MIAMI UNIVERSITY The Graduate School

Certificate for Approving the Dissertation

We hereby approve the Dissertation

of

Huimin Yu

Candidate for the Degree:

Doctor of Philosophy

Elisabeth Widom, Director

William K. Hart, Reader

Michael R. Brudzinski, Reader

Paul B. Tomascak, Reader

Michael W. Crowder, Graduate School Representative ABSTRACT

Li, Hf AND Os ISOTOPE SYSTEMATICS OF AZORES BASALTS AND A NEW MICROWAVE DIGESTION METHOD FOR Os ISOTOPIC ANALYSIS

by Huimin Yu

This dissertation includes three projects related to the isotope geochemistry of ocean island basalts (OIB) in the Azores archipelago. Detailed studies of Hf, Os and Li isotope systematics are combined with Sr-Nd-Pb isotopes and trace elements to investigate the nature and origin of mantle heterogeneity beneath the Azores. In addition, new microwave digestion methods have been developed and tested for dissolution of samples for Os isotope analysis. The first project focuses on Hf-Os isotope systematics of basalts from the Azores Central Group islands (Faial, Pico, São Jorge and Terceira) with HIMU and EM-type 187 188 signatures. Sub-chondritic Os/ Os and ΔεHf signatures on or slightly below the terrestrial εHf - εNd array indicate that the mantle sources of these basalts do not contain significant recycled crustal material. Rather, the sources of the basalts are interpreted to include variable and geographically controlled mixtures of a deeply derived enriched mantle plume, relatively depleted mantle similar to that beneath the Mid-Atlantic ridge, and recycled metasomatized mantle wedge. The second project focuses on assessing the utility of Li isotopes as a tracer of heterogeneous mantle sources. The δ7Li data of Central Group island (Faial, Pico and Terceira) and São Miguel basalts vary only slightly (+3.1 to +4.7‰), and are all within the range of normal MORB, despite large variations in radiogenic isotopes. Nevertheless, the Central Group island basalts have, on average, slightly higher δ7Li than São Miguel, and exhibit positive correlations with Sr and Os isotopes, and negative correlations with Pb, Nd and Hf isotopes, and are consistent with the interpretations of the Hf-Os isotope study. New diffusion modeling furthermore suggests that mantle heterogeneities induced by subduction processes may be maintained in the mantle for timescales of >2.5Ga. The third project assesses the utility of microwave digestion for Os isotopic analysis. Compared to conventional Carius tube digestions, microwave digestion is faster and safer, and allows for the use of HF to achieve complete dissolution of silicate samples. This study demonstrated that microwave digestions successfully achieve spike-sample equilibration, have acceptably low processing blanks, and produce yields >90%.

Li, Hf AND Os ISOTOPE SYSTEMATICS OF AZORES BASALTS AND A NEW MICROWAVE DIGESTION METHOD FOR Os ISOTOPIC ANALYSIS

A DISSERTATION

Submitted to the Faculty of

Miami University in partial

fulfillment of the requirements

for the degree of

Doctor of Philosophy

Department of Geology and Environmental Earth Science

by

Huimin Yu

Miami University

Oxford, Ohio

2011

Dissertation Director: Elisabeth Widom, Ph.D. TABLE OF CONTENTS

Chapter 1: Introduction Introduction 2 References 6

Chapter 2: Hafnium and osmium isotopic systematics of ocean island basalts in the Central Group islands of the Azores Archipelago Abstract 10

Body Text 11 References 39

Chapter 3: Lithium isotope systematics of basalts from the Azores Archipelago: constraints on the origin of the mantle sources Abstract 86 Body Text 88 References 114

Chapter 4: Microwave Digestion Method for Os analysis Abstract 155

Body Text 156 References 169

Chapter 5: Conclusions Summary 181 References 183

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LIST OF TABLES

Chapter 2: Hafnium and osmium isotopic systematics of ocean island basalts in the Central Group islands of the Azores Archipelago

1. Major element, trace element and isotopic ratios of the Central Group island basalts 48

2. The compositions of end-members in Os-Sr-Pb mixing models 59

3. The compositions of end-members in Nd-Hf mixing models 60

Chapter 3: Lithium isotope systematics of basalts from the Azores Archipelago: constraints on the origin of the mantle sources

1. Compositions of Azores basalts samples including lithium concentrations and isotopic compositions 125 2. δ7Li values of four olivine-whole rock pairs 127

Chapter 4: Microwave Digestion Method for Os analysis

1. Comparison of microwave digestion and Carius tube digestion 170

2. Osmium yields of microwave digestion 171

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LIST OF FIGURES

Chapter 2: Hafnium and osmium isotopic systematics of ocean island basalts in the Central Group islands of the Azores Archipelago

1. Map of Azores Archipelago 61 2. 87Sr/86Sr versus 206Pb/204Pb ratios for basalts of the Central Group islands, Azores archipelago 63

3. Total alkalis versus SiO2 and plot of Ni versus MgO for basalts from the Central Group islands 65 4. Variations of major element versus MgO for the basalts from the Central

Group islands 67 5. Chondrite-normalized REE patterns for the Central Group island basalts 69 6. Primitive-mantle normalized trace element patterns for the Central Group island basalts 71

7. Trace element compositions of the Central Group island basalts 73 8. Whole-rock isotope signatures of the Central Group island basalts 75

9. Osmium isotope signatures of the Central Group island basalts 77 10. Calculated mixing trends between enriched Azores mantle plume (AMP) and potential recycle crustal components 79 11. Neodymium and Hf isotope signatures of the Central Group island basalts and calculated mixing trends between enriched Azores mantle plume

(AMP) and potential recycle crustal components 81

12. Model of proposed mantle sources beneath the Azores archipelago 83

Chapter 3: Lithium isotope systematics of basalts from the Azores Archipelago: constraints on the origin of the mantle sources

1. Lithium isotope signatures of Earth reservoirs 128

2. Map of Azores Archipelago 130 87 86 206 204 3. Sr/ Sr versus Pb/ Pb for basalts of the Azores archipelago 132

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4. Deviation of individual analyses from the mean of replicate analyses of

all standards (L-SVEC 50ppb) 134 7 5. Comparison of δ Li values between whole rocks and olivine separates 136 6. Histogram of Li isotopic composition of the Azores basalts and MORB 138 7. δ7Li values of whole rock powder from the Central Group islands versus

radiogenic isotopes 140 7 87 86 8. δ Li versus Sr/ Sr of global OIB 142 9. Osmium concentrations and isotopic compositions of the Central Group

island basalts 144 7 10. Major and trace element compositions versus δ Li for Azores basalts 146 11. Isotopic ratios and major and trace element compositions of the Azores

basalts 148

12. Calculated mixing trends between basalt and pelagic sediment 150 7 13. Diffusion models for Li concentrations and δ Li of recycled material 152

Chapter 4: Microwave Digestion Method for Os analysis

1. The temperature and pressure change during heating of aqua regia in the microwave 172 2. The temperature and pressure change during heating of HF+EtOH+HCl in the microwave 174

3. Distillation system for Os separation 176 4. Osmium concentrations and 187Os/188Os ratios of sample aliquots

digested by Carius tube and microwave 178

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ACKNOWLEDGMENTS

I would like to express my sincerely gratitude to the many people who have helped me during my study at Miami University. The first and the foremost is my advisor Dr. Elisabeth Widom. Everything she taught me over last six years is extremely valuable in my research career and in my whole life. She encouraged me to develop independent thinking skills. I am really grateful to her patience of my scientific writing and guidance when I lost direction during my research. I respect her as an extremely knowledgeable geologist as well as a successful mentor. Thanks to all the people who have served on my committees: Dr. William K. Hart, Dr. Michael R. Brudzinski and Dr. Michael W. Crowder. I owe a special thanks to Dr. Paul B. Tomascak for making the time to come to Oxford for my defense. I would also like to thank Dave Kuentz and John Morton for their assistance in the labs in Miami University. Thank you to Lin Qiu and Roberta Rudnick at University, Richard Carlson at DTM for their assistance in their labs during my visit. I thank all faculty and staff in the Department of Geology at Miami University, especially Cathy Edwards and Jeanne Johnston for their assistance with day to day questions. I am thankful to Amy Gelinas and Qing Meng for their assistance with sample analysis and all of my lab mates, Elise Conte, George Daly, Rebecca Tortorello and Fara Rasoazanamparany to share ideas, discuss my research and help my English. I also would like to thank Shanshan Ji, Yun Luo, Qing Meng, Jing Zhang, Qiuyuan Huang and Shizuko Watanabe for their friendships. Last but not least, I would like to thank my family. My husband, Yongjian Han, always understands me, supports me and encourages me. My daughter Grace Han and my son Steven Han bring a lot of fun and are my source of happiness. I owe a special thanks to my parents Guangshan Yu and Wenqing Shi. Without their support and understanding, I could not finish my doctoral dissertation. These projects were supported by the National Science Foundation (NSF EAR # 0510598 and NSF EAR # 0549552 to Elisabeth Widom) and the Geological Society of America (8696-07 to Huimin Yu).

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CHAPTER 1

Introduction

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Ocean island basalts (OIB) exhibit large variations in their chemical and isotopic compositions, indicating that they are produced by melting of heterogeneous mantle sources and/or modified during ascent through the lithosphere by shallow processes such as melt-rock reaction or assimilation. Several isotopically distinct mantle signatures have been recognized, and it has been proposed that mixing of these mantle components can produce the isotopic heterogeneity observed amongst oceanic basalts globally. The proposed mantle components have been defined based on their end-member compositions in combined Sr-Nd-Pb isotope space, as well as the repeated occurrence of these extreme compositions globally, suggesting that they may represent discrete reservoirs produced by several distinct but common geologic processes, such as subduction of oceanic lithosphere with sediments. Constraining the origins and the locations of heterogeneous mantle components is significant for understanding geodynamic processes, and can potentially contribute insight into several outstanding but fundamental questions such as: how deeply are plates subducted (the upper mantle or core-mantle boundary?); what physical and chemical changes occur at the high temperature and pressure conditions experienced during subduction (mineral phase changes and dehydration reactions?); how do subduction processes (dewatering ± melting) affect element distribution and fluxing amongst Earth reservoirs; and what is the nature of mantle convection (layered-mantle convection versus whole-mantle convection?). However, the origins and locations of the inferred heterogeneous mantle components are still controversial. The Azores archipelago, located in the vicinity of the Mid-Atlantic Ridge (MAR, between ~37-40° N) and the triple junction between the North American, African and Eurasian plates, is an example of a location in which mantle plume-spreading ridge interactions occur, and it has long been recognized not only for tectonic complexities, but for the extreme isotopic heterogeneity and unusual compositions exhibited in basalts on both an archipelago-wide as well as an intra-island scale. Possible mechanisms to produce the isotopic variations in OIB mantle sources have been widely speculated on in the context of the Azores as well as globally, and include a variety or crustal recycling and metasomatic processes such as recycling of altered oceanic crust and sediment (e.g. Chase, 1981; White and Hofmann, 1982; Zindler and Hart, 1986), recycling of oceanic or continental lithospheric mantle (e.g. Sun and McDonough, 1989; Halliday et al., 1995; Niu and O’Hara, 2003; Stracke et al., 2003; Workman et al., 2004), or

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shallow-level interactions of magma with in situ metasomatized mantle lithosphere (e.g. Chauvel et al., 1997; Donnelly et al., 2004; Millet et al., 2009). Most of the previous studies in the Azores have focused primarily on Sr-Nd-Pb isotope variations in the Azores basalts, which are useful for identifying mantle heterogeneity and constraining the number and compositional range of distinct mantle components beneath the Azores. However, difficulties arise in constraining the processes that resulted in the observed heterogeneities using these isotope systems alone, because the parent and daughter elements in these isotope systems are highly incompatible during melting and highly to moderately fluid mobile, leaving ambiguities in the expected parent-daughter ratios following subduction dewatering and/or partial melting of recycled crustal materials. In addition, potentially variable storage times of such materials in the mantle also can lead to present-day isotopic variability, further compounding the difficulties in uniquely identifying the processes responsible for the observed mantle heterogeneities. This dissertation focuses on the application of the radiogenic Hf and Os isotope systems as well as the stable Li isotope system, which are complementary to the Sr-Nd-Pb isotope systems in that they exhibit distinct geochemical behaviors, and thus have the potential to provide unique constraints on the nature and origin of the heterogeneous mantle source components identified in the Azores. This dissertation comprises three sub-projects, two of which address respectively Hf-Os and Li isotope systematics in Azores basalts, and one of which investigates a new dissolution method for Os isotopic analysis of geologic materials including very low Os abundance basalts such as those in the Azores. The first project (Chapter 2) of this dissertation is focused on the Re-Os and Lu-Hf isotope systematics of basalts from four Central Group islands (Faial, Pico, São Jorge and Terceira) of the Azores archipelago, combined with Sr, Nd, Pb isotope and trace element systematics to assess the potential origins of their mantle sources. The Re-Os isotopic system differs from other radiogenic isotope systems in that the parent element Re is a mildly incompatible element during melting, but the daughter element Os is a strongly compatible element that is preferentially retained in the solid residue during mantle melting. Crustal reservoirs thus have much higher Re/Os ratios than depleted mantle, producing differences in 187Os/188Os between crustal and mantle reservoirs that are more extreme than that observed in any other radiogenic isotope system. Therefore, the Os isotopic system provides a complementary tracer for identifying

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potential ancient recycled crustal materials in a mantle source. Furthermore, because Re is extracted from the mantle during melt depletion, ancient lithospheric mantle that experienced a high degree of melt depletion will have distinctly unradiogenic 187Os/188Os ratios (≈0.105-0.127; Walker et al., 1989; Pearson et al., 1995) that, because of the high Os concentrations in the mantle, are not easily modified by subsequent metasomatic processes that may overprint highly incompatible element isotope systems such as Sr, Nd, and Pb. The Os isotope system therefore also can serve as a unique and diagnostic indicator of involvement of ancient depleted recycled mantle in OIB sources. The Lu-Hf isotopic system also exhibits distinct behavior compared to the Sr, Nd and Pb isotope systems, in that the parent element Lu is a heavy rare earth element (HREE), but the daughter element Hf is a high field strength element (HFSE). The distinct geochemical behaviors of Lu and Hf are particularly important in petrogenetic processes that involve the mineral phases garnet or zircon, which have strong preferences for Lu over Hf and Hf over Lu, respectively. Therefore, the Lu-Hf isotope system should be ideally suited as a tracer for discerning the involvement in an OIB mantle source of a variety of potential recycled lithospheric materials, including zircon-bearing terrigenous sediment and garnet-bearing lithospheric mantle, which would develop anomalously low and high 176Hf/177Hf ratios with time, respectively (Vervoort et al., 1999; Salter & White, 1998). The investigation of Hf and Os isotope systematics in the Azores basalts, combined with Sr, Nd, and Pb isotope and trace element systematics, offers the potential to further constrain the origins of the heterogeneous mantle sources of the Central Group islands. The second project (Chapter 3) of this dissertation focuses on Li isotope systematics of basalts from three Central Group islands (Faial, Pico and Terceira) and one of the Eastern Group islands, São Miguel, which has extreme Sr, Nd and Pb isotopic compositions that are distinct from the signatures observed in the Central Group islands and all other OIB. This project aims to test the utility of Li isotopes as a tracer of recycled material and to further evaluate the origins of the mantle sources of these islands by combining Li isotopes with radiogenic isotopes. Lithium isotopes undergo large isotopic fractionations in low-temperature environments, but are relatively homogeneous in the depleted mantle due to limited fractionation at high temperature. Lithium isotopes may thus be a useful tracer for recycled crustal material, because altered oceanic crust and sediments have isotopically higher δ7Li values than depleted mantle due to interaction with heavy-Li sea water (δ7Li = +31‰; Millot et al., 2004), and terrigenous

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sediments have lower δ7Li which correlate with weathering intensity (Rudnick et al., 2004). In addition, in contrast to radiogenic isotopes, Li isotopic compositions are not affected by radioactive decay, so that variable storage times of recycled materials in the mantle should not necessarily affect their isotopic fingerprint. However, some studies suggest that lithium diffuses rapidly in co-existing phases of peridotites and mafic magmas, and thus predict that any heterogeneities in δ7Li imparted to the mantle by recycled material would be erased over very short geologic timescales (perhaps as short as a few years) at magmatic temperatures (Jeffcoate et al., 2007; Parkinson et al., 2007). This is, however, inconsistent with the variable Li isotopic compositions that have been reported recently in MORB (e.g. Elliott et al., 2006; Tomascak et al., 2008; Simons, 2008; Simons et al., 2010) and OIB (e.g. Ryan and Kyle, 2004, Nishio et al., 2005; Chan et al., 2009). The applicability of Li isotopes as a tracer of recycled material in the mantle is thus clearly a complex and unresolved issue, and the present study of Li isotopes in Azores basalts combined with new diffusion modeling aims to provide new constraints. The third project (chapter 4) of this dissertation involves the development of a new digestion method for Os isotopic analysis. Previous digestion methods used for Os isotopic analysis include Teflon bomb digestion, Carius tube digestion and high pressure asher (HPA-S) digestion. These methods generally require long digestion times of up to 48 hours, and the most widely used of these methods (Carius tube digestion) requires custom made and costly consumables (e.g. Carius tubes) as well as equipment that is incompatible with a clean lab (e.g. gas cylinders and torches). In addition, the Carius tube and high pressure asher digestion methods prevent the use of HF due to the glass vessels employed in those digestion procedures, thus it is not possible with these methods to achieve complete digestion of silicate rock samples, for which it is necessary to break the Si-O bonds. In this chapter, a new microwave digestion method is explored and is demonstrated to have high recoveries of Os and low blank levels. This new digestion method is faster, safer and easier compared with previous methods, and enables the use of HF during the digestion process. This project also involved the development of a modified “macro” distillation method for Os purification, based on that developed by Nägler and Frei (1997). The modified technique enable simultaneous purification of multiple (8 or more) samples, under controlled conditions that produce yields of >70%, and that can be easily employed and modified for a variety of geological and cosmochemical applications.

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References Chan, L.H., Lassiter, J.C., Hauri, E.H., Hart, S.R., Blusztajn, J. 2009. Lithium isotope systematics of lavas from the Cook–Austral Islands: Constraints on the origin of HIMU mantle. Earth Planet. Sci. Lett. 277 (3-4), 433–442. Chase, C.G., 1981. Oceanic island Pb: two-stage histories and mantle evolution. Earth Planet. Sci. Lett. 52 (2), 277-284. Chauvel, C., McDonough, W., Guille, G., Maury, R., Duncan, R., 1997. Contrasting old and young volcanism in Rurutu Island, Austral chain. Chem. Geol. 139 (1-4), 125– 143. Donnelly, K.E., Goldstein, S.L., Langmuir, C.H., Spiegelman, M., 2004. Origin of enriched ocean ridge basalts and implications for mantle dynamics. Earth Planet. Sci. Lett. 226 (3-4), 347-366. Elliott, T., Thomas, A., Jeffcoate, A., Niu, Y.L., 2006. Lithium isotope evidence for subduction-enriched mantle in the source of mid-ocean-ridge basalts. Nature 443 (7111), 565-568. Halliday, A.N., Lee, D.C., Tommasini, S., Davies, G.R., Paslick, C.R., Fitton, J.G., James, D.E., 1995. Incompatible trace elements in OIB and MORB and source enrichment in the sub-oceanic mantle. Earth Planet. Sci. Lett. 133 (3-4), 379–395. Hart, S.R., Hauri, E.H., Oschmann, L.A., Whitehead, J.A., 1992. Mantle plumes and entrainment: isotopic evidence. Science 256 (5056), 517-520. Jeffcoate, A.B., Elliott, T., Kasemann, S.A., Ionov, D., Cooper, K., Brooker, R., 2007. Li isotope fractionation in peridotites and mafic melts. Geochim. Cosmochim. Acta 71(1), 202-218. Millet, M.A., Doucelance, R., Baker, J.A., Schiano, P., 2009. Reconsidering the origins of isotopic variations in ocean island basalts: insights from fine-scale study of São Jorge Island, Azores archipelago. Chem. Geol. 265 (3-5), 289–302. Millot, R., Guerrot, C., Vigier, N., 2004. Accurate and High-Precision Measurement of Lithium Isotopes in Two Reference Materials by MC-ICP-MS. Geostandards and Geoanalytical Research, 28: 153–159. doi: 10.1111/j.1751-908X.2004.tb01052.x Nägler, T.F., Frei, R., 1997. “Plug in” Os distillation. Schweiz. Mineral. Petrogr. Mitt. 77, 123-127. Nishio, Y., Nakai, S., Kogiso, T., Barsczus, H.G., 2005. Lithium, strontium, and neodymium isotopic compositions of oceanic island basalts in the polynesian region: constraints on a polynesian HIMU origin. Geochem. J. 39 (1), 91-103. Niu, Y., O’Hara, M.J., 2003. Origin of ocean island basalts: a new perspective from petrology, geochemistry and mineral physics considerations. J. Geophys. Res. 108, 2209. doi:10.1029//2002JB002048. Parkinson, I.J., Hammond, S.J., James, R.H., Rogers, N.W., 2007. High-temperature lithium isotope fractionation: insights from lithium isotope diffusion in magmatic systems. Earth Planet. Sci. Lett. 257 (3-4), 609-621.

