszar et al., (2009) c Models used in PETROLOG: Ol, CPX, PLAG- Danyushevskey, 2001, Melt density- Lange and Carmichael, 1987, Melt viscosity- Bottinga & Weill 1972. Oxygen fugacity set at QFM and then run closed for oxygen d Initial δ18O values of the melt set to pass through the δ18O values measured for Santiago and Fernandina compositions e Equilirium fractionation factors based on the results of Eiler (2001) and Chiba et al., (1989). Δ = 1000*lnα ! 292! Table 3. Simple mixing model compositions Gabbro GI92-8* Composition (Bp)a Seawaterb Major Elements (wt%) SiO2 49.06 41.20 TiO2 3.36 6.12 Al2O3 14.18 9.61 FeO* 12.71 24.69 MnO 0.21 0.33 MgO 5.99 6.46 CaO 10.62 8.85 Na2O 3.03 2.63 K 2O 0.49 0.06 P 2O 5 0.344 0.05 H2O (wt%) 97.5 Ba (ppm) 115 5 0.00117 Nb (ppm) 26.82 2.05 0 Sr (ppm) 302 122 7.8 87 Sr/86Sr 0.7035 0.70253 0.7091 δ18Oo 5.52 2.5/4 0 * δ18Oo of GI92-8 set to the melt value in equilibrium with the average δ18O value of olivine from Fernandina, Floreana and Pinta a Gabbro composition is a prisitine unaltered lower crustal gabbro composition without a strong cumulate trace element signature from Hart et al. (1999). Sr isotope is set to be equal to Genovesa which is the MORB-like component in the Galapagos to simulate a gabbro created at the GSC. Both a δ18O = +2.5 and +4.0‰ are tested to see relative amounts of assimilant needed. b Seawater trace element composition taken from Quinby-Hunt (1983). Isotopic composition taken from (Farrell et al., 1995). ! 293! SYNTHESIS The geochemistry of ocean islands basalts begins with a mantle source derived from complex materials with unknown quantities of recycled components from as deep as the core-mantle boundary (CMB). Separating the signatures of the deep mantle source from the effects of shallow level processes is difficult without high quality data gathered to specifically address the amount of communication occurring between mantle and surface reservoirs. In this dissertation, I have used a large geochemical data set of submarine glass, melt inclusions and whole rock, collected from across the Galapagos Archipelago, to explore the influence of shallow level processes on lavas erupted at the surface and to isolate deep mantle signatures unique to the Galapagos plume. Chapter 1 uses Pb isotopes measured on melt inclusions from Fernandina and Santiago to explore the origin of the “ghost plagioclase” signature. While commonly explained by invoking an ancient recycled gabbro (now eclogite) in the mantle source, the Pb isotope variations of the Santiago and Fernandina inclusions are more easily explained by melt interaction with plagioclase cumulates in the plumbing system of the islands. Chapter 2 sets out to characterize the volatile contents of the different isotopic end-members lavas using new major, trace, and volatile content (H2O, CO2, F, Cl, S) radiogenic and He isotope measurements of 118 submarine glasses collected across the Galapagos Archipelago. The samples geographically and geochemically define 2 arrays that span the signatures of Fernandina and Sierra Negra/Cerro Azul (array 1) and Genovesa and Pinta, Wolf, and Darwin Islands (array 2). Due to the effects of degassing, the samples are found to have significantly lowered CO2 concentrations, with minimal ! 294! lowering of H2O and S, but no observable affect on F or Cl. The samples range from just saturated with a sulfide liquid to sulfide undersaturated. Assimilation of hydrothermally altered material, traced using Cl/K ratios and the magnitude of the Sr anomaly on a primitive mantle normalized diagram (Sr/Sr*), is found to have affected most of the glasses of the study, but only the isotopically depleted glasses have significantly contaminated signatures. Despite the affects of shallow level processes there is still a measurable variation in ratios of volatile and refractory elements with similar incompatibilities between the isotopically unique groups of glasses. The most distinctive differences are found in H2O/Ce ratios and F/Nd ratios. The high 3He/4He Fernandina end-member in the Galapagos is found to have higher water contents than the MORB end-member of array 2. This results in a 5-7x lower viscosity of the Fernandina mantle in comparison to depleted upper mantle which has implications for mantle flow and plume- ridge interaction. The high 3He/4He Fernandina end-member also shows an order of magnitude larger C/3He ratio than what has been measured for the depleted upper mantle and among the highest yet reported for OIB samples. This suggests that the Galapagos mantle component hosting the high 3He/4He ratio is more carbonated than in other settings. Chapter 3 reports on new Fe3+/ΣFe ratios (a proxy for ƒO2) on a subset of glasses measured for major, trace, isotopic and volatile contents in Chapter 2. Utilizing the results from Chapter 2 and the new measurements of Chapter 3, we are able to identify the influence shallow level processes have on Fe3+/ΣFe ratios. Generally, Fe3+/ΣFe ratios increase with magmatic differentiation, decrease with sulfur degassing and increase with assimilation of hydrothermally altered material. After accounting for shallow level ! 295! processes, there is still a measurable variation in Fe3+/ΣFe ratios between the different isotopic end member groups which translates into a magmatic ƒO2 that ranges from ΔQFM = +0.16 to +0.74, between the ranges identified for MORB and arc basalts. In general, the Sierra Negra/Cerro Azul component is more oxidized than the Pinta/Wolf, Darwin component. Both the Siera Negra/Cerro Azul and Pinta mantle components are hypothesized to contain recycled material. The difference in oxidation state between these two sources and the correlation between Fe3+/ΣFe ratios and Pb isotopes suggest a difference in mantle lithology and therefore a difference in the type or influence of the recycled material on ƒO2. The high 3He/4He end-member is found to be the most oxidized. The combination of the high C/3He ratios identified in Chapter 2 and the high Fe3+/ΣFe ratios indicates that the high 3He/4He mantle source is both more carbonated and oxidized than previously thought for a primitive mantle component. In Chapter 4 we present new oxygen isotope of olivine phenocrysts taken from lava collected across the Archipelago. The main isotopic end-members in the Galapagos, (Fernandina, Floreana, and Pinta) have δ18O values within the narrow range found for the upper mantle and, in general, δ18O values do not correlate with radiogenic isotopes. We take this as evidence that the Galapagos mantle has a homogenous δ18O within range of the upper mantle, consistent with previous findings in the Galapagos (Geist et al., 1995, Geist et al., 1998). The positive correlation of Sr/Sr* and Ba/Nb ratios with δ18O values suggest that the maximum oxygen isotope values are the result of plagioclase cumulates. The low δ18O values of Sierra Negra, Alcedo, Darwin, Wolf and Ecuador volcanoes (4.85 ± 0.1‰), do not exhibit geochemical characteristics indicative of assimilation making it difficult to unambiguously identify the occurrence of this process. However, a small ! 296! amount of assimilation of seawater (δ18O = 0‰) and/or hydrothermally altered gabbro without a cumulate component (δ18O = +2.5/4.0‰) can produce the low values for δ18O without a large change to the trace element or radiogenic isotope composition of the melt. These results show that interaction with the hydrothermally altered crust is a widespread phenomenon in the Galapagos Archipelago and should be considered when inferring mantle geochemistry of lavas erupted at the surface. However, the results of this thesis have also isolated mantle signatures that are distinct in the Galapagos from mid- ocean ridge basalts in terms of volatile elements and oxidation state. These results further the conclusion that ocean islands basalts come from a reservoir unique from the depleted upper mantle and have important implications for ability of the mantle to preserve long- lived heterogeneities. This is consistent with plume theory and implies that the mantle is neither a well-mixed homogenous source nor a well-buffered system. FUTURE DIRECTIONS The scientific findings of these chapters pose several immediate questions for future research in the Galapagos as well as having implications for broader areas of research. The following discussion is a summary of those ideas. The results of chapter 1 highlight the importance of presenting isotopic measurements when trying to identify ancient recycled material in the mantle. The ghost plagioclase signature is a well-documented phenomenon that, despite only occurring in a small percentage of samples, can be found in suites of samples from multiple OIBs (Gurenko & Chaussidon, 1995, Hofmann & Jochum, 1996, Yang et al., 1998, Chauvel & Hemond, 2000, Sobolev et al., 2000, Kent et al., 2002, Huang et al., 2005, Ren et al., ! 297! 