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Pearson, D.G., Shirey, S.B., Carlson, R.W., Boyd, F.R., Pokhilenko, N.P., Shimizu, N., 1995. Re-Os, Sm-Nd, and Rb-Sr isotope evidence for thick Archaean lithospheric mantle beneath the Siberian craton modified by multi-stage metasomatism. Geochim. Cosmochim. Acta 59 (5), 959-977. Rudnick, R.L., Tomascak, P.B., Njo, H.B., Gardner, L.R., 2004. Extreme lithium isotopic fractionation during continental weathering revealed in saprolites from South Carolina. Chem. Geol. 212 (1-2), 45-57. Ryan, J.G., Kyle, P.R., 2004. Lithium abundance and lithium isotope variations in mantle sources: insights from intraplate volcanic rocks from Ross Island and Marie Byrd Land (Antarctica) and other oceanic islands. Chem. Geol. 212 (1-2), 125-142. Salters, V.J.M., White, W.M., 1998. Hf constraints on mantle evolution. Chem. Geol. 145 (3-4), 447-460. Simons, K.K., Dixon, J.E., Kingsley, R.H., Holk, G., 2010. Relationship between water and hydrogen isotopes in mantle end-members: preliminary data from the Azores Platform. Goldschmidt 2010, Abstract (A964). Simons, K.K., 2008. Lithium isotopes variability: new constraints on mantle heterogeneity. Ph.D. dissertation, Columbia University. Stracke, A., Bizimis, M., Salters, V.J.M., 2003. Recycling oceanic crust: quantitative constraints. Geochem. Geophys. Geosyst. 4, 8003. doi:10.1029/2001GC000223. Stracke, A., Hofmann, A.W., Hart, S.R., 2005. FOZO, HIMU and the rest of the mantle zoo. Geochem. Geophys. Geosyst. 6, Q05007, doi:10.1029/2004GC000824. Sun, S.S., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: Saunders, A.D., Norry, M.J. (Eds.), Magmatism in the Ocean Basins. Geol. Soc. Am. Spec. Publ. 42, 313–345. Tomascak, P.B., 2004. Developments in the Understanding and Application of Lithium Isotopes in the Earth and Planetary Sciences. Rev. Mineral. Geochem. 55, 153-195. Tomascak, P.B., Langmuir, C.H., le Roux, P.J., Shirey, S.B., 2008. Lithium isotopes in global mid-ocean ridge basalts. Geochim. Cosmochim. Acta 72(6), 1626-1637. Vervoort, J.D., Patchett, J.P., Blichert-Toft, J., Albarede, F., 1999. Relationships between Lu-Hf and Sm-Nd isotopic systems in the global sedimentary system. Earth Planet. Sci. Lett. 168 (1-2), 79-99. Walker, R.J., Carlson, R.W., Shirey, S.B., Boyd, F.R., 1989. Os, Sr, Nd, and Pb isotope systematics of Southern African peridotite xenoliths: implications for the chemical evolution of subcontinental mantle. Geochim. Cosmochim. Acta 53 (7), 1583-1595. White, W.M., Hofmann, A.W., 1982. Sr and Nd isotope geochemistry of oceanic basalts and mantle evolution. Nature 296, 821-825. Workman, R.K., Hart, S.R., Jackson, M., Regelous, M., Farley, K.A., Blusztajn, J., Kurz, M., Staudigel, H., 2004. Recycled metasomatized lithopshere as the origin of the Enriched Mantle II (EMII) end-member: evidence from the Samoan volcanic chain. Geochem. Geophys. Geosyst. 5, Q04008, doi:10.1029/2003GC000623. 7

Zindler, A., Hart. S., 1986. Chemical Geodynamics. Annu. Rev. Earth Planet. Sci. 14, 493-571.

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CHAPTER 2

Hafnium and osmium isotopic systematics of ocean island basalts in the Central Group islands of the Azores Archipelago

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Abstract Ocean island basalts of the Azores archipelago are thought to be produced by hotspot-ridge interaction, and exhibit large chemical and isotopic variations from depleted mantle to HIMU and EM-type signatures. We have analyzed Hf, Os, Sr, Nd and Pb isotopes and trace elements in a suite of young, fresh basalts from four Central Group islands of the Azores archipelago (São Jorge, Terceira, Faial and Pico) to investigate the origins of the heterogeneity. Lack of correlation between radiogenic isotope signatures and MgO suggests that the isotopic variations reflect mantle source heterogeneity. Basalts with HIMU and EM-type signatures in the Central Group island basalts have sub-chondritic 187Os/188Os ratios that are not easily reconciled with significant recycled crustal material in their mantle source. Lack of anomalously low εHf, εNd or

ΔεHf (<-3) of these basalts further argues against a mantle source with a significant ancient recycled crustal component. Basalts from Terceira and some from São Jorge have highly variable and radiogenic 206Pb/204Pb with negative ∆7/4Pb and a narrow range of 87Sr/86Sr. Sub-chondritic 187Os/188Os signatures and ∆εHf≈0 indicate that the mantle source of these lavas does not contain significant recycled oceanic crust, but is consistent with a mantle plume source with a FOZO signature. Faial lavas exhibit an EM-like signature, with positive Δ7/4Pb values, lower 206Pb/204Pb and higher 87Sr/86Sr than São Jorge and Terceira samples, but lower 87Sr/86Sr than lavas with extreme 187 188 EM signatures. The sub-chondritic Os/ Os ratios, slightly negative ΔεHf, and slightly elevated 87Sr/86Sr ratios of Faial lavas indicate limited to no contribution to their source of recycled oceanic crust or pelagic or terrigenous sediments, but are consistent with a component of recycled metasomatized mantle wedge in their source. Pico basalts exhibit a mixing trend between the Faial EM end-member and the Terceira FOZO end-member, suggesting that the mantle source of Pico lavas contains a FOZO plume component, as well as recycled metasomatized mantle wedge. This study has important implications regarding the complexity of processes that lead to development of mantle heterogeneity over small spatial scales, and suggests an important role for metasomatic processes rather than recycled crustal components in the mantle sources of some ocean island basalts.

Keywords: Ocean Island Basalts; Azores; Faial; Pico; São Jorge; Terceira; geochemistry; radiogenic isotopes; trace elements; plume; lithosphere

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1. Introduction Understanding the origin of chemical and isotopic variability of ocean island basalts (OIB) and mid-ocean ridge basalts (MORB) is crucial for deciphering the compositional variability and evolution of the mantle. OIB have chemical and isotopic characteristics distinct from MORB, and most are thought to be generated by hot mantle plumes that are rooted either at the base of the upper mantle or the D” layer at the base of the lower mantle. Based on the characteristics of radiogenic isotope systems (Sr-Nd-Pb), Zindler and Hart (1986) defined four primary isotopic components observed in the oceanic basalt database: DMM (depleted MORB mantle), HIMU (high 206Pb/204Pb with low 87Sr/86Sr), EM1 (high 87Sr/86Sr and low 206Pb/204Pb) and EM2 (high 87Sr/86Sr and intermediate 206Pb/204Pb). Hart et al. (1992) and Stracke et al. (2005) subsequently defined an additional mantle component, FOZO (intermediate 87Sr/86Sr, 143Nd/144Nd, 206Pb/204Pb and high 3He/4He) which was proposed to come from the lower mantle based on it less degassed, high 3He/4He characteristics. The cause of mantle heterogeneity in the source of OIB, however, is still unresolved. The most attractive and commonly invoked mechanism to produce the variation of radiogenic isotopes is the involvement of recycled crustal materials in the mantle source of OIB (e.g. Chase, 1981; White and Hofmann, 1982; Zindler and Hart, 1986). However, an alternative to recycled oceanic crust and sediment in the source of mantle plumes is the involvement of oceanic or continental lithospheric mantle as possible reservoirs for OIB sources (e.g. Sun and McDonough, 1989; Halliday et al., 1995; Niu and O’Hara, 2003; Stracke et al., 2003; Workman et al., 2004). Some recent observations further indicate that the heterogeneous isotopic signatures of some ocean island basalts may reflect shallow-level magma interactions with in situ metasomatized mantle lithosphere (e.g. Chauvel et al., 1997; Donnelly et al., 2004; Millet et al., 2009). Among ocean island basalts globally, the Azores basalts are geochemically unusual (Fig. 1), and exhibit extreme isotopic heterogeneity on both an archipelago-wide as well as an intra-island scale (Fig. 2). Previous research (Hawkesworth et al., 1979; White et al., 1979; Dupré et al., 1982; Davies et al., 1989; Widom and Shirey, 1996; Turner et al., 1997; Widom et al., 1997; Moreira et al., 1999; Schaefer et al., 2002; Widom and Farquhar, 2003; Madureira et al., 2005; Beier et al., 2007; Elliott et al., 2007; Turner et al., 2007; Beier et al., 2008; Jean-Baptiste et al., 2009) has suggested the involvement of at least four source components to explain the isotopic variations in the Azores basalts: local MORB mantle; a regional plume component with

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radiogenic Pb that plots below the Northern Hemisphere Reference Line (NHRL; Hart, 1984) and is exhibited by São Jorge and Terceira samples (young HIMU or FOZO); an extremely enriched component exhibited by São Miguel samples with enriched Sr-Nd isotopic ratios and radiogenic Pb that plots above the NHRL (intermediate between HIMU and EMII signatures); and a component with less radiogenic Pb isotopes and slightly enriched Sr-Nd isotopes (EMI/II signature) that also plot above the NHRL. Millet et al. (2009) proposed that an additional E-MORB component resides in the oceanic basement under São Jorge. However, the origin of these enriched components is still not clear and a variety of possibilities have been proposed, including the influence of a heterogeneous mantle plume with recycled oceanic crust and/or mantle lithosphere (Turner et al., 1997, 2007; Moreira et al., 1999; Widom and Farquhar, 2003; Beier et al., 2007; Elliott et al., 2007; Schaefer et al., 2002; Jean-Baptiste et al., 2009; Prytulak and Elliott, 2009), delaminated metasomatized sub-continental lithospheric mantle (Turner et al., 1997; Widom et al., 1997; Widom and Farquhar, 2003; Millet et al., 2009), and shallow lithospheric processes (Millet et al., 2009). Most of the previous studies have focused primarily on Sr-Nd-Pb isotope variations in the Azores basalts, which are useful for identifying mantle heterogeneity and constraining the number and compositional range of distinct mantle components beneath the Azores. However, difficulties arise in constraining the processes that resulted in the observed heterogeneities using these isotope systems alone, because the parent and daughter elements in these isotope systems are highly incompatible during melting and highly to moderately fluid mobile, leaving ambiguities in the expected parent-daughter ratios following subduction dewatering and/or partial melting of recycled crustal materials. In addition, potentially variable storage times of such materials in the mantle also can lead to present-day isotopic variability, further compounding the difficulties in uniquely identifying the processes responsible for the observed mantle heterogeneities. In this study, we investigate the systematics of two additional isotope systems, Re-Os and Lu-Hf, for which the geochemical behaviors of the parent and daughter elements are distinct from the Rb-Sr, Sm-Nd, and U-Th-Pb isotope systems. The fluid immobile behavior of Lu and especially Hf eliminate some of the ambiguities associated with subduction dewatering of recycled crustal materials, and the large fractionations of these elements in the presence of garnet and zircon make this isotope system particularly diagnostic for evaluating the potential roles of

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recycled terrigenous sediment components and deep-seated recycled or metasomatic components involving garnet. The compatible behavior of Os coupled with the moderately incompatible behavior of the parent element Re lead to extreme differences in the isotopic composition of crustal versus mantle reservoirs, making this isotope system an ideal tracer of involvement of crustal material, either as a recycled component in the mantle or as a shallow assimilation. In this study, Os and Hf isotope systematics are combined with Sr, Nd, Pb isotope and trace element systematics to further evaluate the potential mantle sources and processes leading to heterogeneity in basalts from four of the Azores Central Group islands (Faial, Pico, São Jorge and Terceira), each of which exhibits distinct and variable Sr-Nd-Pb isotopic characteristics.

2. Geological and geochemical setting The Azores archipelago is located in the vicinity of the Mid-Atlantic Ridge (MAR, between ~37-40° N) and the triple junction between the North American, African and Eurasian plates (Fig. 1a). This archipelago is a group of nine islands distributed on both sides of the Mid-Atlantic Ridge. Two islands, Corvo and Flores, lie to the west of the ridge; the other seven islands, including Faial, Graciosa, Pico, São Miguel, São Jorge, Santa Maria, and Terceira, lie to the east of the ridge. To the east of the ridge is a spreading center called the “Terceira Rift”, along which three islands, Graciosa, São Miguel and Terceira, are aligned. The other four islands, including Faial, Pico, São Jorge and Santa Maria, are to the south of this rift. The nine islands of the Azores archipelago can be divided into three groups: the Occidental Group including Corvo and Flores islands; the Central Group including Faial, Graciosa, Pico, São Jorge and Terceira; and the Oriental Group including São Miguel and Santa Maria. The tholeiitic basalts from the MAR along the Azores Platform have distinct geochemical characteristics compared to basalts along normal ridge segments. The Azores Platform tholeiites (APT) are enriched in light rare earth elements (LREE) and large-ion-lithophile elements (LILE) (Schilling, 1975). They also have elevated 87Sr/86Sr (White et al., 1976; White and Schilling, 1978), high 206Pb/204Pb and low 143Nd/144Nd (Dupré and Allè, 1980, 1983; Ito et al., 1987; Dosso et al., 1996). These signatures are interpreted as evidence for an enriched mantle plume beneath the MAR in the region of the Azores. Additional evidence in support of ridge-hot spot interaction in the Azores area includes the topographically elevated spreading ridge (Vogt, 1976; Schilling, 1985; Gente, 1987; Thibaud et al., 1998; Beier et al., 2010), geochemical characteristics of the

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APT (White et al., 1976; Bougault and Treuil, 1980; Schilling et al., 1983; Dosso et al., 1999; Moreira et al., 1999; Simons, 2008; Jean-Baptiste et al., 2009; Millet et al., 2009), and negative gravity anomalies (Detrick et al., 1995; Thibaud et al., 1998; Gente et al., 2003). Recent studies suggest that the mantle plume is currently underlying Terceira Island (Gente et al., 2003; Yang et al., 2006), but the depth of origin of the mantle plume is still controversial (Montelli et al., 2004, 2006; Silveira et al., 2006; Yang et al., 2006). Montelli et al. (2004, 2006) have proposed that the mantle plume originates from the core-mantle boundary, but other studies suggest that it has a relatively shallow origin (250-400 km) within the upper mantle (Silveira et al., 2006; Yang et al., 2006). Gravity and bathymetric data combined with seafloor isochrons deduced from magnetic anomalies indicate that the Azores Platform formed between 20Ma to 7Ma, with the plateau rift developing during migration of the MAR away from the hotspot (Gente et al., 2003). Consistent with a mantle plume, the ages of the islands generally increase with the distance from the ridge, with the oldest basalts found on the eastern-most islands of Santa Maria (8Ma) and São Miguel (4Ma) (Abdel-Monem et al., 1975; Fernandez, 1980; Feraud et al., 1980; Moore, 1990; Moreira et al., 1999). However, the detailed age progression for the archipelago and for each island is complex, and clearly controlled partly by local tectonics (Feraud et al., 1980). Active volcanism occurs across much of the Azores Archipelago, and 26 historical eruptions have been reported in the Azores including submarine eruptions as well as subaerial eruptions on São Miguel, Terceira, São Jorge, Faial and Pico (Pinto Riberia, 2011). The most recent volcanism was the submarine eruption of Seretta, 8.5 km west of Terceira, between 1998-2000 (Forjaz et al., 2001).

3. Samples and analytical methods In this study, we analyzed 74 basalts from the Central Group islands, including 14 from Faial, 11 from Pico, 18 from Terceira and 31 from São Jorge (WAF, WAP, WAT and WASJ series respectively). The São Jorge samples are separated into two groups (Fig. 1b) based on K-Ar and Ar-Ar ages and geographic location, including those from Topo, the oldest formation on the east end of the island (~550Ka to 1.32Ma, Feraud et al., 1980; Hildenbrand et al., 2008), and those from Rosais and Manadas, the intermediate and youngest formations (368Ka to historical volcanic activity; Hildenbrand et al., 2008) on the west end and middle of the island, respectively.

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The samples in this study are primarily alkali basalts with variable crystal contents. Pico samples and some Faial samples are olivine rich (>10% olivine), and most are also clinopyroxene (cpx) rich (>20% cpx). Samples from São Jorge, Terceira and others from Faial are less crystal-rich with <10% by volume crystals, including olivine and cpx. In addition, most samples in this study have microcrystalline matrices including feldspar microphenocrysts between ~50 to 250μm. Whole rock samples were sawed into ~0.5 cm thick slabs, and all surfaces were ground using Si-C paper to remove any metallic residue from the saw blade. Sample slabs were then cleaned in 18.2 mega-Ω E-pure water, including multiple sessions in the ultrasonic bath with rinsing steps in between. After drying in an oven at 110 ˚C, the slabs were wrapped in paper and plastic to avoid metal contamination, and hammered to produce fragments from 0.5 cm to 1 cm. Samples were then chipped in a Sepor alumina jaw crusher, and powdered in a Spex shatterbox using a high purity alumina vessel. Major and trace element analysis, chemical separations and isotopic measurements with the exception of Hf isotopes were all done at in Miami University. Hafnium isotopes were separated and purified in the clean lab at Miami University, and isotopic ratios were measured at the Department of Terrestrial Magnetism (DTM), Carnegie Institution of Washington. Concentrations of major element and Sc, Ba and Cr were obtained by Beckman SpectraSpan V Direct-Current Plasma Atomic Emission Spectrometer (DCP-AES) following the procedures of Katoh et al. (1999). A mixture of 200 mg whole rock powder and 600 mg Li2B4O7 flux was fused at 950 ˚C for 20 minutes, and then dissolved in 5% HNO3. During analysis, a blank and ten international rock standards were used as external standards for calibration. Additional trace element concentrations were obtained by inductively coupled plasma mass spectrometer (ICP-MS) using a Varian quadrupole instrument. With the exception of Pb concentration analysis, samples were dissolved by flux fusion. A mixture of 50 mg whole rock powder and 75 mg of Na-K flux (Na-tetraborate and K-carbonate from Alfa Aesar – Puratronic grade in a 3:2 mixture) was fused in a furnace for 30 minutes at 950°C, and then dissolved in 2% HNO3 and analyzed with 10 external standards and a blank. For Pb concentration analysis, samples were dissolved by acid digestion using high purity, concentrated HF-HNO3. Approximately 50 mg of whole rock powder was dissolved with concentrated HF-HNO3, and following dry down the samples were dissolved in 2% HNO3. Lead concentrations were determined by the standard

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additions method, in which each sample solution was analyzed with two standard addition solutions and a blank. The Sr, Pb, and bulk rare earth element (REE) separations followed the procedures of Snyder (2005). Approximately 100-200 mg of whole rock powder was dissolved in concentrated

HF–HNO3, and then taken up in HBr. Lead was separated and purified by anion exchange chemistry using HBr and HNO3. The residue was further processed for Sr separation by EiChrom Sr-Spec resin following the procedures of Deniel and Pin (2001). Residue from the Sr-Spec column was passed through a cation exchange column, and finally the Sm–Nd separations were performed following the procedures of Pin and Zalduegui (1997) by EiChrom Ln-Spec resin. Isotopic compositions for Sr, Nd and Pb were measured by thermal ionization mass spectrometry (TIMS) on a Finnigan Triton. Purified Pb was loaded on single Re filaments with silica gel and phosphoric acid, and run at 1200℃. Based on the deviations of NBS 981 from Todt et al. (1996), 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb were corrected for fractionation by 0.11% per amu (atomic mass unit). External reproducibility (2SD) for 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb are ±0.015, ±0.020, and ±0.063, respectively, based on long-term, repeat measurements of NBS 981 (n=105). Purified Sr was loaded on single Ta filaments with Ta2O5, and run at ~1400℃. Strontium isotopic ratios were corrected for mass fractionation based on 86Sr/88Sr=0.1194 (exponential law), and the mean value on measurements of NBS 987 (n=106) was 87Sr/86Sr=0.710239±0.000014 (2SD). Purified Nd was loaded onto double Re filaments with phosphoric acid, and run at ~1700℃. Neodymium isotopic ratios were corrected for mass fractionation (exponential law) to 146Nd/144Nd=0.7219, and the mean value for the La Jolla standard was 143Nd/144Nd=0.511846±0.000007 (2SD, n=90). The Re-Os separation chemistry followed the methods described in Shirey and Walker (1995) and Nägler and Frei (1997). Whole rock powder and spike were dissolved and equilibrated in Carius tube by aqua regia at 240 ˚C for 24 to 48 hours. Sample solutions were then transferred from chilled Carius tubes into a distillation system to distill Os from the aqua regia digestion solutions (heated to 115 ˚C) into chilled, concentrated HBr. After 150 minutes of distillation, the Os-bearing HBr trapping solutions were covered and heated at 90 ˚C for two 2- hours to complete reduction of Os to the stable OsBr6 complex. The HBr solutions were then evaporated and purified by a micro-distillation procedure following the methods of Roy-Barman (1993) and Roy-Barman and Allègre (1995). The Re-Os isotopic compositions were analyzed as 16

- OsO3 by negative thermal ionization mass spectrometer (NTIMS) on a Finnigan Triton, and corrected for oxygen isotopes and mass fractionation. The oxygen correction used the isotopic abundances of Nier (1950) and the mass fractionation correction normalized to 192Os/188Os=3.0826. Internal measurement errors range from 0.5% to 2%, and blanks were 0.2 pg. Hafnium was also separated and purified at Miami University. Whole rock powder (150 mg) was dissolved in concentrated HF+HNO3 acid (2:1) in capped Teflon Savillex beakers, heated overnight on a hotplate at 80 ˚C. After drying down, 0.5ml concentrated HNO3 was added to the samples and evaporated, and these steps repeated. Approximately1-2ml concentrated HCl was then added, and capped beakers heated on the hotplate for at least 2 hours, followed by dilution with 18.2 mega-Ω E-pure water. Samples were capped and heated on the hotplate overnight to ensure complete dissolution. Once the sample solutions were completely clear, they were dried down, and redissolved in 1ml 0.5N HCl for chemical separation. The chemical separation process followed the 2-stage method of Connelly et al. (2006), including cation exchange columns followed by EiChrom TODGA chemistry. The cation resin (AG50W×8, 100-200 mesh) was loaded into 2ml columns, and cleaned with 20ml 6N HCl. After preconditioning the resin with 4ml of 0.5N HCl, the samples were loaded in 1ml 0.5N HCl. High field strength elements (HFSE) including Hf and Ti were collected with 7ml 0.5N HCl + 0.15N HF, and dried down.

The samples were then redissolved in 0.85ml 3.5N HNO3–0.06N H3BO3 to prepare for TODGA chemistry. Columns were filled with 0.2ml TODGA resin, and washed with 2ml 10.5N HNO3,

6ml 1N HNO3-0.35N HF and 3ml 0.05N HCl. The residue of HF was removed from the resin with 1ml 3.5N HNO3-0.45N H3BO3 and the resin preconditioned with 2ml 3.5N HNO3. The

HFSE solution was then loaded onto the columns in 0.85ml 3.5N HNO3-0.06N H3BO3. After washing Ti from the resin with 3.5ml 3.5N HNO3, Hf was collected with 6ml 1N HNO3-0.35N

HF. The Hf samples were finally dried and re-dissolved with 0.4N HNO3 (~2ml) to prepare for analysis by Multi-Collector Inductively Coupled Plasma Mass Spectrometery (MC-ICP-MS) using a VG P-54 at DTM. During the course of analysis, the 176Hf/177Hf value for the JMC-475 Hf standard was 0.282126±0.000014 (2SD, n=37).

4. Results 4.1 Major and trace elements

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Major element, trace element and isotopic compositions of samples in this study are reported in Table 1. On an alkali vs. SiO2 classification diagram, most samples are basalts or trachybasalts; two samples plot in the area of basaltic trachyandesite and one in the basanite field (Fig. 3a). The

Azores samples range in SiO2 from 43.5 to 53.2 wt.% and MgO from 2.49 to 17.04 wt.% (Table 1), and exhibit a trend of decreasing Ni with decreasing MgO (Fig. 3b). Primitive magmas in the Azores have been inferred to have MgO contents of ~12% (Beier et al., 2006), and samples with MgO higher than 12 wt.% are considered to result from accumulation of olivine and clinopyroxene, and samples with MgO contents below 12 wt.% to have been affected by fractional crystallization. Fig. 3b shows that most of the Pico samples have substantially higher MgO wt.% than the other samples (up to 17.04 wt.%), consistent with their high crystal contents (ankaramites with >20% olivine and clinopyroxene crystals). Variations of major element versus

MgO contents are shown in Fig. 4. CaO exhibits a positive correlation with MgO; Al2O3, Na2O,

K2O, Al2O3/CaO and to some extent total Fe display negative correlations with MgO. For high

MgO samples, TiO2 increases with decreasing MgO, but for samples with <5 wt.% MgO, TiO2 and total Fe decrease with decreasing MgO, indicating the onset of titanomagnetite fractionation.

Based on K2O contents, two compositional groups can be distinguished among the Central

Group islands: samples from São Jorge and Terceira have lower K2O contents and samples from

Faial and Pico have higher K2O contents for a given MgO content. Terceira samples also have lower Al2O3 concentrations at a given MgO content compared to the other three Central Group islands. Chondrite-normalized rare earth elements (REE) patterns of the Central Group island samples are shown in Fig. 5. The samples from the different Central Group islands are similar in their REE patterns, and all of them exhibit strong light REE (LREE) enrichments with (La/Yb)N ranging from 6-14. However, the moderately incompatible middle REE (MREE) are less enriched, and ratios such as (Gd/Yb)N ratios are nearly constant in all samples (2-3). The most notable difference amongst the Central Group islands is that some Faial and São Jorge samples have stronger enrichments in the LREE than samples from Pico and Terceira. For example,

(Ce/Yb)N ratios range from 6.5-11 in Faial and São Jorge samples, generally higher than the range of 6-9 in Pico and Terceira. Primitive mantle-normalized trace element patterns for the Central Group islands are compared with average N-MORB, E-MORB, HIMU, EMI and EMII-type basalts in Fig. 6. Trace elements patterns of the Faial samples are similar to those of

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EMI/EMII basalts, with more enriched Cs, Rb and Ba (higher Cs/Yb, Rb/Yb and Ba/Yb ratios) than typical HIMU basalts, whereas the São Jorge and Terceira samples have trace elements patterns similar to HIMU basalts. Pico samples exhibit a signature intermediate between those of typical EMI/EMII and HIMU basalts.

4.2 Sr-Nd-Pb-Os-Hf isotopes Figures 2, 8 and 9 illustrate the Sr, Nd, Pb, Hf and Os isotope systematics of the Central Group island samples analyzed in this study. The Sr-Nd-Pb isotope variations are generally consistent with previous work (White et al., 1979; Davies et al., 1989; Turner et al., 1997; Moreira et al., 1999; Millet et al., 2009), although our data set does not include the lowest 206Pb/204Pb ratios (~19.01) found previously in some Faial samples (Fig. 8). 206Pb/204Pb ratios in the samples from this study range from19.05 to 20.37, with more limited variation in 208Pb/204Pb (38.73-39.40) and especially 207Pb/204Pb (15.53-15.66). Pico, Terceira and São Jorge exhibit large intra-island variations of 206Pb/204Pb (19.48-19.98, 19.49-20.00, and 19.34-20.37 respectively), whereas the range of 206Pb/204Pb in Faial samples is comparatively narrow (19.05-19.44). Fig. 8a, b show that Pico, Terceira and most of the São Jorge samples plot to the right of the NHRL in both 208Pb/204Pb-206Pb/204Pb and 207Pb/204Pb-206Pb/204Pb, but all of Faial samples and two São Jorge samples that have elevated 87Sr/86Sr plot to the left side of NHRL, consistent with the results of Millet et al. (2009). Also in accord with the results of Millet et al. (2009), the São Jorge samples exhibit two sub-parallel trends in 208Pb/204Pb-206Pb/204Pb (Fig. 8b), and for a given 206Pb/204Pb, samples from Topo (oldest formation, designated as “São Jorge-O”) have lower 208Pb/204Pb ratios than samples from Rosais and Manadas (intermediate and youngest formations, designated as “São Jorge-Y”). Compared to the large range in 206Pb/204Pb, the Sr, Nd and Hf isotope variations of the Central Group island samples are relatively limited. Nevertheless, Faial and Pico samples have slightly elevated 87Sr/86Sr (0.7035-0.7040) and lower 143N/144Nd (0.51284-0.51295) and 176Hf/177Hf (0.28292-0.28308) than São Jorge and Terceira samples (0.7033-0.7038, 0.51288-0.51298 and 0.28295-0.28311 respectively, Fig.8), and the Central Group island samples exhibit two trends in Sr-Nd isotope space (Fig. 8c). Samples from São Jorge and Terceira exhibit a slight positive correlation between Sr and Nd isotopes, whereas samples from Faial and Pico exhibit a strong negative correlation. Similar distinctions are also found in other

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isotope systems. Fig. 8e shows that there is a slight positive correlation between 206Pb/204Pb and 143Nd/144Nd in the Terceira and older São Jorge samples, and a strong positive correlation in the Faial samples. Pico samples form a mixing trend from one end of the Terceira trend to one end of the Faial trend. Together, the Central Group island samples produce a linear array in Nd-Hf isotope space that has a steeper slope than the global MORB and OIB array, and trends towards an enriched composition below the mantle array (Fig. 8d). Samples from each island show minor but systematic differences in Hf and Nd isotopic compositions: the São Jorge samples from the older formation (Topo) as well as most Terceira samples have relatively depleted (higher) Hf and Nd isotopic signatures; the São Jorge basalts from the younger formations (Rosais and Manadas) have lower 143Nd/144Nd and 176Hf/177Hf values than Topo samples, but higher than Faial and Pico basalts. The Pico and Faial basalts overlap in composition, but Faial basalts extend to the lowest 143 144 176 177 Nd/ Nd and Hf/ Hf values. Terceira and most São Jorge samples have ΔεHf values close to 0, and plot on the global MORB and OIB array (εHf=1.59εNd+1.28, Chauvel et al., 2008; Fig.

8d), except for two São Jorge samples that have lower ΔεHf (-2.5 and -2.6) similar to Faial and Pico, and that also plot in the range of Faial and Pico samples for Sr, Nd and Pb isotopes. Correlations between 206Pb/204Pb and 176Hf/177Hf ratios for the Central Group samples (Fig. 8f) highlight the distinctions between islands in Hf isotope signatures, with the Pico and younger São Jorge samples (Rosais and Manadas) exhibiting a mixing trend from one end of the Terceira trend to one end of the Faial trend. Osmium concentrations of the Central Group island samples range from 1 ppt to 260 ppt, and display a positive correlation with both MgO and Ni content, consistent with the compatible nature of Os. 187Os/188Os isotope ratios span a range from 0.122 to 0.177 (Table 1). The Os isotope ratios do not appear to vary systematically between islands, but rather are correlated with Os concentration. In particular, at low Os concentrations, there is a large variation of Os isotope ratios from highly radiogenic to sub-chondritic values, but the higher Os concentration samples display relatively limited variation. Almost all of the samples with 187Os/188Os>0.13 (except WAF2 and WAT5) have Os concentrations <10ppt. Osmium isotopic compositions of basalts are very sensitive to shallow-level contamination, which can produce radiogenic 187Os/188Os ratios (Reisberg et al., 1993; Martin et al., 1994). Therefore, in the following discussion, we focus only on the Os isotopic data from samples with >10 ppt Os (Widom et al., 1996).

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Variations of 187Os/188Os with Os concentration and 87Sr/86Sr, 143Nd/144Nd and 206Pb/204Pb isotopes are shown in Fig. 9. Most samples from the Central Group islands have Os isotopic signatures in the range of depleted mantle (Snow and Reisberg, 1995) or primitive mantle (Alard et al., 2005), despite the radiogenic signatures and large variations in Pb isotopes. Only two samples with Os concentrations lower than 25ppt have 187Os/188Os ratios higher than primitive mantle (187Os/188Os>0.129; WAF2 and WAT5 in Fig. 9). No samples from Pico or Faial islands had strongly sub-chondritic 187Os/188Os ratios (<0.122), in contrast to the finding of Schaefer et al. (2002). The samples from Faial exhibit broadly positive and negative correlations, respectively, between 187Os/188Os and Sr and Nd isotopes (Fig. 9c and 9d), but such correlations are not observed in the samples from the other islands.

5. Discussion The Sr-Nd-Pb isotope signatures of the Central Group island basalts are consistent with a mixture of DMM with HIMU/FOZO and EM components in their sources. With the exception of Pb isotopes, the variations of the other isotopes of the Central Group island lavas are relatively limited compared to the global ocean island basalt range including that observed in São Miguel. Nevertheless, the lavas from each of the Central Group islands exhibit distinct isotopic signatures, potentially reflecting distinct and heterogeneous mantle sources (Fig. 2 and Fig. 8).

5.1 Fractional crystallization As described previously, the large variations of major and trace elements in Azores basalts demonstrate an important role for fractional crystallization and potentially variable degrees of partial melting. Before discussing the origin of different mantle components in the sources of the Central Group island basalts, it is important to evaluate the magmatic histories of these samples by further examining their major and trace element variations. The strong positive correlation between Ni and MgO illustrates the role of fractional crystallization in the Central Group island samples (Fig. 3b). Olivine and clinopyroxene phenocrysts, feldspar microphenocrysts and Fe-Ti oxides are found in all of the Central Group island samples, and it is important to evaluate the effects of fractional crystallization on the trace elements. Beier et al. (2006) suggested that relatively evolved lavas with MgO wt.% <5% have generally experienced plagioclase and Fe-Ti oxide fractionation. The lack of significant Eu

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anomalies in the samples from this study (Fig. 5 and Fig. 6) indicate that plagioclase was not a dominant fractionating mineral phase, but the compatible behavior of Fe and Ti in the more evolved samples indicates an important role for fractionation of Fe-Ti oxides in samples with <5 wt.% MgO (Fig. 4). The crystal-rich anakaramites with MgO>12 wt.% likely result from olivine and clinopyroxene accumulation (Beier et al., 2006). In this study, many highly incompatible element ratios, including U/Pb, Th/U, Ba/Th as well as ratios of LREE/HREE (e.g. La/Yb and Ce/Yb) and the ratios of HFSE (e.g. Nb/Zr), exhibit large variations. However, these ratios do not vary systematically with MgO content (Fig. 7), confirming that they were not significantly affected by processes of fractional crystallization or crystal accumulation.

5.2 Partial melting Partial melting also can potentially affect trace element systematics. Based on the strongly LREE-enriched patterns observed in Azores basalts (e.g. Fig. 6), which are controlled by the presence of garnet in the mantle source, basalts from the Central Group island lavas have been suggested to reflect 3%-5% partial melting of mantle peridotite in the presence of 6%-10% residual garnet (Bourdon et al., 2005), and differences in the degree of LREE-enrichment between basalts from São Miguel and Terceira have been interpreted to reflect subtle variations in degree of melting (Beier et al., 2008). Titanium is also a particularly useful tracer for investigating relative degrees of partial melting, because it is a moderately incompatible element that is sensitive to degree of melting but not affected by variable source enrichment (Klein and Langmuir, 1987) or variable pressure (Beier et al., 2008). Although fractionation of Fe-Ti oxides may modify the TiO2 content of a melt, samples with MgO wt.%>5% do not exhibit evidence for

Fe-Ti oxide crystallization. Faial and Pico samples have lower TiO2 than São Jorge and Terceira, which implies a slightly higher degree of melting beneath Faial and Pico (Elliott et al., 2007), possibly consistent with the location of Faial and Pico closer to the mid-Atlantic ridge and in a region of thinner lithosphere. However, a lack of correlation of highly incompatible trace element ratios that are unaffected by garnet stability, such as ratios involving HFSE (e.g. Nb/Zr,

Ba/Nb) with TiO2 (or other major element affected by depth of melting such as Al2O3 or SiO2) for the relatively less evolved samples with MgO wt.%>5% suggests that distinct incompatible trace element signatures in the Central Group island basalts (Fig. 6 and 7) are likely controlled primarily by source heterogeneity rather than variable degrees of melting.

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5.3 Shallow-level contamination In order to link observed trace element and isotopic signatures in basalts to mantle source compositions, it is also critical to assess the potential role of shallow level crustal assimilation. It was demonstrated above that incompatible trace element ratios do not correlate with indices of differentiation such as MgO, which might be expected if crustal assimilation or assimilation-fractional crystallization (AFC) played an important role in the petrogenesis of the Azores basalts. However, more subtle effects of assimilation should be further evaluated. Osmium isotope systematics are highly sensitive to shallow-level crustal contamination, and are therefore ideal for identifying even a minor crustal influence. As discussed previously, samples with low Os abundances (<10 ppt) have radiogenic Os isotope signatures (187Os/188Os > 0.130) that likely do reflect crustal assimilation. Even minor contamination of basaltic magma by oceanic crust or volcanic edifice basalts that have been influenced by seawater-derived Os, and/or assimilation of small amounts of ocean floor sediments with radiogenic 187Os/188Os, can substantially elevate 187Os/188Os in lavas with low Os concentrations (Reisberg et al., 1993; Martin et al., 1994; Widom and Shirey, 1996). The Azores oceanic lithosphere ranges in age to 50Ma (Searle, 1980), but even ~50Ma old oceanic crust would develop highly radiogenic 187Os/188Os ratios of 0.2 to 4.3 given the observed range in MORB 187Re/188Os ranging from 100 to 5000 (Shirey and Walker, 1998). Seawater alteration would not change this estimate substantially, given that seawater ratios through time show a narrower range in 187Os/188Os (Peucker-Ehrenbrink and Ravizza, 2000). Based on similar Os concentrations in the low abundance Azores basalts and typical oceanic crust, as little as 0.7-3% assimilation would significantly increase the 187Os/188Os of the Azores basalts from 0.122 to values greater than the observed range for samples with >10 ppt Os. It is also common for marine sediments to have ~30-50 ppt Os (Hattori et al., 2003; Kendall et al., 2009) and radiogenic 187Os/188Os (~1) derived from seawater (Levasseur et al., 1998), so even less sediment assimilation could be accomodated. For sediments with 50ppt Os and 187Os/188Os=1, 0.5% sediment assimilation would increase the 187Os/188Os ratios of the basalts above observed values, assuming the lavas have 30ppt Os, typical for many of the Azores basalts. Mn-rich sediments are substantially more enriched in Os, with concentrations as high as 1000 ppt (Reisberg et al., 1993), such that as little as 0.2% assimilation by typical Azores basalts would result in 187Os/188Os ratios >0.14. Therefore, even

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minor shallow-level contamination would affect the 187Os/188Os signatures of the basalts. With the exception of two samples (WAF2, WAT5), samples with Os>10 ppt have sub-chondritic 187Os/188Os ratios (Fig. 9), indicating that they escaped even relatively minor (<0.5%) late-stage contamination. The two samples (WAF2, WAT5) with slightly supra-chondritic 187Os/188Os ratios have Os concentrations lower than 30 ppt, and their Os isotopic compositions may therefore have been impacted by minor (<0.2%) marine sediment assimilation. Oxygen isotope signatures in olivines from several samples of Central Group island basalts also place constraints on the role of sediment assimilation; δ18O values ranging from 4.87‰ to 5.14‰ (Turner et al., 2007), all below MORB mantle values of 5.2‰, rule out any significant contamination by marine sediment, which is characterized by extremely high δ18O values (15 to 25‰; Eiler et al., 1997). The lack of Os isotope evidence for even very minor crustal assimilation in most samples suggest that the large ranges in Sr, Nd, Pb and Hf isotopic signatures in the Azores basalts are likely to reflect those of their mantle sources. The isotopic compositions of magmas would not be significantly disturbed even if contaminated by 0.5% marine sediment. For example, 0.5% marine sediment addition would increase the 87Sr/86Sr of a magma from 0.70350 to 0.70355 and decrease 206Pb/204Pb from 19.10 to 19.09 (Hart et al., 1992; Widom et al., 1997), both insignificant changes, assuming a magma with [Sr]=500 ppm, [Pb]=2 ppm (similar to basalts in this study), and marine sediment with [Sr]=600 ppm, 87Sr/86Sr=0.71, [Pb]=25 ppm and 206Pb/204Pb=18.9 (Weaver, 1991). Marine sediment typically has higher concentrations of Sr and Pb, and larger variations of radiogenic isotopes than altered oceanic crust, so contamination by altered oceanic crust should have even less of an impact than marine sediment on the isotopic compositions of the Azores basalts. Similar mass balance arguments can be made for Nd and Hf isotopes, and correlations of Sr and Pb isotopes with Hf and Nd isotopes (Fig. 8) suggest that variations in Nd and Hf isotopes also are independent of crustal assimilation processes. Therefore, the large variations observed in radiogenic isotope ratios in the Central Group island basalts can be attributed to significant mantle source heterogeneity.

5.4 Mantle heterogeneity beneath the Central Group islands Major and trace element systematics of the Central Group island basalts illustrate that they have experienced extensive fractional crystallization and variable degrees of partial melting, but

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insignificant shallow-level crustal assimilation. Isotopic variations and variations in some highly incompatible trace element ratios are independent of these processes, and instead reflect the involvement of distinct mantle source components. Basalts from São Jorge and Terceira have radiogenic 206Pb/204Pb ratios and display a HIMU/FOZO-like signature, whereas Faial basalts have elevated 87Sr/86Sr characteristic of an EMI/EMII signature, and Pico basalts show a mixing trend between these HIMU/FOZO and EM components (Fig. 2).

5.4.1 Source of Terceira and São Jorge basalts The São Jorge and Terceira basalts in this study have 206Pb/204Pb values ranging from 19.4 to 20.4 and 19.5 to 20.0, respectively, with negative Δ7/4Pb (expect for a few that plot in the Faial range, Fig. 8a) and slightly elevated 87Sr/86Sr (Dupré et al., 1982; Turner et al., 1997; Moreira et al., 1999), suggesting a HIMU component in their mantle source. Several previous studies proposed that subducted ancient oceanic crust may be the source of the HIMU mantle component (Zindler et al., 1982; Weaver, 1991; Chauvel et al., 1992; Hofmann, 1997; Stracke et al., 2003), due to the development high U/Pb and Th/Pb ratios during subduction dewatering and loss of soluble Pb, which would develop radiogenic Pb isotope signatures with time. In contrast, Halliday et al. (1995) have distinguished two types of HIMU components based on lead isotope and incompatible trace element systematic of OIB and MORB, neither of which they attribute to recycled oceanic crust. One of these HIMU components is characterized by moderately high 206Pb/204Pb ratios (<20.5) and elevated 87Sr/86Sr but low 207Pb/204Pb signatures (to the right of the NHRL), and is inferred to represent a “young” HIMU reservoir produced within one hundred million years by metasomatism of in situ oceanic lithospheric mantle by small degree basaltic melts. The other HIMU component, referred to as the “old” or “extreme” HIMU reservoir, is characterized by higher 206Pb/204Pb (206Pb/204Pb>20.5) and high 207Pb/204Pb signatures such that it lies along the NHRL, and is inferred to have been produced by old metasomatized oceanic lithospheric mantle that has been recycled back into the mantle (Halliday et al., 1995). The São Jorge and Terceira basalts are characterized by Pb isotope signatures that fall primarily to the right of the NHRL (Fig. 8), and in this sense are similar to the young HIMU identified by Halliday et al. (1995). However, such signatures could potentially be produced by any process that results in an enrichment in U relative to Pb in the source, including introduction of relatively young recycled oceanic crust or the involvement of metasomatized lithospheric

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mantle, provided it is recent enough to produce significant radiogenic 206Pb but not 207Pb ingrowth. Consideration of elemental systematics, as well as Os and Hf isotopes can potentially help to distinguish between these processes. If the São Jorge and Terceira lavas were produced from a source containing significant subducted oceanic crust, they should exhibit relative depletions in K and other large ion lithophile elements (LILE: e.g. Cs, Rb, Ba and Sr), because these elements are highly fluid-mobile and would be released from the recycled oceanic crust during subduction dehydration (Weaver, 1991). Conversely, Nb and Ta (and other HFSE) are highly insoluble during subduction dehydration due to their high charge and small ionic radius (Weaver, 1991), and subducted oceanic crust would largely retain Nb and Ta during subduction dehydration. Fig. 6 shows that the São Jorge and Terceira basalts have trace element patterns similar to the HIMU end-member. São Jorge and Terceira basalts do have slightly lower K (and Rb) contents than basalts from Faial and Pico for a given MgO concentration (Fig. 4) and Nb concentration (Fig. 7d), and most of them have lower K/Nb (~160-240) and Rb/Nb (~0.35-0.60) than E-MORB (~253 and 0.61 respectively, Sun and McDonough, 1989). These elemental features are potentially consistent with a component of recycled oceanic crust in the mantle source of São Jorge and Terceira basalts, but this hypothesis can be further evaluated using Os and Hf isotope data.