2005, Maclennan, 2008, Sobolev et al., 2011) and MORB (Kamenetsky et al., 1998, Danyushevsky et al., 2003, Danyushevsky et al., 2004). Only a few of these studies present isotopic measurements in addition to major and trace elements, yet many claim there is recycled gabbro in the mantle source of the area being studied. Without isotopic measurements these conclusions are ambiguous, as shown in Chapter 1. Therefore it will be worthwhile to return to these areas where the ghost plagioclase signature has been recognized and measure these inclusions for Pb isotopes. The results of Chapters 2 and 3 present interesting characterizations of volatile contents and oxidation state of the high 3He/4He end-member, the trace element enriched end-member characterized by the lavas of Pinta (low 3He/4He ratios, high Th/La, Δ7/4 and Δ8/4), and the trace element depleted end-member anchored by the lavas of Genovesa, which represent 3 of the 4 major isotopic end-members in the Galapagos. The fourth end-member identified in the submarine glasses has the composition of Sierra Negra/Cerro Azul, which, while enriched in comparison to MORB, has a signature that is isotopically intermediate to Floreana, Pinta, and Fernandina (White et al., 1993, Blichert- Toft & White, 2001, Harpp & White, 2001, Geist et al., 2006, Saal et al., 2007, Geist et al., 2008, Kurz et al., 2009, Harpp et al., 2014). This signature is therefore either the product of mixing of melts from several end-members or a unique but isotopically intermediate mantle source. The Floreana end-member, as a result, remains uncharacterized for volatiles or oxidation state. Floreana is dynamically and geochemically unique in the Galapagos. The eruption style is the most explosive, lavas are much more alkaline and it has the most radiogenic Sr and Pb isotopes measured in the archipelago (McBirney et al., 1969, Bow & Geist, ! 298! 1992, White et al., 1993, Harpp & White, 2001, Lyons et al., 2007, Harpp, 2014). The explosive eruption style, in particular, has led to the hypothesis that the Floreana mantle component is volatile-rich in comparison to the other components of the Galapagos plume. If the composition of Sierra Negra/Cerro Azul is the result of mixing between melts from the Floreana, Pinta and Fernandina end-members, then the higher F/Nd ratio and lower H2O/Ce ratio than either the Pinta or Fernandina end-members may be the result of the influence of Floreana on the geochemical signature of the lavas. This is inconsistent, however, with the hypothesis that the Floreana mantle source is enriched in volatiles. A means to test this would be to measure a combination of submarine glass with the isotopic composition of Floreana and sub-aerially collected olivine-hosted melt inclusions from the Floreana Island for volatile contents. The former of these 2 measurements is planned for the summer of 2015. In addition to volatile measurements, measuring Fe3+/ΣFe ratios for samples with the isotopic composition of Floreana would also provide an interesting future direction for research. In the study of Cottrell and Kelley (2013), submarine glass from mid-ocean ridges that show enrichments in Pb isotopes were found to be more reduced in terms of Fe3+/ΣFe ratios than normal MORB samples (Fe3+/ΣFe = 0.12 in comparison to the average value of 0.16 found for MORB (Cottrell & Kelley, 2011, Le Voyer et al., 2015)). The hypothesis put forth by the authors was that carbon subducted into the mantle prior to 2.3 Ga would be reduced due to the presence of an anoxic atmosphere during this time period (Kasting et al., 1993), resulting in reduced carbon domains in the mantle. The samples with low Fe3+/ΣFe ratios, as a result, are purported to be sampling this ancient reduced domain. An obvious test of this hypothesis would be to measure isotopically ! 299! HIMU samples, like the Floreana component, for Fe3+/ΣFe ratios. The HIMU mantle source is hypothesized to contain material that was at the surface at 1.8-2.3 Ga (e.g. Hanyu et al., 2011, Cabral et al., 2013, Hanyu et al., 2013). Because the Floreana component is atypical in comparison to other HIMU plumes (Blichert-Toft & White, 2001, Harpp & White, 2001, Harpp et al., 2014), it would be essential to combine these measurements with measurements from a separate hotspot such as the Austral Island chain, to better characterize the ƒO2 of the HIMU mantle source. The results of Chapters 2 and 3 also make the Galapagos a unique location to measure highly siderophile elements (HSE: Os, Ir, Ru, Pt, Pd, Rh, Re, and Au). Residual sulfide in the source, crystallization of sulfide from the melt and degassing of sulfur exerts a strong control on the variation of HSE in a melt (Barnes et al., 1985, Bockrath et al., 2004, Mungall & Su, 2005, Ballhaus et al., 2006, Fonseca et al., 2007, Brenan, 2008, Dale et al., 2009, van Acken et al., 2010). The presence of sulfide in the source and melt, as well as the partitioning of HSE into sulfide, is dependent on sulfur content and fO2 of the source (Mavrogenes & O'Neill, 1999, Holzheid & Grove, 2002, Liu et al., 2007, Righter et al., 2008). The subset of samples from the Galapagos that have been measured for both volatile concentrations and Fe+3/ΣFe ratios are thus uniquely constrained for S content and ƒO2 which are the controlling parameters for HSE fractionation. Variations can be correlated with isotopic values making the Galapagos an ideal location to explore the HSE budget of the mantle. Measuring Re-Os and Pt-Os isotopes on samples characterized for volatile elements and Fe3+/ΣFe ratios is also an interesting direction for research in the Galapagos. For the Re-Os isotope system (187Re – 187Os + β-; t1/2 = 4.16 x 1010) and the Pt-Os isotope ! 