5.4.1.1 Osmium isotope constraints on crustal recycling The Os isotopic system is a complementary tracer to Sr and Pb isotopes for assessing potential recycled crustal material in a mantle source. Oceanic crust has much higher Re/Os than mantle peridotite, because Os is compatible and Re moderately incompatible during mantle melting. Oceanic crust will therefore develop an extremely radiogenic Os isotope signature with time, and if recycled into the mantle in sufficient quantity, could produce a hybrid mantle source with supra-chondritic 187Os/188Os. HIMU OIB from the South Pacific, including those from Rurutu, Tubuai and Mangaia (Hauri and Hart, 1993) and the Cook-Austral islands (Schiano et al., 2001; Lassiter et al., 2003) have been found to have strongly supra-chondritic 187Os/188Os (0.128-0.159), consistent with their mantle source containing fragments of recycled oceanic crust. However, São Jorge and Terceira basalts differ distinctly from these extreme HIMU lavas in that, with the exception of the samples with Os abundance <10 ppt, most samples have sub-chondritic

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rather than supra-chondritic 187Os/188Os ratios. In addition, if there were a significant component of ancient recycled oceanic crust in the mantle source of the São Jorge and Terceira lavas, co-variations between Os isotopes and other radiogenic isotopes (especially Pb) would be expected. Such co-variations have been found in some extreme HIMU lavas, including OIB from the South Pacific (Hauri and Hart, 1993; Lassiter et al., 2003). However, the São Jorge and Terceira lavas do not exhibit correlations between Os isotopes and other isotope systems (Fig. 9), which may allow constraints to be placed on the maximum potential contribution of recycled oceanic crust to their mantle sources. Given the low Os concentrations in oceanic crust relative to mantle peridotite, it has been proposed that less than 10% of recycled oceanic crust mixed into the mantle would not significantly disturb the Os isotope signature of the hybrid source (e.g. Widom and Shirey, 1996). In this study, we further model the expected mixing trends between a hypothetical mantle plume component similar to the enriched mantle plume of Hart et al. (1992) and Shirey and Walker (1998), which we assume as an approximation of the predominant (or “common”; Millet et al., 2009) mantle component in the Azores, to which variable recycled oceanic crust may be incorporated either with or without sediments. Based on the present-day parent:daughter ratios and radiogenic isotopic compositions of oceanic crust, we can calculate their isotopic ratios at any time in the past. Using current estimates of parent and daughter element mobility during subduction dewatering, we can then calculate the modified parent:daughter ratios and the expected present-day radiogenic isotope ratios that such subducted components would develop. The isotopic compositions of the “common” Azores mantle component, recycled oceanic crust, recycled terrigenous sediment, pelagic sediment and GLOSS are calculated by this method for 0.5Ga and 2Ga recycled materials, assuming element mobilities during subduction dewatering as estimated by Stracke et al. (2003). The compositions of these potential end-members are listed in Table 2. As discussed previously, the fluid mobilities of Rb-Sr and U-Pb are high, leading to uncertainties in the expected trace element and isotopic compositions of recycled crustal materials. The situation for the Re-Os isotope system is somewhat mitigated in that Os is generally considered to be immobile in hydrous slab-derived fluids based on experimental data (Xiong and Wood, 1999; Xiong and Wood, 2000) as well as evidence from comparisons of metabasalts and metagabbros with their likely oceanic crustal protoliths (Becker, 2000; Dale et

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al., 2009). The mobility of Re in slab fluids may be more significant than Os, with up to 40-60% of the Re released during subduction dewatering in some cases (Becker, 2000; Dale et al., 2009). However, evidence from mantle xenoliths from active subduction zones indicates inconsequential loss of Re during subduction dewatering (Brandon et al., 1996; Widom et al., 2003). Regardless, the mobilities of both Re and Os are apparently much less significant than that of alkalis, alkaline earths and Pb, thus minimizing the relative uncertainties in calculated Os isotopic compositions of recycled crustal components compared to Sr and Pb. Constraining the Os isotopic composition of ancient recycled oceanic crust is somewhat complicated by the fact that there is a tremendous range in Re/Os ratios in modern MORB, which regardless of potential minor modification of Re/Os ratios during subduction dewatering, could produce an extreme range of present-day Os isotope signatures. Nevertheless, we can calculate a range of potential compositions for oceanic crust recycled at various times in the past, and the effect of mixing such crust into a mantle source, based on the observed ranges in 187Re/188Os (325 to 1000) and Os concentrations (2 to 30 ppt) in typical modern MORB (e.g. Becker, 2000). Of note is the fact that Re/Os ratios in MORB are negatively correlated with Os concentration (Schiano et al., 1997), such that recycling of these diverse crustal compositions into the mantle does not produce drastically different signatures for a given age of the recycled crust (e.g. Widom and Shirey, 1996; Widom, 1997). The results of the mixing calculations demonstrate that substantial fractions of recycled oceanic crust can be mixed into the mantle without significantly altering the Os isotopic composition. Fig. 10a and b show calculated Pb-Os and Sr-Os isotope mixing trends between the “common” Azores mantle component and recycled oceanic crust representing the range in Re/Os ratios and Os concentrations of typical MORB (shown as OC-1 and OC-2) and different recycling ages (0.5Ga and 1Ga). Despite the very radiogenic Os isotope signatures of the recycled crust, up to 30% of 0.5Ga and 10% of 1Ga recycled oceanic crust could be incorporated, respectively, without significantly altering the Os isotopic composition of the mantle source. However, if relatively young subducted oceanic crust (0.5Ga) were mixed into the Azores mantle, incorporation of >50% of such crust would be required to produce the most radiogenic 206Pb/204Pb ratios found amongst the São Jorge and Terceira basalts (~20.4), which would produce substantially elevated Os isotopic signatures. Furthermore, even such a large fraction of recycled crust would not produce the most radiogenic 87Sr/86Sr amongst the São Jorge and

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Terceira basalts (0.704). If older recycled oceanic crust (1Ga) were mixed into the Azores mantle plume, ~15% recycled crust could produce basalts with 206Pb/204Pb and 87Sr/86Sr isotopic signatures in the range of the most radiogenic São Jorge and Terceira basalts. Recycled oceanic crust with moderate Os concentration and low 187Re/188Os (OC-2) would impart slightly elevated 187Os/188Os signatures to the source, but recycled oceanic crust with lower Os concentration and high 187Re/188Os (OC-1) would not disturb the Os isotope system significantly. In order to make the figures clear, mixing trends between the Azores mantle plume and 2Ga recycled oceanic crust are not shown. However, calculations demonstrate that mixing 4% of 2 Ga recycled oceanic crust with the OC-1 composition can produce the Sr and Pb isotope signatures of the Central Group island basalts. Therefore, mixing with ~15% 1Ga or 4% 2Ga recycled oceanic crust with relatively low Os concentration and high Re/Os (OC-1) into the enriched plume of Azores could potentially produce Sr, Pb and Os isotope signatures consistent with the enriched component in the mantle source of São Jorge and Terceira basalts.

5.4.1.2 Hafnium isotope constraints on crustal recycling Hafnium isotopes can also provide relatively robust and complementary evidence to constrain the potential role of crustal recycling in the HIMU-type sources of the São Jorge and Terceira basalts, given the limited effects of subduction dewatering on Lu and Hf and the added sensitivity to potential sediment contributions. Previous studies have demonstrated that ancient oceanic crust would develop low 176Hf/177Hf values for a given 143Nd/144Nd value and fall significantly below the global MORB-OIB trend (εHf=1.59εNd+1.28, Fig. 11a; Chauvel et al., 2008). Hafnium isotope signatures of some extreme HIMU OIB (Cook-Austral islands) (e.g. Ballentine et al., 1997, Salters and White, 1998, Lassiter et al., 2003) do in fact have negative

ΔεHf values of ~-2 (ΔεHf defined as the difference between measured εHf values and those predicted by the global Hf-Nd correlation, Johnson and Beard, 1993), consistent with a model in which their mantle source contains a component of ancient recycled oceanic crust. Most of the

São Jorge and Terceira lavas plot on the global MORB-OIB array (ΔεHf ≈0, Fig. 8d), which is not consistent with their mantle source containing significant ancient recycled oceanic crust. As discussed in the previous models, ~15% of 1Ga or ~4% 2Ga recycled oceanic crust mixed into the Azores mantle could produce the radiogenic Pb and Sr isotope signatures of São Jorge and Terceira. However, such a mixture would also affect the Hf isotope system. Fig. 11b shows the

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expected mixing trends between Azores mantle with 1Ga and 2Ga recycled oceanic crust, assuming average compositions for recycled oceanic crust and mantle, and minor modifications to the Lu/Hf ratios during subduction dewatering based on data of Stracke et al. (2005) (Table 3). This model illustrates that recycled oceanic crust, regardless of the age or amount incorporated into the mantle, cannot produce the Hf-Nd variations observed in the São Jorge and Terceira basalts. The recycled oceanic crust would produce a steeper Hf-Nd trend resulting in lower εHf at a given εNd. The combined results from Os and Hf isotopes therefore suggest that recycled oceanic crust is unlikely to be the source of the enriched component in the mantle beneath São Jorge and Terceira. An alternative possibility is that the mantle source of the São Jorge and Terceira basalts contains a mixture of recycled oceanic crust coupled with pelagic and/or terrigenous sediment. Chauvel et al. (2008) proposed that the global MORB and OIB Hf-Nd array is produced by mixing depleted mantle peridotite and 20-30% of a recycled oceanic crustal package containing 0-6% sediment. Fig. 11c shows calculated mixing trends between the predominant Azores mantle plume component and 1Ga recycled oceanic crust with 3% of various sediment compositions, including terrigenous sediment (TS), two types of pelagic sediment (PS-1 and PS-2) and global subducting sediment (GLOSS). The sediment compositions (Table 3) were chosen to represent terrigenous sediment with extremely low 176Hf/177Hf, pelagic sediments with low 176Lu/177Hf and either extremely low 176Hf/177Hf (PS-1) or high 176Hf/177Hf (PS-2), and the average composition GLOSS, in order to evaluate all of the potential isotopic variations resulting from mixtures of oceanic crust and sediments. Only when the mantle plume contains a large amount (>50%) of the mixture of 1Ga recycled oceanic crust plus 3% pelagic sediment type 2, does the mixing trend approach the Nd-Hf trend of the São Jorge and Terceira basalts (Fig. 11c). However, such a large percentage of recycled oceanic crust is inconsistent with the sub-chondritic 187Os/188Os ratios, which limits the recycled component to 15%. Note that decreasing the percentage of sediment in the subducted crustal package would result in mixing arrays that are too steep, and increasing the percentage of sediment (for example, to 6%) still requires >50% recycled crustal material to reproduce the observed trends of the São Jorge and Terceira basalts. Based on the above models, we can thus exclude the possibility of a mixture of oceanic crust and sediment contained in the mantle source of São Jorge and Terceira basalts.

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One additional crustal recycling model that can be tested is that for mixing pure sediment into the Azores mantle source, decoupled from subducted oceanic crust. Fig. 10g and 10h show calculated mixing trends between Azores mantle plume and terrigenous and pelagic sediment with ages of 1Ga. The mixing trends between mantle plume and terrigenous and pelagic sediment with ages of 0.5Ga and 2Ga (not shown) are similar to the mixing trends shown here. Mixing <1% sediment would produce the radiogenic 87Sr/86Sr in the basalts without disturbing 187Os/188Os (Fig. 10h). However, more than 20% ancient (2Ga) terrigenous sediment would be required to produce the most radiogenic 206Pb/204Pb (Table 1; not shown in the figures) found in the São Jorge and Terceira basalts, which would produce radiogenic 187Os/188Os that is inconsistent with the observed sub-chondritic 187Os/188Os (Fig. 10g), as well as distinctly heavy oxygen isotope signatures that are inconsistent with existing data from Turner et al. (2007).

5.4.1.3 Alternative models for the origin of source of Terceira and São Jorge basalts Metasomatized oceanic lithosphere is another possible young HIMU reservoir (Halliday et al., 1995). Based on the measured U/Pb ratios of the São Jorge and Terceira samples (0.35 to 0.65), which reflect the maximum possible U/Pb ratio in their mantle source, 500Ma to 1Ga would be required to produce the range in 206Pb/204Pb from 19.4 to 20.2, assuming the original 206Pb/204Pb ratio of the depleted mantle source was 18.8 (Elliott et al., 2007). However, the Azores oceanic lithosphere is relatively young (<50Ma, Searle, 1980), and therefore could not be the source of the more radiogenic 206Pb/204Pb ratios observed in São Jorge and Terceira lavas. The high 3He/4He signature of Terceira basalts (Moreira et al., 1999; Jean-Baptiste et al., 2009) suggests that Terceira is fed by a mantle plume containing a relatively undegassed reservoir from deep mantle. Elevated 20Ne/22Ne ratios in Terceira basalts also support the suggestion that the mantle plume contains material from the deep mantle (Madureira et al., 2005), similar to the proposed FOZO end-member. Although no He or Ne isotope measurements have been performed for São Jorge lavas, similarity of the Sr and Pb isotopic data of the Rosais and Manadas lavas with Terceira lavas suggest that they share the same FOZO mantle plume source. Compared to extreme HIMU OIB which require a source enriched in U and Th relative to Pb without elevated Rb/Sr, FOZO has lower 206, 207, 208Pb/204Pb, higher 208Pb/206Pb and 87Sr/86Sr (Stracke et al., 2005). In Fig. 2, almost all of the Rosais and Manadas lavas (intermediate and youngest formations) from São Jorge are in the range of FOZO. Terceira lavas also generally fall

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within the FOZO field, but show a slightly positive trend between 206Pb/204Pb and 87Sr/86Sr, with one end-member in the range of Azores platform tholeiites (APT) and the more radiogenic 206Pb/204Pb samples slightly elevated in 87Sr/86Sr compared to Rosais and Manadas lavas. The 176Hf/177Hf ratios of Terceira, Rosais and Manadas lavas are also in the range of FOZO (0.28295 to 0.28310) and extend to values higher than those of HIMU (0.2828 to 0.2830, Stracke et al., 2005). These data thus further suggest that the mantle sources of the Rosais and Manadas (São Jorge) and Terceira lavas are distinct from typical end-member HIMU mantle, and are instead consistent with a model in which they were produced by mixing of an enriched, deeply derived heterogeneous mantle plume (FOZO) with relatively depleted mantle characteristic of that beneath the Mid-Atlantic ridge. In contrast, the Topo lavas have higher 87Sr/86Sr at a given 206Pb/204Pb than FOZO and fall outside of the mixing trend between depleted Mid-Atlantic ridge mantle and FOZO, suggesting that the Topo mantle source may involve an additional enriched component. Some Topo basalts appear to tap a source similar to the Faial basalts, with high 87Sr/86Sr. However, most of the Topo lavas exhibit negative ∆7/4Pb and ∆8/4Pb, in contrast to the positive values observed in Faial (Fig. 8a and 8b). The Topo lavas are also inconsistent with a HIMU source due to their relatively elevated 87Sr/86Sr signatures. Although the origin of Topo mantle source remains unresolved, the results of this study are consistent with recent studies by Millet et al. (2009) and Pinto Ribeira (2011), which all reveal this distinct Topo component in the mantle source of the Central Group island basalts that was apparently tapped only in the earliest stages of the evolution of São Jorge.

5.4.2 Nature of the Azores Central Group EM component Faial basalts have higher 87Sr/86Sr and lower 206Pb/204Pb, 143Nd/144Nd and 176Hf/177Hf than the Terceira and São Jorge lavas, and most samples plot to the left of the NHRL (Fig. 2, 8). Faial 143 144 206 204 87 86 samples also have negative ΔεHf values down to -2. The Nd/ Nd, Pb/ Pb and Sr/ Sr ratios of the Faial samples reflect an EM-type signature, but do not extend to either extreme EMI or EMII compositions (Fig. 2, 8). Possible origins for EM-type components include pelagic and/or terrigenous sediments subducted with recycled oceanic crust, (Zindler and Hart, 1986 and reference therein, Weaver, 1991; Woodhead and Devey, 1993), recycled sub-continental or sub-oceanic lithospheric mantle (McKenzie and O’Nions, 1983; Sun and McDonough, 1989;

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Schaefer et al., 2002; Turner et al., 2007); or recycled mantle wedge (Eiler et al., 1997; Lassiter et al., 2003). Pelagic sediments have high Pb concentrations with low U/Pb ratios, and high 87Sr/86Sr with high Rb/Sr ratios (Ben Othman et al., 1989). Mixtures of recycled oceanic crust and pelagic sediments can produce mantle sources with low 206Pb/204Pb ratios and high 87Sr/86Sr ratios. Also, ancient oceanic crust and pelagic sediment both have high Re/Os ratios, which, if recycled in sufficient quantity, would produce mantle sources with radiogenic 187Os/188Os ratios. Although the Faial lavas do not have supra-chondritic 187Os/188Os ratios such as observed in Hawaiian lavas with EM signatures, they do exhibit a positive trend between 87Sr/86Sr and 187Os/188Os. Figs. 10c-f show mixing trends between the Azores mantle plume source and 0.5Ga and 1Ga oceanic crust with either pelagic or terrigenous sediment in a ratio of 97:3. The Os isotope signature of these mixtures are mainly controlled by aging, thus different ratios of oceanic crust to pelagic/terrigenous sediment (from 99:1 to 95:5) do not change the Os isotopic composition significantly. If the mantle source of the Faial basalts contained a mixture of 1Ga recycled oceanic crust coupled with sediment, ~7% or less of this mixture could potentially produce Sr, Pb and Os isotope signatures consistent with the enriched component in the mantle source of Faial basalts, including basalts with 87Sr/86Sr ranging from 0.7036 to 0.7041, and a slightly positive trend between Sr and Os isotopes. However, such a model is not supported by the Nd-Hf isotope systematics (Fig. 11c), as it fails to produce εNd signatures as low as the most enriched Faial basalts. However, because Faial basalts do not show any evidence for a FOZO component, and do not have 3He/4He values in the MORB-mantle range (Moreira et al., 1999; Jean-Baptiste et al., 2009), it is possible that their mantle source involves mixing of crustal components into a depleted (DMM-like) rather than an enriched plume source. Nevertheless, Fig. 11c illustrates that mixtures of DMM-like mantle with 1Ga recycled oceanic crust, with or without sediment, also fails to reproduce the Nd-Hf signatures of the more enriched Faial basalts. Continental or oceanic lithospheric mantle also are possible enriched components in the source of EM-type lavas (Milner and LeRoex, 1996; Schiano et al., 2001; Schaefer et al., 2002; Niu and O’Hara, 2003; Turner et al., 2007). Ancient melt depleted sub-continental lithospheric mantle is known to have distinctly sub-chondritic Os isotope signatures due to long-term Re depletion (e.g. Walker et al., 1989; Pearson et al., 1995; Shirey and Walker, 1998). Ancient oceanic lithospheric mantle also would be expected to exhibit sub-chondritic Os isotope

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signatures. In contrast to the results of Schaefer et al. (2002), all of the Faial and Pico samples in this study with Os abundances >25 ppt have 187Os/188Os ratios with the range of depleted MORB mantle and primitive upper mantle range (0.122-0.130). The range of 187Os/188Os ratios of Faial and Pico samples (>25 ppt) is 0.125-0.129 in this study, and 0.125-0.127 in Widom and Shirey (1996). None of these samples were found to have the strongly sub-chondritic Os isotope signatures reported by Schaefer et al. (2002). The Faial sample with strongly sub-chondritic 187Os/188Os ratio found by Schaefer et al. (2002) has a relatively low Os concentration (16.4 ppt) that makes the isotope ratio sensitive to blank, spike, and interference corrections, and until reproduced, there is not strong evidence for an ancient lithospheric mantle component in the mantle source of Faial and Pico. On the other hand, mixing with relatively young or relatively undepleted sub-continental or sub-oceanic lithoshperic mantle could produce the DMM-like Os isotopic signatures of Faial lavas. Previous studies have proposed that either sub-continental or sub-oceanic lithospheric mantle may be present in the upper mantle sources of some OIB, including the Azores (Widom and Shirey, 1996; Schaefer et al., 2002; Turner et al., 2007) and Kerguelen (Hassler and Shimizu, 1998; Ingle et al., 2002), and possibly the Cook-Australs (Schiano et al., 2001) and Iceland (Hanan and Schilling, 1997; DeBaille et al., 2009). If similar metasomatic processes operated on sub-oceanic and sub-continental lithosphere, they would be difficult to distinguish. Both sub-continental and sub-oceanic lithosphere mantle have sub-chondritic 187Os/188Os ratios (e.g. Walker et al., 1989; Pearson et al., 1995; Shirey and Walker, 1998), high 87Sr/86Sr and variable 206Pb/204Pb ratios (Hawkesworth et al., 1990; Lee et al., 1994), which would span the isotopic variation of Faial basalts. The EM type basalts of Iceland have similar sub-chondritic 187Os/188Os ratios to the Faial and Pico basalts, and are proposed to have a component of sub-continental lithosphere in their mantle source (DeBaille et al., 2009). However, mixing with sub-continental and sub-oceanic lithosphere would produce a negative correlation of Sr-Os isotopes as found in the Iceland basalts, instead of the positive correlation found in the Faial basalts, arguing against sub-continental or sub-oceanic lithosphere in the mantle source beneath Faial. Another possible source of EM-type basalts is subduction-modified sub-arc mantle wedge that has been recycled back into the mantle via viscous coupling to the subducted slab (Eiler et al., 1997; Lassiter et al., 2003; Donnelly et al., 2004). As a result of seafloor alteration and aging, subducted oceanic crust would have radiogenic 87Sr/86Sr. If metasomatized by slab-derived fluid

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and/or melt, mantle wedge would have enriched trace element and radiogenic 87Sr/86Sr signatures. Previous studies have shown that some arc mantle xenoliths have super-chondritic 187Os/188Os ratios (~0.13-0.14) due to interaction with radiogenic slab-derived melts or fluids (e.g. Brandon et al., 1996; McInnes et al., 1999; Widom et al., 1999). During subduction dehydration, up to ~50% of the Re may be transferred from the subducted slab to the overlying mantle wedge (Becker, 2000), which may also contribute to the development of super-chondritic 187Os/188Os ratios over time. Although none of the Faial basalts exhibit super-chondritic Os isotope signatures, mixing between depleted mantle and metasomatized mantle wedge, characterized by higher 87Sr/86Sr and super-chondritic 187Os/188Os, could produce the positive Sr-Os isotopic trend observed in the Faial basalts (Fig. 9c). In addition, oceanic sediments have a large range of εHf from ~-50 to 15 and εNd from ~-25 to 10 (Chauvel et al., 2008). Although Hf is thought to be immobile in a strictly hydrous slab fluid, even very minor transfer of terrigenous sediment-derived Hf to the mantle wedge could produce a distinctly low εHf and negative ΔεHf signature in the wedge. Alternatively, a minor component of sediment melt, in which Hf would be moderately mobile (up to 45% loss; Stracke et al., 2003), could also overprint the mantle wedge signature as evidenced by the large variation of εHf (~-10 to 20) and εNd (~-8 to 12) in arc lavas (e.g. White and Patchett, 1984; Pearce et al., 1999; Chauvel et al., 2008; GEOROC database). Trace element data can provide additional insights into potential processes involved in the development of the Faial enriched mantle source. Metasomatism of mantle wedge due to subduction dewatering of oceanic crust and sediment might be expected to produce elevated ratios of LILE and LREE relative to HFSE due to the high and moderate solubilities of the former, and insolubility of the latter in hydrous slab-derived fluids. Therefore, mantle wedge that has been metasomatized by slab-derived fluids might be predicted to have higher LILE/HFSE and LREE/HFSE ratios than pre-subduction MORB-mantle (Weaver, 1991). In contrast, Faial basalts have lower La/Nb ratios (0.7-0.9) than N-MORB (1.07) and arc lavas (Weaver, 1991). However, previous studies have shown that orthopyroxene, clinopyroxene, garnet and olivine have Nb crystal/liquid distribution coefficients larger than those of La (Kelemen et al., 1993), indicating that reactions between subduction generated melt and peridotite mantle wedge can produce arc lavas with depleted Nb relative to La, leaving a residual mantle wedge enriched in Nb. If such mantle wedge were recycled in the source region of Faial, it could produce basalts

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with enriched Nb. Elliott et al. (2006) proposed another possibility, which is that the formation of hydrous minerals might change the behavior of some highly incompatible elements. For example, amphibole in the Zagarbad peridotite in the Red Sea, which has been argued to have a subduction signature (Brooker et al., 2004; Elliott et al., 2006), is a significant host for Nb, raising the Nb concentration and causing low La/Nb ratios in the metasomatized mantle wedge. amp/liq During mantle melting, Rb is incompatible in amphibole with DRb =0.023, but K is amp/liq compatible with DK =9.6 (Halliday et al., 1995). Therefore, if amphibole is residual in the mantle source of the Central Group island basalts during partial melting, magmas with low

K2O/Rb ratios would be produced. However, the relatively high K2O contents at a given MgO for Faial and Pico basalts compared to those of Terceira and São Jorge (Fig. 4) could reflect stabilization of amphibole during slab fluid metasomatism of mantle wedge followed by wholesale melting of the amphibole component during OIB formation. This would also explain the relatively lower La/Nb ratios, as well as higher Ba concentrations in Faial basalts. Pico basalts have Sr-Nd-Pb-Hf isotopic signatures that produce mixing trends between Faial and Terceira samples (Fig. 2 and Fig. 8), although their Pb isotopic compositions (206Pb/204Pb, 207Pb/204Pb and 208Pb/204Pb) are mainly in the range of the less radiogenic Terceira samples and they have negative Δ7/4Pb values. The Pico basalts therefore may be produced by mixing of enriched components in the mantle sources of both Faial and Terceira. One end-member of the Pico samples shares the high 87Sr/86Sr ratios, low 206Pb/204Pb, 143Nd/144Nd and 176Hf/177Hf ratios of the enriched Faial samples. As discussed previously, this could represent recycled enriched mantle wedge in the mantle source of Faial. The other end-member of the Pico samples has lower 87Sr/86Sr ratios and higher 206Pb/204Pb, 143Nd/144Nd and 176Hf/177Hf ratios, consistent with the young HIMU or FOZO component identified in the mantle source of São Jorge and Terceira. 3 4 Previous studies have shown that the He/ He ratios of Pico samples are similar (8.5±0.5Ra,

Moreira et al., 1999) or slightly higher (9.4±0.9Ra, Jean-Baptiste et al., 2009) than MORB values

(8±1Ra), but much lower than Terceira basalts (up to 13.5±0.7Ra, Jean-Baptiste et al., 2009). If the elevated 3He/4He ratios of Terceira samples reflect the composition of a deep, relatively undegassed mantle plume, then the mantle plume does not contribute significantly to the source of Pico lavas, but rather may be mixed with a substantial component of the enriched mantle wedge that contributes to the Faial source.

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5.4.3 Model of possible mantle dynamics beneath the Azores Fig. 12 shows a cartoon model of the proposed mantle dynamics beneath the Central Group islands of Azores based on results from this study. The high 3He/4He (Moreira et al., 1999; Jean-Baptiste et al., 2009) and elevated 20Ne/22Ne (Madureira et al., 2005) combined with the seismic evidence (Gente et al., 2003; Yang et al., 2006) indicate that the Azores mantle plume comes from the lower mantle with a relatively primitive or undegassed composition, which is reflected by the FOZO signatures of Terceira and some São Jorge basalts. Depleted mantle-range 3He/4He ratios in Faial lavas indicate that the Faial mantle source has no significant contribution from the relatively undegassed FOZO plume material. The most likely explanation for this is that the Azores mantle plume serves only as a heat source in this region of the archipelago, leading to partial melting of the upper mantle beneath Faial (the red region surrounding the plume in Fig. 12), which is heterogeneous and contains recycled mantle wedge that produces magmas with EM-type signatures. Subduction metasomatized mantle wedge, viscously coupled and dragged to depth by a subducting slab, is expected based on relative density considerations to become decoupled from the downgoing slab at a depth of ~60-80 km (e.g. Furukawa, 1993; Abers et al., 2006; Wada et al., 2008), and thus remaining in the shallow upper mantle. Such hydrous metasomatized mantle is envisioned as the source of the Faial basalts. Pico basalts are produced largely by the same source as Faial basalts, but with a minor plume component, producing a mixing trend between EM and FOZO type signatures.

6 Conclusions As has previously been observed, basalts from the Central Group islands of Azores archipelago exhibit variable trace element and isotopic signatures that indicate derivation from distinct and heterogeneous mantle sources. Despite the large variations in geochemistry, almost all of the samples from the Central Group islands have sub-chondritic 187Os/188Os ratios, which suggest that the enriched components in their mantle sources do not contain significant crustal material. Modeling of the relationships between Os-Sr-Pb and Hf-Nd isotopes further argue against a significant role for recycled oceanic crust ± pelagic/terrigenous sediment in the mantle sources of the Azores Central Group islands. 206 204 87 86 Large variation of Pb/ Pb, slightly elevated Sr/ Sr, ΔεHf≈0 and sub-chondritic 187Os/188Os signatures in São Jorge and Terceira basalts are instead interpreted to reflect the

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FOZO signature characteristic of a deep-rooted mantle plume. Faial basalts have slightly 206 204 87 86 elevated Pb/ Pb, high Sr/ Sr (up to 0.704), negative ΔεHf (-1 ~ -2), sub-chondritic 187Os/188Os and exhibit a positive correlation between 87Sr/86Sr and 187Os/188Os, interpreted as a component of recycled metasomatized mantle wedge with an EM signature in the Faial mantle source. Stabilization and subsequent melting of an amphibole-bearing metasomatized mantle wedge can explain the relatively K2O-rich and low La/Nb signatures of the Faial basalts relative to those of Terceira and São Jorge. Pico basalts are likely produced by mixing between the EM component beneath Faial and the FOZO plume that sources the São Jorge and Terceira basalts.

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Table 1. Major and trace element and isotopic compositions of the Azores Central Group island basalts WAF1a WAF 2 WAF 3 WAF 7 WAF 10 WAF10r WAF 11 WAF 12 SiO2 46.38 48.78 46.39 48.00 47.27 47.56 48.50 TiO2 2.11 2.80 3.14 2.76 2.54 2.82 2.82 Al2O3 12.32 16.01 14.30 15.48 15.35 16.17 16.88 Total Fe 10.85 10.07 11.38 10.18 9.91 10.17 9.64 MnO 0.15 0.16 0.18 0.16 0.16 0.16 0.16 MgO 13.51 7.30 8.94 8.24 9.27 7.84 6.51 CaO 10.90 9.38 11.02 9.91 10.98 10.15 9.23 Na2O 2.54 3.45 2.96 3.41 3.02 3.38 3.97 K2O 0.89 1.46 1.14 1.30 1.07 1.26 1.79 P2O5 0.34 0.58 0.56 0.55 0.43 0.49 0.50 LOI -0.42 -0.21 -0.54 -0.27 -0.40 -0.58 -0.49 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor* 0.991 1.002 1.005 1.003 1.004 1.006 1.005 Li 4.8 6.6 5.6 6.2 5.3 6.0 7.0 Be 1.2 1.7 1.3 1.6 1.2 1.4 1.9 Sc 33 24 31 25 29 27 25 21 V 210 272 332 284 272 259 271 255 Cr 700 159 310 217 433 426 152 168 Co 59 40 49 43 45 47 42 35 Ni 339 111 143 120 171 201 132 82 Cu 63 38 53 40 46 45 44 33 Zn 81 101 107 103 82 89 91 93 Ga 15 21 20 21 19 18 21 22 Rb 18 32 24 27 22 22 26 37 Sr 410 572 506 595 528 551 591 696 Y 20 27 27 26 24 24 26 27 Zr 143 270 214 238 178 180 212 275 Nb 26 43 38 42 33 37 40 51 Cs 0.18 0.29 0.25 0.27 0.24 0.20 0.24 0.30 Ba 227 397 311 332 291 286 322 482 La 21 36 28 33 25 25 28 40 Ce 47 76 61 70 53 52 61 82 Pr 5.7 9.3 7.8 8.6 6.8 7.0 7.8 9.9 Nd 24 37 33 35 28 30 32 39 Sm 5.6 7.8 7.4 7.6 6.3 6.1 7.2 7.9 Eu 1.7 2.4 2.3 2.4 2.0 2.0 2.3 2.4 Gd 4.9 6.8 6.6 6.7 6.4 6.0 7.1 7.7 Tb 0.77 1.04 1.03 1.01 0.87 0.90 0.95 1.01 Dy 4.3 5.6 5.6 5.4 4.8 4.8 5.2 5.3 Ho 0.84 1.07 1.07 1.02 0.92 0.90 1.00 1.03 Er 2.1 2.8 2.8 2.6 2.4 2.2 2.6 2.7 Tm 0.28 0.37 0.37 0.34 0.32 0.32 0.35 0.36 Yb 1.7 2.3 2.2 2.1 1.9 1.9 2.1 2.3 Lu 0.24 0.34 0.33 0.31 0.28 0.29 0.30 0.32 Hf 3.8 5.8 4.8 5.2 4.3 4.4 5.1 6.3 Ta 1.8 3.2 2.8 3.1 2.3 2.4 2.8 3.5 Pb 1.72 1.72 2.03 Th 2.1 3.8 2.6 3.2 2.5 2.5 2.8 4.1 U 0.63 1.14 0.73 0.96 0.76 0.72 0.85 1.14 87Sr/ 86Sr 0.70388 0.70380 0.70373 0.70384 0.70378 0.70380 0.70373 0.70393 143Nd/144Nd 0.512860 0.512875 0.512888 0.512866 0.512876 0.512888 0.512879 0.512842 εNd 4.3 4.6 4.9 4.5 4.6 4.9 4.7 4.0 206Pb/204Pb 19.579 19.330 19.407 19.523 19.387 19.403 19.443 19.320 207Pb/204Pb 15.628 15.637 15.611 15.636 15.616 15.635 15.623 15.634 208Pb/204Pb 39.262 39.159 39.027 39.206 39.120 39.184 39.126 39.058 176Hf/177Hf 0.28296 0.28300 0.28303 0.28296 0.28296 0.28299 0.28299 0.28295 εHf 6.6 8.0 9.0 6.8 6.8 7.8 7.7 6.3 187Os/188Os 0.1269 0.1313 0.1555 0.1296 0.1236 0.1252 0.1352 Os ppt 124.9 21.2 4.0 4.7 17.8 21.6 1.9 48

WAF 14 WAF 15 WAF15r WAF 24 WAF 25 WAF 29 WAF 30 WAF31 SiO2 47.93 47.99 48.58 48.50 47.21 46.65 47.55 TiO2 2.75 2.59 2.66 2.86 2.57 2.86 2.79 Al2O3 16.32 15.31 15.95 16.83 14.45 15.55 15.76 Total Fe 9.85 9.74 9.74 9.60 11.17 11.30 10.70 MnO 0.16 0.16 0.16 0.16 0.17 0.17 0.16 MgO 7.48 9.22 7.60 6.45 9.75 8.30 8.28 CaO 9.79 10.32 9.43 9.25 10.78 10.86 9.62 Na2O 3.62 3.11 3.80 3.97 2.72 2.90 3.29 K2O 1.64 1.17 1.56 1.76 0.78 0.95 1.36 P2O5 0.47 0.40 0.53 0.61 0.40 0.45 0.49 LOI -0.48 0.07 -0.39 -0.35 -0.58 -0.27 0.09 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor 1.005 0.999 1.004 1.004 0.986 1.003 1.000 Li 6.6 5.4 6.5 6.9 4.9 5.6 5.2 Be 1.7 1.2 1.7 1.8 1.0 1.1 1.4 Sc 23 29 26 24 22 31 29 25 V 267 232 230 248 256 277 293 251 Cr 188 431 426 238 159 346 289 242 Co 40 44 43 39 35 55 47 40 Ni 106 184 201 119 78 245 124 148 Cu 34 44 43 36 37 116 52 38 Zn 89 86 88 92 94 94 100 91 Ga 21 19 18 21 22 19 20 21 Rb 33 26 25 35 37 17 20 28 Sr 680 512 540 623 721 470 555 523 Y 25 24 25 26 27 24 27 28 Zr 247 203 206 245 274 160 182 243 Nb 47 36 40 43 51 30 35 42 Cs 0.38 0.36 0.27 0.40 0.37 0.19 0.15 0.25 Ba 451 320 318 454 484 236 291 615 La 36 26 27 35 39 22 27 30 Ce 74 56 58 73 81 47 56 64 Pr 9.0 7.1 7.8 8.9 9.7 6.0 7.1 8.2 Nd 35 29 33 35 38 25 29 33 Sm 7.4 6.4 6.6 7.2 7.9 5.8 6.7 7.6 Eu 2.3 2.0 2.1 2.2 2.4 1.9 2.1 2.3 Gd 7.1 6.4 6.3 7.1 7.5 5.9 6.7 7.3 Tb 0.94 0.87 0.90 0.96 0.99 0.83 0.93 1.00 Dy 5.0 4.7 4.9 5.1 5.2 4.6 5.1 5.5 Ho 0.97 0.92 0.93 0.98 1.02 0.92 1.01 1.07 Er 2.52 2.40 2.26 2.58 2.68 2.42 2.63 2.82 Tm 0.34 0.32 0.33 0.34 0.36 0.33 0.37 0.38 Yb 2.0 2.0 2.0 2.1 2.2 2.0 2.2 2.3 Lu 0.30 0.29 0.31 0.30 0.32 0.29 0.32 0.34 Hf 5.7 4.8 4.9 5.7 6.2 3.9 4.4 5.7 Ta 3.2 2.5 2.6 3.0 3.4 2.0 2.4 2.9 Pb 2.95 2.12 1.36 2.33 Th 3.6 2.7 2.8 3.7 4.0 2.2 2.7 3.2 U 1.06 0.81 0.82 1.07 1.18 0.65 0.73 0.96 87Sr/ 86Sr 0.70396 0.70381 0.70381 0.70402 0.70392 0.70381 0.70391 0.70366 143Nd/144Nd 0.512816 0.512891 0.512902 0.512822 0.512836 0.512874 0.512873 0.512902 εNd 3.5 4.9 5.1 3.6 3.9 4.6 4.6 5.2 206Pb/204Pb 19.247 19.436 19.440 19.307 19.299 19.316 19.299 19.336 207Pb/204Pb 15.648 15.632 15.635 15.642 15.626 15.621 15.618 15.606 208Pb/204Pb 39.051 39.191 39.204 39.163 39.018 39.111 39.087 39.062 176Hf/177Hf 0.28293 0.28298 0.28300 0.28291 0.28295 0.28296 0.28300 0.28302 εHf 5.5 7.4 7.9 5.0 6.2 6.7 8.0 8.9 187Os/188Os 0.1365 0.1264 0.1307 0.1284 0.1284 0.1336 0.1226 Os ppt 2.6 11.4 4.6 11.3 58.2 5.0 15.7 49

AF4 AF15 AGR14 WAP3 WAP6 WAP8 WAP9 WAP10 SiO2 48.04 47.88 48.66 47.39 47.63 46.84 47.63 48.58 TiO2 2.66 2.67 2.44 2.23 2.65 2.47 2.13 1.31 Al2O3 15.78 16.04 16.43 11.50 13.22 11.86 11.87 7.72 Total Fe 10.52 10.44 10.71 11.18 10.64 10.81 10.14 8.21 MnO 0.16 0.16 0.17 0.18 0.17 0.18 0.16 0.14 MgO 7.74 7.69 7.25 11.84 10.43 12.05 13.21 17.04 CaO 9.29 9.29 9.07 11.84 10.29 12.03 10.95 14.79 Na2O 3.73 3.76 3.60 2.49 3.12 2.44 2.47 1.50 K2O 1.57 1.57 1.16 0.85 1.33 0.86 0.98 0.44 P2O5 0.51 0.51 0.53 0.50 0.51 0.46 0.44 0.27 LOI -0.28 -0.36 -0.39 -0.15 -0.21 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor 1.008 0.993 1.001 1.003 1.004 1.004 1.001 1.002 Li 5.3 5.8 4.7 5.2 3.0 Be 1.3 1.7 1.3 1.4 0.7 Sc 24 23 22 35 28 37 33 52 V 240 239 239 255 226 248 191 123.28 Cr 247 234 147 610 631 651 1029 2537.00 Co 41 40 38 54 50 57 56 58 Ni 137 132 103 209 220 255 335 387 Cu 34 35 22 77 59 58 58 29 Zn 101 96 94 97 97 92 83 62 Ga 20 20 20 17 18 16 15 10 Rb 36 36 26 18 27 17 20 9 Sr 672 659 585 422 500 417 422 230 Y 27 26 28 23 25 23 21 13 Zr 248 242 185 154 203 159 156 76 Nb 49 48 41 30 40 29 28 13 Cs 0.27 0.31 0.20 0.21 0.25 0.19 0.22 0.10 Ba 440 422 358 220 347 225 285 112 La 37 36 28 23 31 22 23 10 Ce 75 73 58 51 68 52 51 23 Pr 9.6 9.4 7.8 6.3 8.1 6.4 6.2 2.9 Nd 39 38 33 27 33 27 25 13 Sm 7.5 7.3 6.9 6.3 7.3 6.4 5.8 3.2 Eu 2.4 2.3 2.4 2.0 2.2 2.0 1.8 1.0 Gd 6.8 6.7 6.7 5.5 6.2 5.5 5.0 2.9 Tb 0.98 0.95 1.00 0.86 0.95 0.86 0.77 0.48 Dy 5.1 5.1 5.3 4.8 5.2 4.8 4.3 2.7 Ho 0.98 0.96 1.01 0.92 1.00 0.92 0.82 0.53 Er 2.4 2.3 2.5 2.2 2.5 2.3 2.0 1.3 Tm 0.35 0.34 0.36 0.31 0.34 0.31 0.29 0.18 Yb 2.1 2.1 2.2 1.8 2.0 1.9 1.7 1.1 Lu 0.33 0.31 0.34 0.26 0.29 0.27 0.24 0.16 Hf 5.9 5.7 4.4 4.1 5.1 4.1 4.1 2.0 Ta 3.2 3.1 2.7 2.1 2.8 2.1 2.0 0.9 Pb 4.64 3.59 2.41 1.44 1.79 0.80 Th 3.7 3.7 2.8 2.3 3.2 2.2 2.3 0.9 U 1.01 1.03 0.78 0.48 1.00 0.74 0.67 0.29 87Sr/ 86Sr 0.70401 0.70400 0.70338 0.70383 0.70382 0.70363 0.70393 0.70378 143Nd/144Nd 0.512838 0.512838 0.512968 0.512855 0.512860 0.512896 0.512839 0.512901 εNd 3.9 3.9 6.4 4.2 4.3 5.0 3.9 5.1 206Pb/204Pb 19.052 19.185 19.834 19.687 19.563 19.977 19.519 19.887 207Pb/204Pb 15.645 15.645 15.641 15.573 15.560 15.598 15.588 15.571 208Pb/204Pb 38.905 39.051 39.275 39.151 39.054 39.335 39.093 39.068 176Hf/177Hf 0.28292 0.28293 0.28307 0.28301 0.28297 0.28303 0.28296 0.28300 εHf 5.2 5.5 10.6 8.3 7.1 9.1 6.6 8.2 187Os/188Os 0.1513 0.1432 0.1761 0.1250 0.1262 0.1249 0.1255 0.1244 Os ppt 3.2 3.1 10.7 65.3 62.6 22.5 139.6 94.1 50