300! system (190Pt – 186 Os + α; t1/2 = 4.5 x 1011), Re, Pt and Os are all HSE, but Re and Pt are both moderately compatible elements, while Os is a highly incompatible element (Hauri & Hart, 1997, Shirey & Walker, 1998, Fonseca et al., 2007, Mallmann & O'Neill, 2007). To generate coupled, high time-integrated Re/Os and Pt/Os ratios and therefore high 187 Os/188Os and 186 Os/188Os ratios in basaltic glass and whole rock we measure at the surface requires specific conditions for fractionation. Combining variations in Os isotopes with trace elements, other radiogenic isotopes, volatile elements and Fe3+/ΣFe ratios, as a result, allows us to examine the mantle for components such as recycled, crustal materials, contributions from the core, or a cumulate pile at the core-mantle boundary that has trapped liquid metal from the outer core (Ireland et al., 2011). For 187 example, Floreana, could be expected to show high Os/188Os ratios (up to 0.17) predicted for ancient recycled oceanic crust and lithosphere and in the range of other HIMU OIBs (Day, 2013). However, if the atypical Sr and Hf isotopes are due to the 187 presence of recycled sediment (which would contain low Os/188Os isotopes in relation to the isotopes expected of recycled lithosphere due to the moderate compatibility of Re) 187 lower Os/188Os ratios (between 0.14-0.17) could be expected. Fernandina, in 186 comparison, may show more radiogenic Os/188Os ratios in comparison to chondritic values (0.1198373 ± 49) and the other regional mantle end-members if the component housing the high 3He/4He component has trapped liquid from the outer core which would have the high concentration of Pt needed to produce high 186Os/188Os ratios (Ireland et al., 187 2011). Therefore the combination of Os/188Os and 186 Os/188Os ratios with measurement of relative and absolute abundances of HSE can elucidate characteristic differences between primitive mantle sources in terms of HSE and therefore inform on the evolution ! 301! of HSE elements in the mantle. Os isotopes are also highly sensitive tracers of interaction with the crust, which would continue to build on the results of this thesis. The results of chapter 4 provide an interesting snapshot of oxygen isotope variation in the Galapagos, with compelling results that require further investigation. The origin of low δ18O values, in particular, could be better examined with a more focused study. In addition to measuring the olivine grains for their major and minor element compositions, focusing in on a pair of volcanoes, such as Sierra Negra and Cerro Azul that have similar trace element and radiogenic isotope signatures but variable δ18O, and collecting both subaerial and submarine samples would provide a means to test hypotheses for the origin of the low δ18O component. If the light 18O component is a pyroxenite in the mantle source, this should produce olivine with enrichents in Ni and Si and depletions in Mn and Ca due to the formation of pyroxenite from eclogite melts reacting with peridotite (Sobolev et al., 2005, Sobolev et al., 2007, Vidito et al., 2013). Sierra Negra, which has shown lower δ18O values than Cerro Azul, would therefore be predicted to show a measurably higher Ni and Si content and lower Mn and Ca content than olivine from Cerro Azul. If the light component, however, is apart of the volcanic edifice or lithosphere then the Fo contents of olivine from Sierra Negra should not show unusual composition and be correlated with δ18O values (Wang & Eiler, 2008, Genske et al., 2013). If submarine samples from Sierra Negra with melt inclusions can be identified, it would also be interesting to combine oxygen isotope measurements with XANES measurements of Fe3+/ΣFe ratios. The Sierra Negra/Cerro Azul component examined in Chapter 3, showed the largest variation in Fe3+/ΣFe ratios of any of the groups with few ! 302! compelling correlations between Fe3+/ΣFe ratios and indicators of shallow level processes. However, the results of Chapter 4 suggest that melt from Sierra Negra has interacted with hydrothermally altered lithosphere. 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