WAP12 WAP25 WAP28 WAP29 WAP30 WAP33a AP1 AP4 SiO2 47.63 46.99 47.44 48.24 47.43 47.43 46.78 47.20 TiO2 2.46 2.01 2.47 2.22 2.70 2.62 2.64 2.77 Al2O3 13.85 10.11 12.75 12.93 14.71 13.35 13.18 14.66 Total Fe 10.21 10.83 10.78 10.01 9.92 10.55 11.54 11.39 MnO 0.17 0.18 0.18 0.17 0.16 0.17 0.16 0.16 MgO 10.56 14.40 9.82 10.72 8.73 10.81 11.21 9.43 CaO 10.65 12.33 12.30 11.30 10.84 11.15 10.67 10.00 Na2O 2.95 2.12 2.89 2.97 3.79 2.61 2.65 3.01 K2O 0.97 0.71 1.00 1.04 1.17 0.93 0.86 0.95 P2O5 0.54 0.32 0.38 0.40 0.55 0.39 0.31 0.43 LOI -0.49 -0.28 -0.50 -0.32 -0.23 -0.28 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor 1.005 1.003 1.003 1.003 0.981 1.003 1.009 0.995

Li 5.3 4.5 5.2 5.7 6.5 4.4 Be 1.4 1.1 1.4 1.5 1.7 1.4 Sc 28 41 37 32 31 31 27 24 V 239 208 276 228 284 248 269 230 Cr 513 1187 485 630 258 550 612 343 Co 50 60 50 48 39 51 54 49 Ni 239 300 181 202 84 230 325 186 Cu 61 70 67 54 58 61 65 34 Zn 90 102 93 89 99 91 97 101 Ga 17 14 17 17 20 18 17 18 Rb 20 15 20 19 26 20 17 20 Sr 458 344 463 468 547 501 540 522 Y 23 20 24 23 26 23 22 25 Zr 162 128 168 164 229 173 159 208 Nb 32 22 33 31 36 32 33 40 Cs 0.23 0.15 0.23 0.23 0.37 0.20 0.09 0.17 Ba 255 194 266 280 318 261 228 240 La 24 18 25 24 31 27 22 24 Ce 54 40 57 53 66 59 46 52 Pr 6.5 4.9 7.0 6.4 8.2 7.2 6.3 7.1 Nd 27 21 29 26 33 30 27 30 Sm 6.4 5.1 6.6 6.2 7.3 6.8 5.6 6.2 Eu 2.0 1.6 2.0 1.9 2.3 2.0 1.9 2.1 Gd 5.5 4.5 5.7 5.4 6.4 5.8 5.5 6.2 Tb 0.88 0.73 0.90 0.86 1.00 0.90 0.81 0.92 Dy 4.9 4.2 5.0 4.8 5.4 5.0 4.4 4.9 Ho 0.93 0.80 0.96 0.92 1.02 0.94 0.80 0.91 Er 2.3 2.0 2.4 2.3 2.6 2.3 1.9 2.2 Tm 0.32 0.27 0.33 0.31 0.34 0.32 0.28 0.32 Yb 1.9 1.6 2.0 1.9 2.1 1.9 1.6 1.9 Lu 0.27 0.23 0.28 0.27 0.32 0.27 0.25 0.29 Hf 4.2 3.4 4.4 4.2 5.0 4.5 4.0 4.8 Ta 2.2 1.6 2.3 2.2 2.7 2.4 2.1 2.6 Pb 1.39 1.91 1.62 1.57 1.66 Th 2.3 1.7 2.5 2.5 3.0 2.7 2.0 2.6 U 0.76 0.51 0.79 0.76 0.94 0.83 0.60 0.84

87Sr/ 86Sr 0.70373 0.70382 0.70368 0.70384 0.70379 0.70378 0.70379 0.70350 143Nd/144Nd 0.512878 0.512872 0.512885 0.512881 0.512884 0.512871 0.512881 0.512918 εNd 4.7 4.6 4.8 4.7 4.8 4.5 4.7 5.5 206Pb/204Pb 19.721 19.943 19.851 19.734 19.888 19.478 19.646 19.836 207Pb/204Pb 15.581 15.604 15.575 15.572 15.632 15.581 15.653 15.633 208Pb/204Pb 39.104 39.366 39.184 39.096 39.398 39.033 39.350 39.334 176Hf/177Hf 0.28302 0.28304 0.28299 0.28300 0.28300 0.28296 0.28299 0.28304 εHf 8.6 9.3 7.6 8.1 8.0 6.6 7.8 9.6 187Os/188Os 0.1237 0.1255 0.1262 0.1219 0.1276 0.1227 0.1273 0.1220 Os ppt 13.2 259.2 20.3 12.4 12.2 73.3 63.3 14.8 51

AP7 AP14 WASJ2 WASJ3 WASJ4 WASJ5 WASJ6 WASJ7 SiO2 46.63 49.71 46.77 46.39 46.69 44.84 47.94 43.47 TiO2 2.75 2.00 3.46 3.09 3.19 3.78 3.08 4.20 Al2O3 13.47 14.11 16.76 15.30 16.30 15.74 17.20 15.19 Total Fe 11.28 10.74 12.20 11.88 11.42 12.82 11.71 13.65 MnO 0.16 0.15 0.18 0.18 0.16 0.17 0.20 0.17 MgO 10.47 9.58 5.78 8.02 7.23 7.50 4.38 8.70 CaO 10.88 10.15 8.84 10.36 10.09 10.22 8.12 10.27 Na2O 2.93 2.76 3.81 2.81 3.30 3.28 4.61 2.89 K2O 1.02 0.57 1.41 1.35 1.11 1.04 1.67 0.93 P2O5 0.41 0.24 0.78 0.63 0.52 0.61 1.09 0.51 LOI -0.25 -0.03 -0.37 -0.40 -0.67 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor 1.009 1.002 0.998 0.994 0.995 0.994 0.998 0.997 Li 6.3 6.1 4.9 Be 1.6 1.6 1.4 Sc 28 25 25 25 28 V 256 212 290 317 337 Cr 604 491 232 121 201 Co 50 46 44 47 57 Ni 234 233 115 85 105 Cu 47 42 47 33 32 Zn 104 90 106 121 123 Ga 17 17 22 22 22 Rb 19 11 26 21 20 Sr 526 374 675 625 639 Y 25 22 28 30 28 Zr 193 124 267 250 247 Nb 39 21 51 46 46 Cs 0.18 0.11 0.30 0.29 0.30 Ba 255 149 380 265 245 La 25 14 41 31 31 Ce 53 31 85 70 69 Pr 7.2 4.3 10.4 9.1 9.0 Nd 31 19 41 38 37 Sm 6.4 4.6 8.4 8.6 8.3 Eu 2.1 1.7 2.6 2.7 2.6 Gd 6.2 4.9 7.2 7.4 7.1 Tb 0.90 0.77 1.07 1.13 1.09 Dy 4.7 4.2 5.7 6.0 5.8 Ho 0.91 0.80 1.08 1.14 1.08 Er 2.2 2.0 2.8 2.9 2.7 Tm 0.31 0.29 0.36 0.37 0.35 Yb 1.9 1.7 2.2 2.3 2.1 Lu 0.28 0.26 0.32 0.33 0.30 Hf 4.5 3.1 6.2 5.9 6.0 Ta 2.5 1.4 3.5 3.3 3.4 Pb 1.91 1.05 1.67 1.67 Th 2.4 1.3 3.8 2.5 2.7 U 0.75 0.44 1.11 0.92 0.97 87Sr/ 86Sr 0.70371 0.70352 0.70345 0.70386 0.70344 0.70342 0.70345 0.70343 143Nd/144Nd 0.512904 0.512952 0.512927 0.512883 0.512931 0.512929 0.512912 0.512909 εNd 5.2 6.1 5.6 4.8 5.7 5.7 5.3 5.3 206Pb/204Pb 19.712 19.601 19.878 19.363 19.826 20.081 19.962 19.960 207Pb/204Pb 15.649 15.635 15.604 15.623 15.605 15.622 15.615 15.606 208Pb/204Pb 39.340 39.137 39.327 39.119 39.252 39.370 39.349 39.350 176Hf/177Hf 0.28301 0.28308 0.28295 0.28301 0.28302 εHf 8.4 11.0 6.2 8.4 8.9 187Os/188Os 0.1629 0.1419 0.1373 0.1719 Os ppt 2.0 5.4 3.3 1.9 52

WASJ8 WASJ9 WASJ10 WASJ11 WASJ12 WASJ13 WASJ14 WASJ15 SiO2 45.97 48.15 46.66 49.69 48.57 45.93 49.09 48.56 TiO2 3.60 3.07 3.54 2.88 3.15 2.93 2.87 2.07 Al2O3 17.35 17.15 16.67 17.44 16.58 15.01 16.06 23.72 Total Fe 12.10 11.38 12.29 11.20 11.94 12.56 11.40 7.47 MnO 0.17 0.19 0.17 0.18 0.21 0.17 0.16 0.11 MgO 5.71 4.66 6.13 4.25 4.22 9.99 6.21 2.49 CaO 9.38 8.16 8.77 7.25 8.04 9.93 8.87 11.04 Na2O 3.75 4.55 3.83 4.58 4.69 2.34 3.71 3.36 K2O 1.24 1.72 1.25 1.73 1.47 0.73 1.06 0.81 P2O5 0.71 0.97 0.69 0.79 1.14 0.41 0.57 0.38 LOI -0.59 -0.56 -0.48 -0.49 -0.41 0.88 -0.41 0.08 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor 1.000 0.997 0.994 0.998 0.996 0.986 0.993 0.996 Li 6.1 4.9 5.8 4.2 Be 1.7 1.2 1.7 1.1 Sc 19 27 20 12 V 260 264 249 148 Cr 72 348 153 6 Co 38 55 38 18 Ni 59 230 80 14 Cu 31 45 37 32 Zn 115 107 114 74 Ga 22 19 22 23 Rb 25 13 24 18 Sr 685 459 546 845 Y 32 26 33 23 Zr 288 190 257 170 Nb 53 35 46 27 Cs 0.31 0.20 0.22 0.33 Ba 316 236 289 227 La 35 24 34 21 Ce 78 52 72 45 Pr 10.1 6.8 8.9 5.8 Nd 42 28 36 25 Sm 9.2 6.7 8.0 5.9 Eu 2.9 2.2 2.6 2.2 Gd 7.9 6.1 7.2 5.3 Tb 1.21 0.98 1.14 0.85 Dy 6.4 5.4 6.2 4.6 Ho 1.22 1.03 1.20 0.88 Er 3.1 2.6 3.1 2.2 Tm 0.40 0.34 0.40 0.29 Yb 2.5 2.1 2.5 1.8 Lu 0.36 0.30 0.35 0.25 Hf 6.6 4.7 5.9 4.1 Ta 3.6 2.5 3.1 1.9 Pb 1.96 Th 3.0 2.2 3.1 1.6 U 1.04 0.83 1.09 0.63 87Sr/ 86Sr 0.70343 0.70343 0.70341 0.70343 0.70374 0.70359 0.70368 0.70376 143Nd/144Nd 0.512921 0.512919 0.512929 0.512922 0.512966 0.512932 0.512940 0.512939 εNd 5.5 5.5 5.7 5.5 6.4 5.7 5.9 5.9 206Pb/204Pb 19.953 19.950 19.884 19.930 19.884 19.895 19.683 19.393 207Pb/204Pb 15.610 15.617 15.617 15.635 15.593 15.611 15.591 15.531 208Pb/204Pb 39.335 39.359 39.304 39.401 39.016 39.258 38.955 38.734 176Hf/177Hf 0.28301 0.28306 0.28304 0.28311 εHf 8.5 10.1 9.5 12.1 187Os/188Os 0.1344 0.1259 Os ppt 9.4 53.5 53

WASJ16 WASJ17 WASJ18 WASJ19 WASJ20 WASJ21 WASJ22 WASJ23 SiO2 53.24 47.78 46.48 46.98 45.80 45.90 49.43 46.55 TiO2 2.17 3.47 3.22 3.10 3.10 3.37 2.91 3.43 Al2O3 16.99 16.84 15.01 15.04 15.38 16.23 16.98 16.77 Total Fe 9.76 12.44 11.49 11.14 12.25 12.32 11.40 12.41 MnO 0.18 0.19 0.17 0.16 0.18 0.18 0.19 0.18 MgO 2.83 4.62 9.72 8.88 8.24 6.78 4.48 5.43 CaO 6.21 8.11 9.78 10.19 10.37 10.54 7.52 8.80 Na2O 5.27 4.56 2.62 2.90 3.17 3.23 4.55 4.06 K2O 2.51 1.27 0.98 1.08 1.01 0.97 1.73 1.50 P2O5 0.84 0.72 0.54 0.52 0.51 0.48 0.81 0.87 LOI -0.15 -0.63 0.02 -0.32 -0.30 0.74 -0.54 -0.55 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor 1.000 0.997 0.993 0.990 0.999 0.981 0.996 0.994 Li 10.3 8.4 5.5 5.6 5.5 6.1 Be 2.7 2.3 1.5 1.6 1.6 1.6 Sc 13 17 26 27 22 34 V 134 243 265 280 266 301 Cr 5 4 322 357 172 81 Co 17 32 49 47 48 43 Ni 8 16 216 171 108 53 Cu 16 68 46 47 34 30 Zn 128 134 104 96 111 109 Ga 27 25 20 20 20 20 Rb 56 29 18 24 23 20 Sr 501 679 537 562 647 664 Y 54 39 27 27 28 28 Zr 475 338 264 267 255 227 Nb 77 56 43 45 41 35 Cs 0.50 0.34 0.08 0.27 0.31 0.28 Ba 513 347 289 293 303 230 La 62 42 32 36 34 27 Ce 131 90 69 77 72 60 Pr 16.1 11.4 8.9 9.5 9.0 7.9 Nd 64 48 37 38 37 34 Sm 14.1 10.4 7.8 7.9 8.0 7.8 Eu 4.1 3.3 2.4 2.4 2.5 2.5 Gd 12.4 9.3 6.8 6.9 7.1 6.9 Tb 1.97 1.46 1.04 1.04 1.10 1.09 Dy 10.7 8.0 5.5 5.5 5.9 5.9 Ho 2.08 1.53 1.05 1.06 1.12 1.13 Er 5.4 3.9 2.7 2.7 2.9 2.9 Tm 0.72 0.51 0.35 0.36 0.38 0.38 Yb 4.5 3.1 2.2 2.2 2.4 2.3 Lu 0.65 0.46 0.32 0.32 0.34 0.34 Hf 12.1 7.1 5.7 5.9 5.6 5.1 Ta 5.1 4.0 3.1 3.3 3.0 2.6 Pb 3.41 2.47 1.75 1.77 Th 6.8 4.0 3.0 3.3 3.1 2.3 U 2.13 1.45 0.95 1.06 1.08 0.86 87Sr/ 86Sr 0.70362 0.70370 0.70350 0.70374 0.70376 0.70376 0.70342 0.70343 143Nd/144Nd 0.512952 0.512960 0.512916 0.512889 0.512950 0.512977 0.512917 0.512915 εNd 6.1 6.3 5.4 4.9 6.1 6.6 5.4 5.4 206Pb/204Pb 19.531 20.056 19.691 19.339 19.804 20.367 19.912 19.897 207Pb/204Pb 15.547 15.615 15.630 15.612 15.619 15.659 15.612 15.612 208Pb/204Pb 38.873 39.154 39.344 39.012 39.043 39.376 39.324 39.327 176Hf/177Hf 0.28306 0.28311 0.28302 0.28296 0.28307 0.28309 εHf 10.2 11.9 8.7 6.6 10.6 11.3 187Os/188Os 0.1254 0.1263 0.1324 Os ppt 92.6 25.3 3.1 54

WASJ24 WASJ25 WASJ26 WASJ27 WASJ28 WASJ29 WASJ30 WASJ31 SiO2 45.81 45.97 45.94 45.00 47.91 45.12 45.69 44.95 TiO2 3.71 3.58 3.51 2.82 3.05 3.44 3.56 3.97 Al2O3 16.77 16.90 16.32 13.20 16.74 14.88 16.40 16.79 Total Fe 12.66 13.02 12.55 11.91 11.75 12.15 12.63 12.52 MnO 0.18 0.19 0.19 0.17 0.20 0.17 0.18 0.17 MgO 5.78 5.34 6.52 12.02 4.57 9.17 6.45 6.19 CaO 9.12 8.87 9.94 11.30 8.35 10.56 9.18 10.02 Na2O 3.80 3.79 3.18 2.40 4.54 2.99 3.82 3.49 K2O 1.43 1.39 1.16 0.75 1.70 1.03 1.29 1.23 P2O5 0.74 0.96 0.69 0.41 1.20 0.48 0.81 0.67 LOI -0.28 -0.18 0.06 -0.56 -0.38 -0.47 -0.31 -0.44 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor 0.995 0.998 0.989 1.002 1.001 1.001 0.997 0.997 Li 4.5 5.3 Be 1.2 1.4 Sc 31 26 V 258 315 Cr 366 277 Co 65 52 Ni 236 145 Cu 27 41 Zn 98 109 Ga 18 21 Rb 16 21 Sr 506 550 Y 25 24 Zr 202 238 Nb 34 38 Cs 0.16 0.20 Ba 200 249 La 23 27 Ce 56 60 Pr 6.9 7.7 Nd 30 32 Sm 7.0 7.3 Eu 2.2 2.3 Gd 5.8 6.4 Tb 0.92 0.98 Dy 5.1 5.2 Ho 0.98 0.99 Er 2.5 2.5 Tm 0.32 0.32 Yb 1.9 2.0 Lu 0.27 0.29 Hf 5.0 5.3 Ta 2.3 2.9 Pb 1.72 Th 2.2 2.5 U 0.80 0.83 87Sr/ 86Sr 0.70345 0.70342 0.70351 0.70349 0.70343 0.70348 0.70343 0.70342 143Nd/144Nd 0.512926 0.512919 0.512922 0.512930 0.512916 0.512907 0.512920 0.512901 εNd 5.6 5.5 5.5 5.7 5.4 5.2 5.5 5.1 206Pb/204Pb 19.890 19.858 19.790 19.855 19.955 19.794 19.985 19.909 207Pb/204Pb 15.605 15.624 15.603 15.601 15.609 15.636 15.618 15.611 208Pb/204Pb 39.280 39.306 39.228 39.209 39.330 39.359 39.374 39.334 176Hf/177Hf 0.28303 0.28301 εHf 9.0 8.3 187Os/188Os 0.1238 Os ppt 8.6 55

WASJ32 ASJ3 ASJ7 WAT1 WAT2 WAT3 WAT4 WAT5 SiO2 45.78 45.09 45.21 46.74 47.01 47.75 47.02 49.41 TiO2 3.52 3.47 4.05 4.22 3.62 1.90 2.68 2.28 Al2O3 17.19 14.51 16.81 14.67 14.55 13.36 13.83 15.15 Total Fe 12.10 12.56 12.80 13.77 13.51 9.68 11.35 9.86 MnO 0.18 0.17 0.17 0.21 0.22 0.15 0.17 0.16 MgO 5.71 9.51 5.92 5.02 5.49 11.80 9.66 8.10 CaO 9.39 10.20 9.60 10.19 9.18 11.54 11.53 10.09 Na2O 3.93 2.98 3.50 3.45 3.67 2.55 2.51 3.12 K2O 1.44 1.05 1.26 1.21 1.23 0.98 0.75 1.44 P2O5 0.77 0.45 0.67 0.52 1.51 0.29 0.50 0.39 LOI -0.53 -0.32 0.10 -0.15 -0.18 -0.08 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor 1.011 0.999 0.993 1.003 0.997 0.994 0.991 0.999 Li 5.7 4.3 4.3 6.2 Be 1.6 1.2 1.0 1.7 Sc 26 20 27 31 31 29 V 286 272 416 242 271 255 Cr 326 88 21 780 508 383 Co 53 41 45 54 52 41 Ni 178 56 32 324 216 136 Cu 42 32 36 77 57 52 Zn 112 126 131 69 86 86 Ga 20 23 22 15 17 19 Rb 21 25 25 23 15 30 Sr 581 758 540 381 459 389 Y 27 33 34 20 25 25 Zr 233 309 236 127 137 193 Nb 46 62 50 29 29 37 Cs 0.16 0.25 0.29 0.23 0.15 0.10 Ba 263 310 340 270 289 419 La 27 37 33 19 22 32 Ce 59 81 71 40 49 59 Pr 8.2 11.2 9.0 4.9 6.4 7.7 Nd 35 47 37 20 28 30 Sm 7.2 9.6 8.5 4.7 6.8 6.5 Eu 2.4 3.1 2.7 1.5 2.2 2.0 Gd 7.1 8.8 7.5 4.1 5.8 5.9 Tb 1.00 1.24 1.21 0.66 0.90 0.95 Dy 5.3 6.6 6.6 3.8 5.1 5.4 Ho 0.98 1.21 1.30 0.77 0.99 1.07 Er 2.4 2.9 3.4 2.1 2.6 2.9 Tm 0.34 0.40 0.45 0.27 0.32 0.38 Yb 2.0 2.4 2.7 1.7 2.0 2.4 Lu 0.31 0.36 0.40 0.24 0.27 0.34 Hf 5.5 7.1 5.7 3.2 3.4 4.9 Ta 3.0 4.2 3.5 1.9 2.0 2.7 Pb 1.56 1.97 1.98 1.56 1.07 Th 2.7 3.6 3.0 2.0 1.8 3.4 U 0.86 1.12 1.08 0.61 0.63 0.97 87Sr/ 86Sr 0.70343 0.70350 0.70343 0.70360 0.70357 0.70344 0.70355 0.70350 143Nd/144Nd 0.512913 0.512917 0.512924 0.512950 0.512950 0.512913 0.512948 0.512891 εNd 5.4 5.4 5.6 6.1 6.1 5.4 6.0 4.9 206Pb/204Pb 19.909 19.682 19.729 19.982 19.908 19.768 19.766 19.598 207Pb/204Pb 15.616 15.642 15.630 15.611 15.601 15.605 15.589 15.574 208Pb/204Pb 39.341 39.286 39.252 39.278 39.253 39.272 39.120 39.096 176Hf/177Hf 0.28300 0.28300 0.28305 0.28304 0.28307 0.28300 εHf 8.0 8.2 9.9 9.5 10.6 8.1 187Os/188Os 0.1528 0.1252 0.1272 0.1379 Os ppt 3.8 0.8 111.9 18.2 11.0 56

WAT6 WAT7 WAT8 WAT9 WAT10 WAT11 WAT12 WAT13 SiO2 47.56 47.61 47.45 47.09 51.87 50.26 49.47 47.24 TiO2 3.49 3.30 3.74 3.59 2.79 3.31 3.12 3.01 Al2O3 14.22 14.03 13.95 14.22 15.90 14.75 15.00 13.95 Total Fe 12.73 12.24 13.50 12.90 11.06 12.75 12.20 11.99 MnO 0.20 0.19 0.23 0.20 0.21 0.21 0.22 0.18 MgO 6.19 7.19 5.35 6.43 3.45 3.91 4.54 8.21 CaO 9.83 10.18 9.34 10.22 7.42 8.01 8.37 10.52 Na2O 3.57 3.36 3.81 3.36 4.56 4.40 4.38 3.14 K2O 1.23 1.15 1.30 1.07 1.83 1.49 1.34 0.89 P2O5 0.98 0.75 1.33 0.91 0.91 0.90 1.37 0.88 LOI -0.67 -0.53 -0.52 -0.62 -0.22 -0.35 -0.47 -0.47 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor 1.008 1.003 0.998 1.000 0.999 1.000 0.997 0.997 Li 5.4 8.4 4.6 Be 1.2 2.2 1.0 Sc 26 18 27 V 294 283 283 Cr 282 3 278 Co 41 28 44 Ni 102 5 120 Cu 42 17 34 Zn 113 145 102 Ga 21 26 20 Rb 22 31 19 Sr 461 576 494 Y 31 44 31 Zr 184 302 169 Nb 35 63 35 Cs 0.07 0.22 0.12 Ba 453 397 457 La 33 53 33 Ce 71 112 72 Pr 9.2 14.1 9.4 Nd 39 57 40 Sm 9.2 13.1 9.4 Eu 3.1 3.9 3.4 Gd 8.3 11.8 8.5 Tb 1.30 1.84 1.31 Dy 7.0 10.0 7.0 Ho 1.36 1.91 1.33 Er 3.4 4.9 3.4 Tm 0.44 0.63 0.42 Yb 2.6 3.8 2.4 Lu 0.37 0.53 0.34 Hf 4.7 8.4 4.1 Ta 2.7 4.7 2.5 Pb 1.54 Th 2.7 5.8 2.2 U 0.76 1.92 0.70 87Sr/ 86Sr 0.70344 0.70356 0.70341 0.70342 0.70354 0.70357 0.70354 0.70345 143Nd/144Nd 0.512942 0.512930 0.512941 0.512944 0.512950 0.512942 0.512954 0.512946 εNd 5.9 5.7 5.9 6.0 6.1 5.9 6.2 6.0 206Pb/204Pb 19.684 19.673 19.603 19.687 19.975 20.002 19.828 19.617 207Pb/204Pb 15.581 15.592 15.569 15.578 15.609 15.612 15.583 15.592 208Pb/204Pb 39.127 39.136 39.019 39.126 39.231 39.226 39.111 39.087 176Hf/177Hf 0.28305 0.28307 0.28308 εHf 9.7 10.7 10.8 187Os/188Os 0.1286 0.1288 Os ppt 4.4 12.9 8.3 4.2 4.0 1.0 9.9 57

WAT14 WAT15 WAT16 WAT17 WAT18 AT19 AT28 AT30 SiO2 47.64 46.53 46.67 47.33 47.55 49.34 47.94 47.05 TiO2 3.40 3.28 3.80 3.62 3.51 3.25 1.93 3.82 Al2O3 14.33 13.64 14.15 14.14 14.32 15.00 13.61 14.05 Total Fe 12.42 12.51 13.72 13.23 12.75 12.49 9.72 13.65 MnO 0.20 0.19 0.20 0.19 0.21 0.23 0.15 0.21 MgO 6.45 8.23 5.86 6.22 6.20 4.41 11.52 5.65 CaO 10.13 10.46 9.91 10.23 9.85 8.15 11.29 9.55 Na2O 3.44 3.21 3.59 3.31 3.50 4.34 2.55 3.63 K2O 1.10 0.94 1.04 0.96 1.11 1.42 1.01 1.25 P2O5 0.90 1.01 1.05 0.76 1.02 1.38 0.28 1.14 LOI -0.63 -0.52 -0.37 -0.21 -0.61 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 Norm Factor 1.006 0.997 1.000 0.994 0.999 1.010 0.999 1.011 Li 4.3 5.0 Be 1.1 1.2 Sc 29 27 17 32 24 V 306 335 206 258 310 Cr 384 103 20 781 134 Co 43 38 23 48 36 Ni 127 49 13 290 51 Cu 43 29 16 72 30 Zn 103 116 134 73 127 Ga 19 21 23 14 20 Rb 22 21 29 24 26 Sr 504 540 700 399 534 Y 34 37 52 20 42 Zr 176 189 288 126 206 Nb 40 41 68 32 52 Cs 0.27 0.21 0.25 0.22 0.12 Ba 576 514 658 283 663 La 35 38 54 19 42 Ce 78 83 115 39 91 Pr 10 11 16 5 13 Nd 44 48 70 22 57 Sm 10.1 10.8 14.8 4.4 11.9 Eu 3.9 3.8 5.2 1.5 4.2 Gd 8.8 9.6 14.4 4.5 11.6 Tb 1.33 1.45 2.00 0.66 1.60 Dy 7.0 7.6 10.3 3.7 8.3 Ho 1.32 1.44 1.91 0.72 1.57 Er 3.3 3.6 4.5 1.8 3.7 Tm 0.41 0.44 0.63 0.27 0.51 Yb 2.5 2.7 3.6 1.6 3.0 Lu 0.35 0.38 0.55 0.25 0.45 Hf 4.3 4.6 6.9 3.2 4.9 Ta 2.7 2.9 4.4 2.1 3.4 Pb 1.21 2.13 1.70 1.51 Th 2.1 2.4 4.1 2.1 2.9 U 0.72 0.75 1.31 0.58 0.87 87Sr/ 86Sr 0.70346 0.70 347 0.70350 0.70350 0.70343 0.70356 0.70344 0.70342 143Nd/144Nd 0.512939 0.512935 0.512940 0.512938 0.512940 0.512960 0.512921 0.512958 εNd 5.9 5.8 5.9 5.8 5.9 6.3 5.5 6.2 206Pb/204Pb 19.702 19.492 19.830 19.729 19.694 19.832 19.689 19.589 207Pb/204Pb 15.585 15.567 15.621 15.590 15.591 15.618 15.592 15.593 208Pb/204Pb 39.149 39.012 39.256 39.120 39.163 39.206 39.225 39.084 176Hf/177Hf 0.28305 0.28306 0.28308 0.28303 0.28308 εHf 9.8 10.0 11.0 9.0 11.0 187Os/188Os 0.1245 0.1457 0.1774 0.1277 0.1447 Os ppt 5.4 8.1 1.4 54.1 3.4 Norm factor* represents the normalizing factor used to correct the major element totals to 100%. 58

Table 2: The compositions of end-members in the Os-Sr-Pb mixing models

187Re/188Os 187Os/188Os 238U/204Pb 206Pb/204Pb 87Rb/86Sr 87Sr/86Sr Os ppt Pb ppm Sr ppm OC1 1000a 0.127c 37f 18.0g 0.2i 0.7022g 2 0.1f 97f OC2 325a 0.127c 37f 18.0g 0.2i 0.7022g 30 0.1f 97f PS 500b 0.7d 37f 18.9h 0.2i 0.709i 30k 20i 600i TS 500b 1.2e 20i 18.9h 0.5i 0.715i 30k 10i 300i AMP 0.122c 19.1g 0.7035j 3100l 0.02m 10m OC1 is recycled oceanic crust with low Os concentration and high 187Re/188Os ratio; OC2 is recycled oceanic crust with higher Os concentration and low 187Re/188Os ratio; PS is pelagic sediment; TS is terrigenous sediment; AMP is enriched mantle plume. 187Re/188Os ratios are based on the upper and lower limits for dehydrated, recycled oceanic crust as estimated by Becker (2000). 238U/204Pb and 87Rb/86Sr ratios are estimated ratios after subduction as calculated by Weaver (1991) and Becker et al. (2000). a: Becker (2000); b: Bendall et al. (2009); c: Shirey and Walker (1998); d: Peucher-Ehrenbrink (1995); e: Widom et al. (1999); f: Becker et al. (2000); g: Hart et al. (1992); h:Plank and Langmuir (1998); i: Weaver (1991); j: Widom et al., 1997; k: Kendall et al. (2009); l: Morgan (1986); m: Salters and Stracke (2004).

59

Table 3: The compositions of end-members in the Nd-Hf mixing models

176Lu/177HfB 176Lu/177HfA 176Hf/177Hf 147Sm/144NdB 147Sm/144NdA 143Nd/144Nd Lu Hf Sm Nd AMP 0.28310 0.513000 0.07 0.31 0.44 1.4 DMM 0.28316 0.513088 0.07 0.31 0.44 1.4 OC 0.027 0.039 0.28316 0.199 0.201 0.513088 0.45 1.8 2.7 7.5 TS 0.005 0.007 0.28147 0.096 0.097 0.511320 0.14 2.8 2.0 13 PS-1 0.006 0.009 0.28214 0.114 0.115 0.511969 0.43 7.1 4.7 25 PS-2 0.065 0.095 0.28290 0.136 0.138 0.512392 1.6 2.4 28 125 GLOSS 0.014 0.021 0.28283 0.130 0.131 0.512180 0.33 2.2 4.6 21

“B” indicates compositions before subduction and “A” indicates compositions after subduction. Isotopic compositions of enriched mantle plume (AMP) are from Stracke et al. (2005) and DMM and oceanic crust (OC) are from Chauvel et al. (2008), and the concentration data are from Sun and McDonough (1989). Isotopic composition data for pre-subduction terrigenous sediment (TS), two types of pelagic sediment (PS-1 and PS-2) and GLOSS are from Plank and Langmuir (1998), Vervoort et al. (1999), and Stracke et al. (2003). Isotopic composition and concentration data for the post-subduction end-members are calculated using the trace element mobility of Stracke et al. (2003). All of the Lu, Hf, Sm and Nd concentrations listed in this table are post-subduction concentrations. PS-1 has lower 176Lu/177Hf and 176Hf/177Hf than PS-2.

60

Figure 1. (a) Map of the Azores Archipelago (modified after Moreira et al., 1999) and (b) map of São Jorge. The Azores archipelago is located between ~37-40°N, in the vicinity of the Mid-Atlantic Ridge (MAR) and the triple junction between the North American, African and Eurasian plates. The archipelago comprises nine islands situated on both sides of the mid-Atlantic ridge. The samples in this study are from the Central Group islands of Faial, Pico, São Jorge and Terceira. The samples from São Jorge are separated into two groups: Topo samples shown as open triangles, and Rosais and Manadas samples shown as filled triangles. Numbers in (b) designate WASJ-series sample numbers.

61

Figure 1

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Figure 2. 87Sr/86Sr versus 206Pb/204Pb ratios for basalts of the Central Group islands, Azores archipelago. APT designates Azores platform tholeiite field (data from White et al., 1976; White and Schilling, 1978; Dupré and Allè, 1980, 1984; Ito et al., 1987; Dosso et al., 1996), and FOZO data are from Stracke et al. (2005). São Jorge-O represent the São Jorge samples from the oldest formation (Topo), and São Jorge-Y represent the samples from the intermediate and youngest formations (Roasais and Manadas). Azores basalts exhibit heterogeneity on an intra-island and archipelago-wide scale, ranging from relatively depleted signatures in the Azores Platform MORB to relatively strong HIMU and EM signatures in the Azores island basalts. Black symbols show the data from literature, with shapes denoting the different islands as in this study (White et al., 1979; Davies et al., 1989; Turner et al., 1997; Moreira et al., 1999; Millet et al., 2009). The numbered end-members are from Millet et al. (2009): 1 is HIMU, 2 is source of E-MORB, 3 is source of Mid-Atlantic ridge and 4 is the mantle source of Faial and Pico basalts.

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Figure 2

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Figure 3. (a) Total alkalis vs. SiO2 diagram for samples in the Central Group islands (after Le Bas et al., 1986). (b) Plot of Ni versus MgO concentrations illustrating the fractionation trend of the Azores lavas (Beier et al., 2010).

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Figure 3

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T Figure 4. Variations of SiO2, CaO, Al2O3, Na2O, TiO2, K2O, Fe2O3 and Al2O3/CaO versus MgO for the basalts from the Central Group islands. Samples with low MgO T contents (<5%) exhibit decreasing TiO2 and Fe2O3 with decreasing MgO, indicating the onset of titanomagnetite fractionation. Two distinct magma groups can be recognized based on K2O at a given MgO content: a lower K2O group defined by samples from São

Jorge and Terceira and a higher K2O group defined by samples from Faial and Pico.

Terceira samples also have lower Al2O3 concentrations compared to the other three islands.

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Figure 4

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Figure 5. Chondrite-normalized REE patterns for Central Group island basalts, using chondrite values from Sun and McDonough (1989). All samples are LREE-enriched. The lack of negative Eu anomalies suggests a limited role for plagioclase fractionation.

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Figure 5

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Figure 6. Primitive-mantle normalized trace element patterns for the Central Group island basalts, using primitive mantle compositions from McDonough and Sun (1995). N-MORB and E-MORB data are from Sun and McDonough (1989); HIMU data are from Newsom et al. (1986), Palacz and Saunders (1986) and Willbold and Stracke (2006); EMI data are from Newsom et al. (1986) and Willbold and Stracke (2006); EMII data are from Newsom et al. (1986), Palacz and Saunders (1986) and Dostal et al, (1982).

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Figure 6

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Figure 7. Trace element ratios including (a) Nb/Zr, (b) La/Yb and (c) Ce/Yb versus wt.% MgO; and (d) K ppm versus Nb ppm. Lack of correlations of Nb/Zr, La/Yb and Ce/Yb ratios with wt.% MgO suggests that variable degrees of melting did not affect such trace element ratios in the Central Group island basalts. Relationship of K versus Nb reveals two distinct compositional trends, in which São Jorge and some Terceira samples have higher Nb concentrations than Faial and Pico samples for a given K content.

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Figure 7

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Figure 8. Whole-rock Sr-Nd-Pb-Hf isotope relationships in the Central Group island basalts. (a) 207Pb/204Pb and (b) 208Pb/204Pb versus 206Pb/204Pb, (c) 143Nd/144Nd versus 87 86 143 144 206 204 176 177 Sr/ Sr, (d) εHf versus εNd, (e) Nd/ Nd versus Pb/ Pb and (f) Hf/ Hf versus 206Pb/204Pb. Black symbols show literature data (White et al., 1979; Davies et al., 1989; Turner et al., 1997; Moreira et al., 1999; Millet et al., 2009), the open triangles show the São Jorge samples from Topo labeled as “São Jorge-O”, and the filled triangles show the São Jorge samples from Rosais and Manadas (for both of the data from this study and from literature) labeled as “São Jorge-Y”. MAR is Mid-Atlantic ridge and the data are from Beier et al. (2008) and references therein. FOZO data are from Stracke et al. (2005). The numbered end-members are same as in the Fig. 2.

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Figure 8

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Figure 9. Variation of 187Os/188Os with (a) Os concentration, (b) 87Sr/86Sr, (c) 143Nd/144Nd, and (d) 206Pb/204Pb for samples with Os>10ppt.

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Figure 9

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Figure 10. Calculated mixing trends between Azores mantle plume (AMP) and potential recycled crustal components. (a) and (b) mixtures of AMP and 0.5Ga and 1Ga recycled oceanic crust with variable 187Re/188Os and Os concentration. OC-1 has 187Re/188Os=1000 and [Os]=2 ppt; OC-2 has 187Re/188Os=325 and [Os]=30 ppt. The compositional data for OC-1 and OC-2 are same for all figures. (c) and (d) mixtures of AMP mixing with 0.5Ga and 1Ga recycled oceanic crust+3% pelagic sediment (PS). (e) and (f) mixtures of AMP mixing with 0.5Ga and 1Ga recycled oceanic crust+3% terrigenous sediment (TS). (g) and (h) mixtures of AMP mixing with 1Ga pelagic and terrigenous sediment with tick marks every 10%. The tick marks in the other figures represent crustal fractions of 1%, 2%, 3%, 4%, 5%, 10%, and then increments of 10%. Compositional data of AMP and each end-member are listed in Table 2.

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Figure 10

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Figure 11. (a) εHf and εNd values of oceanic basalts and other components (modified after Chauvel et al., 2008). The yellow array shows samples from the Central Group islands. (b) Mixing trends between Azores mantle plume (AMP) and 1Ga and 2Ga recycled oceanic crust (OC). (c) Mixing trends between AMP and potential recycled component. End-member 1 is 1Ga oceanic crust (OC), end-member 2 is 1Ga oceanic crust plus 3% GLOSS, end-member 3 is 1Ga oceanic crust plus 3% terrigenous sediment (TS), and end-members 4 and 5 are 1Ga oceanic crust plus 3% pelagic sediment (PS-1 and PS-2, respectively). The mantle array follows the relationship εHf = 1.51εNd + 1.39. All of compositional data for the enriched mantle plume, present-day oceanic crust, terrigenous sediment, pelagic sediment and GLOSS are listed in Table 3.

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Figure 11

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Figure 12. Model of proposed mantle sources beneath the Azores archipelago (modified after Donnelly et al., 2004), in which melting of a high 3He/4He FOZO-like mantle plume produces the Terceira and São Jorge basalts, and heat from the plume causes partial melting of surrounding heterogeneous upper mantle with blobs of recycled metasomatized mantle wedge to produce Faial basalts (no plume material). Pico basalts are produced by mixing between the Faial and Terceira sources, with a limited contribution of plume material.

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Figure 12

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CHAPTER 3

Lithium isotope systematics of basalts from the Azores Archipelago: constraints on the origin of the mantle sources

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Abstract

Basalts from the Azores Archipelago and MORB from the nearby Azores Platform exhibit extreme chemical and isotopic variations attributed to the influence of a heterogeneous mantle plume, with compositions ranging from depleted mantle (DMM) to HIMU, EMI and EMII signatures. In order to assess the utility of Li isotopes as a mantle source tracer and to better constrain the origin of heterogeneous mantle beneath the Azores, we have analyzed Li isotopes in a suite of young, fresh, MgO-rich basalts and olivine separates from São Miguel and three Central Group islands including Pico, Faial and Terceira. Despite large variations in radiogenic isotope signatures in the Azores samples (e.g. 206Pb/204Pb=19.3 to 20.1), δ7Li varies only slightly (+3.1 to +4.7‰), and is within the range for global and North Atlantic MORB (Simons, 2008; Tomascak et al., 2008). Larger variations in δ7Li values such as those reported previously for some EMII, EMI and HIMU ocean island basalts (+3‰ to +7‰; Ryan and Kyle, 2004; Nishio et al., 2005) were not observed in this study. Nevertheless, basalts from the Central Group islands with EM-type isotope signatures are, on average, slightly higher in δ7Li than the São Miguel samples, and they exhibit positive correlations with 87Sr/86Sr and 187Os/188Os and negative correlations with 206Pb/204Pb, 176Hf/177Hf and 143Nd/144Nd. Lithium isotopes do not correlate with indices of fractionation such as MgO, suggesting that the δ7Li correlations with radiogenic isotopes may represent subtle variations in mantle source signatures. The correlations of δ7Li with radiogenic isotopes are consistent with a young recycled metasomatized mantle wedge component with elevated δ7Li beneath the Central Group islands. Lithium isotopes in the São Miguel basalts are essentially homogeneous despite highly variable and radiogenic 87Sr/86Sr, 206Pb/204Pb and 187Os/188Os signatures, which suggests either that their enriched source does not contain typical recycled, altered oceanic crust, or that long storage times of such crust have diffusively homogenized its δ7Li signatures. However, diffusion calculations suggest that heterogeneous Li isotopic compositions of enriched recycled crustal components in the mantle can be maintained over timescales >2.5Ga if the recycled material has a sufficient radius of ~8 km.

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Keywords: Lithium isotopes; Azores Archipelago; ocean island basalts; mantle heterogeneity; crustal recycling; mantle wedge

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1. Introduction With the exception of hydrogen and deuterium, 6Li and 7Li are the two stable isotopes that have the largest relative mass difference, and large isotopic fractionation can be produced in low-temperature environments. Lithium isotope ratios are heterogeneous in crustal materials and altered lithospheric mantle, but relatively homogeneous in the depleted mantle due to limited fractionation at high temperature (Fig. 1). With the development of multi-collector inductively coupled plasma mass spectrometery (MC-ICP-MS), Li isotopic compositions can be analyzed precisely. The δ7Li values of fresh MORB are relatively constant at ~+3-5‰ (Chan et al., 1992; Tomascak, 2008; Elliott et al., 2006). Altered oceanic crust and sediments have high δ7Li values due to interaction with extremely heavy-Li sea water (δ7Li ≌ +31‰; Millot et al., 2004), and terrigenous sediments have lower δ7Li caused by weathering, and their δ7Li values decrease with weathering intensity (Rudnick et al., 2004). The mantle has lower lithium concentrations than crust (Tomascak, 2004), because during mantle melting, lithium is a moderately incompatible element (e.g., Brenan et al., 1998). Therefore, minor recycling of crustal material, whose isotopic signatures are significantly different from the depleted mantle, should modify the lithium isotopic composition of the mantle. Also, some studies have found that during subduction dewatering, residual minerals preferentially retain 6Li, causing the slab-derived fluids to have elevated δ7Li (Chan and Kastner, 2000; Tomascak et al., 2002; Benton et al., 2004; Brooker et al., 2004), and possibly resulting in an isotopically heavy hydrated mantle wedge (Elliott et al., 2006) with a complementary dehydrated slab residue characterized by relatively low δ7Li (Zack et al., 2003; Nishio et al., 2004; Wunder et al., 2006; Simons et al., 2010a). Light Li isotopic signatures have been found in Alpine eclogites (δ7Li=-11 to +5‰) and have been interpreted as the residue of subducted oceanic crust and sediments (Zack et al., 2003). Such fractionations during subduction zone processing have led to the suggestion that the Li isotope system could be a highly sensitive tracer of the residue of recycled and dewatered oceanic crust and sediments, which may be deeply recycled and incorporated into mantle plumes, and which would be isotopically distinct from subduction-modified mantle (Zack et al., 2003; Nishio et al., 2004; Elliott et al., 2006; Wunder et al., 2006).

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However, Marschall et al. (2007) argued that the low δ7Li values of eclogite found by Zack et al. 2003 are more likely explained by kinetic fractionation of the Li isotopes from the country rocks during diffusion rather than subduction dehydration. Modeling results of Marschall et al. (2007) suggest that the dehydration process can only account for a decrease in δ7Li of ≤3‰, which cannot explain the extremely low δ7Li values of eclogite (δ7Li=-11‰, Zack et al., 2003). In a study of Li isotope systematics of lavas from the Cook-Austral islands, Chan et al. (2009) further suggest that even after subduction dehydration, the residue of recycled oceanic crust might still retain a “heavy” Li-isotope fingerprint despite some preferential loss of 7Li, producing HIMU lavas with high rather than low δ7Li (up to 6.2‰). This is consistent with the modeling of Marschall et al. (2007) and more recent modeling by Vlastélic et al. (2009). Experimental results suggest that lithium diffuses rapidly in co-existing phases in peridotites (olivine, clinopyroxene, garnet) and phenocrysts in mafic magmas (clinopyroxene and olivine), and both Li concentration and isotope profiles in crystals suggest that Li equilibration may occur over time scales as short as a few years or less at magmatic temperatures (Jeffcoate et al., 2007; Parkinson et al., 2007). Lithium diffuses far faster than other elements (e.g. Na, K and Rb) in minerals and silicate melts, and the diffusion rates of Li differ between minerals (e.g. Li diffuses faster in clinopyroxene than in olivine; Parkinson et al., 2007). In addition, based on experimental data, the diffusion rates of 6Li and 7Li differ for a given mineral, with 6Li diffusing 3-4% faster than 7Li (Giletti and Shanahan, 1997; Richter et al., 2003; Lundstrom, 2003; Coogan et al., 2005; Parkinson et al., 2007). The rapid and differential diffusion rates of 6Li relative to 7Li lead to strong inter- and intra-mineral Li isotopic fractionation in some mantle xenoliths, and may explain Li isotopic heterogeneity in terrestrial and extraterrestrial igneous rocks from a variety of settings including peridotite and mafic melt (Jeffcoate et al., 2007), arc lavas (Parkinson et al., 2007), mantle xenoliths (Rudnick and Ionov, 2007; Ionov and Seitz, 2008), the Trinity peridotite (Lundstrom et al., 2005), alkaline igneous system (Marks et al., 2007), lunar basalts (Barrat et al., 2005) and phenocrysts from meteorites (Beck et al., 2006). Extreme Li isotopic fractionation (from +7.6 to -19.9‰) has also been found between country rocks and pegmatite (Teng et al., 2006a). These disequilibria are apparently caused by rapid diffusion of Li isotopes at relatively low temperature.

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Therefore, the Li isotopic signal of residual recycled oceanic crust, sediment or slab fluid added to the mantle may not be preserved for geologically significant time periods, but rather may be rapidly homogenized by diffusion. These results challenge the assumption that the Li isotope system would be a sensitive tracer of recycled materials in the mantle. Most of the above studies have focused on mantle peridotites and eclogites. Data for Li isotopes in ocean island basalts (OIB) are still relatively sparse and difficult to interpret. OIB are known to have mantle sources that are highly heterogeneous with respect to radiogenic isotope systems, including proposed enriched end-member compositions HIMU, EMI and EMII, which are generally considered to result from addition of various recycled crustal materials to depleted MORB-type mantle (DMM) (Gast et al., 1964; Zindler and Hart, 1986). Despite substantial variability in radiogenic isotope signatures, most OIB have similar δ7Li values to fresh MORB (+3‰ to +5‰, average is +3.2‰; Tomascak, 2004; Seitz et al., 2007). However, some HIMU and EM-type OIB have been reported to exhibit a larger variation of δ7Li values than MORB (+3‰ to +7‰; Ryan and Kyle, 2004, Nishio et al., 2005; Chan et al., 2009). In addition to these OIB data, Li isotopic data for E-MORB samples from the East Pacific Rise (+3.1‰ to +5.2‰, Elliott et al., 2006; +1.6‰ to +5.5‰, Tomascak et al., 2008) and Mid-Atlantic Ridge (+3.2‰ to 5.7‰, Simons, 2008; Simons et al., 2010b; +2.8‰ to +5.2‰, Tomascak et al., 2008) also have been shown to have slightly anomalous Li isotopic signatures, attributed either to recycled materials in the shallow upper mantle or to seawater contamination during magma storage in shallow crustal level chambers. These results question whether the anomalous Li isotopic signatures of recycled materials are in fact rapidly erased by diffusion. One possibility is that the fractionation observed in experiments between coexisting minerals occurs in nature syn- to post-magmatism instead of in the mantle environment, which could explain the observation that the inter-mineral δ7Li disequilibria in xenoliths are dependent of the type of volcanic rock hosting the xenoliths (Ionov and Seitz, 2008). Whether or not it is possible for Li isotopic signatures of recycled materials to be preserved in the mantle is unclear, and thus the question of the utility of Li isotopes as a tracer of recycled material in the mantle is clearly a complex and unresolved issue. Our study aims to further address the above issues by focusing on Li isotope systematics in the Azores OIB with large Sr-Nd-Pb

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isotope variations that demonstrate significant mantle heterogeneity, and for which recycled material has been proposed.

2. Geological and geochemical setting The Azores Archipelago is located between ~37-40° N, in the vicinity of the Mid-Atlantic Ridge (MAR) and the triple junction between the North American, African and Eurasian plates (Fig. 2). The geochemical characteristic of the tholeiitic basalts from the MAR along the Azores Platform differ from those along normal ridge segments. The Azores Platform tholeiites exhibit enrichment of light rare earth elements (LREE) (Schilling, 1975), high concentrations of large-ion-lithophile elements (LILE), enriched 87Sr/86Sr ratios (White et al., 1976; White and Schilling, 1978), high 206Pb/204Pb and low 143Nd/144Nd isotopic ratios (Dupré and Allègre, 1980; Hamelin et al., 1984; Ito et al., 1987; Dosso et al., 1996), which are interpreted as evidence for an enriched mantle plume beneath the MAR. The Azores Archipelago is a group of nine islands distributed on both sides of the Mid-Atlantic Ridge: Corvo and Flores to the west of the ridge; and Faial, Pico, São Jorge, Graciosa, Terceira, São Miguel and Santa Maria to the east of the ridge (Fig. 2). To the east of the ridge, the Terceira Rift defines the active spreading center, and three islands (Graciosa, São Miguel and Terceira) are aligned along it. The other four islands lie to the south of this rift. The ages of these nine islands generally increase with distance from the ridge, consistent with a mantle plume, although the details are complex and controlled partly by local tectonics (Feraud, et al., 1980). The oldest rocks are found in Santa Maria (8Ma) and São Miguel (4Ma) islands (Abdel-Monem et al., 1975; Fernandez, 1980; Feraud et al., 1980; Moore, 1990; Moreira et al., 1999), but recent and historic volcanism has occurred on many of the islands. The most recent volcanism was the submarine eruption of Seretta, 8.5 km west of Terceira between 1998-2000 (Forjaz et al., 2001). The lava compositions of the Azores Archipelago range from basalts to trachytes, and some peridotite nodules are found in basalts from Pico, Faial and Terceira islands (Mitchell-Thome, 1976; França et al., 1995). Numerous studies suggest that the Azores region relates to ridge-hot spot interaction (Gente et al., 2003), based on the anomalously thick crust at the spreading ridge (Vogt,

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1976; Schilling, 1985; Gente, 1987; Thibaud et al., 1998; Beier et al., 2010), the enriched basalt geochemistry (Schilling, 1975; White et al., 1976; Bougault and Treuil, 1980; Schilling et al., 1983; Dosso et al., 1999; Moreira et al., 1999; Jean-Baptiste et al., 2009; Millet et al., 2009), and gravity anomalies (Detrick et al., 1995; Thibaud et al., 1998). The gravity and bathymetric data combined with seafloor magnetic anomalies indicate that the Azores Platform was formed by ridge-hotspot interaction between ~20-7Ma, and that the rift in the plateau was caused by migration of the MAR away from the hotspot (Gente et al., 2003). Based on schematic reconstructions of the relative movement between the Mid-Atlantic Ridge and a hot spot, Gente et al. (2003) suggest that the hotspot is currently underlying Terceira Island. However, the depth of origin of the mantle plume is still highly controversial, and several seismic tomography studies of the Azores region have led to a variety of interpretations (Montelli et al., 2004, 2006; Silveira et al., 2006; Yang et al., 2006). The Azores mantle plume was proposed to be one of only six mantle plumes globally that originate at the core-mantle boundary based on images from both P-wave (Monteli et al., 2004) and S-wave (Montelli et al., 2006) velocities. However, Silveira et al. (2006) argued based on another S-wave velocity study that the Azores mantle plume is restricted to the upper mantle. Yang et al. (2006) suggest further, based on P-wave velocity modeling derived from teleseismic body waves, that the hotspot underlying Terceira Island is a cylindrical, low-velocity volume at 250-400 km depth centered northeast of Terceira, and is deflected beneath Terceira by shallow regional upper mantle shear. The alkali basalts of the Azores Archipelago exhibit extreme heterogeneity on both archipelago-wide and intra-island scales based on Sr-Pb isotopes (Fig. 3), and our new isotopic data are consistent with previous work (Widom et al., 1992, 1997; Moreira et al., 1999; Beier et al., 2007, 2008; Elliott et al., 2007; Millet et al., 2009). Sr-Nd-Pb isotopic data indicate that the mantle source of the Azores basalts include DMM with HIMU and EM components (Turner et al., 1997; Widom et al., 1997). However, the origins of these HIMU and EM components are still not clear. Previous workers proposed that the enriched components of the São Miguel mantle source are due to subducted terrigenous sediments (Hawkesworth et al., 1979; Turner et al., 1997), delaminated subcontinental lithosphere from the opening of the North Atlantic (Davies et al., 1989; Widom et al.,

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1997; Moreira et al., 1999), recycled oceanic crust with evolved compositions (such as a subducted seamount; Beier et al., 2007), or enriched (E-MORB type) under-plated basalt that infiltrated the oceanic mantle lithosphere (Elliott et al., 2007). For the other islands from the Central Group (Faial, Pico, São Jorge and Terceira), recycled oceanic crust (Prytulak and Elliott, 2009) or recycled Archaean oceanic mantle lithosphere (Schaefer et al., 2002; Turner et al., 2007; Jean-Baptiste et al., 2009) in their mantle source, or shallow assimilation of delaminated SCLM have been proposed to explain the isotopic variations (e.g. Faial and some of São Jorge samples; Millet et al., 2009). So far, no Li isotope data have been published for basalts from the Azores Archipelago, although Li isotopes in nearby Mid-Atlantic Ridge (MAR) basalts have been published in previous studies. Tomascak et al. (2008) found that the mean values of Li isotopes in MAR basalts overlap the N-MORB range within uncertainty, and three of these samples are from the Azores Plateau (+2.2 to +4.5‰). In this study, there was lack of systematic co-variation between Li isotopes and elemental or radiogenic isotopic parameters. Another study by Simons (2008) also found that fresh glass samples from the Azores Platform lavas have a small range in δ7Li (+3 to +5‰, similar to fresh MORB), 7 but showed that δ Li signatures correlate with trace element ratios (K/Ti, H2O/Ce) and radiogenic isotopes (87Sr/86Sr and 143Nd/144Nd), which is interpreted to reflect mantle source heterogeneity beneath the MAR. In this study, we investigate the Li isotope signatures of the Azores island basalts to further investigate the extent and potential causes of Li isotopic heterogeneity (or lack thereof), and to potentially help constrain the origins of the mantle sources and magmatic processes beneath the Azores archipelago region.

3. Samples and analytical methods In this study, we selected 11 young, fresh basalt samples including four from São Miguel (WASM series), and seven from the Central Group islands of Terceira, Faial and Pico (WAT, WAF and WAP series, respectively) for Li isotopic analysis. Basalts from Pico, São Miguel and one basalt from Faial (WAF1a) are olivine rich (>10% olivine); WASM5, WASM17 and WAP8 are also clinopyroxene (cpx) rich (>20% cpx). WAF30

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and basalts from Terceira have sparse olivine and cpx (<5% olivine and cpx). Feldspar is found in all samples as microphenocrysts (~50 to 250μm). We collected the freshest possible samples in the field and removed all possibly weathered surfaces in the laboratory. Whole rock samples were sawed into 0.5 cm thick slabs, and all surfaces were ground using Si-C paper to remove any metallic residue from the saw blade. Sample slabs were then cleaned in 18.2 mega-Ω E-pure water, including multiple sessions in the ultrasonic bath with rinsing in between. After drying in an oven (110℃), the slabs were wrapped in paper and plastic, and chipped by hammer to 0.5-1.0 cm fragments. Samples were then chipped in a Sepor alumina jaw crusher, and powdered in a Spex shatterbox using a high purity alumina vessel. MgO-rich samples were selected to represent the observed range of variations of Sr and Pb isotopic signatures on each island. Major and trace element analysis, chemical separations and Sr, Nd and Pb isotopic measurements were done at Miami University. Major and some trace element concentrations (Sc, Ba and Cr) of whole rocks were obtained by Beckman SpectraSpan V Direct-Current Plasma Atomic Emission Spectrometer (DCP-AES) following the procedures of Katoh et al. (1999).

Approximately 200 mg of whole rock powder was mixed with ~600 mg Li2B4O7 flux and dissolved in 5% HNO3 after fusing at 950 ˚C for 20 minutes. Ten international rock standards and one blank were used as external standards for calibration. The other trace element concentrations were obtained by inductively coupled plasma mass spectrometer (ICP-MS) using a Varian quadrupole instrument with the exception of Pb concentration. Approximately 50 mg of whole rock powder was mixed well with ~75 mg of Na-K flux (Na-tetraborate and K-carbonate from Alfa Aesar-Puratronic grade in a 3:2 mixture) and fused in a furnace for 30 minutes at 950 ˚C, and then dissolved in 2% HNO3 and analyzed with ten external standards and one blank. For Pb concentration analysis, samples were dissolved by acid digestion using high purity, concentrated HF-HNO3. Approximately 50 mg of whole rock powder was dissolved with concentrated HF-HNO3, and following dry-down the samples were dissolved in 2% HNO3. Pb concentrations were determined by the standard additions method, in which each sample solution was analyzed with two standard addition solutions and a blank. Methods for chemical separation and analysis of radiogenic isotopes by TIMS were described in detail in Chapter 2. The Sr, Nd and Pb 94

analysis followed the procedures of Snyder (2005) and references therein, Re-Os isotope analysis followed the procedures of Shirey and Walker (1995) and Nägler and Frei (1997), and Hf isotope analysis followed the procedures of Connelly et al. (2006). Lithium isotopes were analyzed at the Geochemistry Laboratory of the University of Maryland, College Park, following the procedures of Teng et al. (2006b). Lithium isotopes were analyzed on whole rock powder for all samples, and four samples were selected for additional analysis of Li isotopes in olivine separates. The olivine crystals from each of these samples were handpicked under a binocular microscope, and inclusions and adhering glass were avoided. Approximately 25mg of whole rock powder or ~50mg of olivine separates were dissolved by a ~3:1 mixture of concentrated HF and

HNO3, covered and heated by hotplate (T<120℃) overnight until the samples were largely dissolved (minor residue only). After drying down, the sample residues were re-dissolved overnight in ~5ml concentrated HNO3, and dried down. Then, the samples were heated in concentrated HCl on the hotplate until the samples were completely dissolved and the solutions appeared completely clear, generally one to two days. After being dried down again and re-dissolved in appropriate amounts of 4N HCl, the samples were ready for chromatographic separation. Before samples were loaded onto the columns, they were put in the ultrasonic bath for 10 minutes to ensure that all of the samples were fully dissolved, and then centrifuged. The sample solutions, with ~50ng Li in 1ml of 4M HCl, were then loaded onto the first separation column. Lithium was eluted through three sets of columns, each of which contained 1ml of cation-exchange resin (BioRad AG50W×12). The procedures for separations and purification of Li on each of the three columns are based on methods of Morguti and Nakamura (1998a), and effectively separate Li from other elements with >98% yield (Teng et al., 2006b). The first column was conditioned with 1ml of 4N HCl, and the sample dissolved in 1ml of 4N HCl and loaded into the column, and the Li collected in 9ml of 2.5N HCl. After drying down, the collected Li was redissolved in 1.5ml of 0.15N HCl and loaded onto the second column after the column equilibrated with 1ml MQ H2O, and the Li collected in 30ml of

0.15N HCl. The third column was conditioned with 1ml MQ H2O, and the sample loaded in 1ml 0.15N HCl, and then Li collected in 16 ml of 30 vol.% ethanol in 0.5N HCl. If

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necessary (see below), the sample solutions were passed through the third column twice to fully separate Na from Li. Samples were analyzed on a Nu-Plasma Multi-Collector-Inductively Coupled Plasma-Mass Spectrometer (MC-ICP-MS) (Belshaw et al., 1998), following the procedures of Teng et al. (2004). Before Li analysis, the Na/Li signal ratio of each solution was evaluated by mass spectrometery. If the Na/Li signal ratio was ≥5, the solution was passed through the third column again until the Na/Li signal ratio was <5. Purified Li solutions were introduced into the Ar plasma by an auto-sampler (ASX-100 Cetac Technologies) and through a desolvating nebulizer (Aridus Cetac Technologies) with a PFA spray chamber and micro-nebulizer (Elemental Scientific Inc.). 7Li and 6Li were measured statically using two Faraday cups. Before and after each sample, the L-SVEC standard (Flesch et al., 1973), with a Li concentration of 50ng/ml and close to those of the sample solutions, was analyzed. At least two other Li standards were analyzed during the course of an analytical session, including the in-house standard Li-UMD-1 (a purified Li solution from Alfa Aesar) and IRMM-016 (Qi et al., 1997). The isotopic data are reported relative to the L-SVEC standard as δ7Li 7 7 6 7 6 (δ Li=[( Li/ Li)Sample/( Li/ Li)L-SVEC-1]×1000) with an uncertainty estimated to be ±0.6‰ (see below). The uncertainty of the Li concentration measurements was <±10% (1ζ) (Teng et al., 2007).

4. Results In this study, whole rock powder and olivine crystals were divided into three groups and analyzed using MC-ICP-MS over the course of three days. Whole rock powders for 11 samples were analyzed on day one (sample group 1), and values for standards of this analytical session (L-SVEC 50ppb) are reported in Fig. 4. Deviations from the mean of replicate analyses of all standards plotted within a range of ±0.6‰, with the exception of the first standard (Fig. 4). The standards run during the sample analysis period (labeled in Fig. 4) have an even narrower deviation about the mean of replicate analyses (±0.3‰), indicating good instrumental stability. Based on long-term deviations (2ζ) of Li standard solutions and rock solutions over a couple years, the external precision of this instrument has been estimated to be ≤±1.0‰ to ±0.6‰ (Teng et al., 2007). In this study we use

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±0.6‰ as the uncertainty represented by the sample error bars, given the demonstrated reproducibility for standard analyses during our analytical time period. Lithium concentrations and isotopic compositions of whole rock samples are reported in Table 1 with major element, trace element, and Sr, Nd, Pb, Hf and Os isotopic compositions. We also investigated potential Li isotopic fractionation between whole rock and olivine separates. Sample group 2 included four pairs of olivine and associated whole rock powder, one pair each from samples WAP8, WAP9, WAP10 and WASM24. The Li isotopes of these four pairs were analyzed by MC-ICP-MS on day two. For three pairs, the Li isotopic data of whole rock powder and respective olivine separates agreed very well within ±0.6‰ (WASM24, WAP9 and WAP10; Table 2 and Fig. 5). However, the Li isotopic ratio of the olivine separate from WAP8 was 3‰ higher than the respective whole rock sample (Fig. 5, WAP8-1 and WAP8-1replicate). In this analysis, the Li concentration of the WAP8 whole rock powder was higher than that determined for this same sample when run with sample group 1 as well as that obtained by quadrupole ICP-MS (>±20%), indicating there may have been a large fractionation between 6Li and 7Li during sample processing, or that WAP8 may have been contaminated during dissolution or chemical separation prior to the sample group 2 analysis. Therefore, we analyzed the WAP8 whole rock powder and olivine separate again in the sample group 3 (WAP8-2 and WAP8-2r). In this analysis, the Li isotopic data for the whole rock powder and the olivine separate agreed with each other within ±0.35‰ (Fig. 5), and the Li concentration of the whole rock powder was consistent with the data from the group 1 analysis and the ICP-MS results. These results suggest that the data from the group 3 analysis of WAP8 (whole rock powder and olivine separates) are more reliable than the group 2 analysis. Lithium abundances in the selected Azores samples have a relatively small variation (3-6 ppm), and are in the range of other oceanic island basalts (3-19 ppm; Ryan and Kyle, 2004). The δ7Li values of these Azores basalts have restricted variability (+3.0‰ and +4.7‰). Fig. 6 shows that the Li isotopic values are similar to those of global fresh MORB (+3‰ to +5‰, average +3.4‰; Chan et al., 1992; Tomascak, 2004, 2008; Elliott et al., 2006). Although within analytical uncertainty, the mean δ7Li value of the Central Group islands (+4.14‰) is slightly higher than that of the São Miguel samples (+3.46‰)

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(Fig. 6a). Similarly, there are also slight differences (albeit within analytical uncertainty) in the means between the Central Group islands, decreasing from Faial (+4.55‰) to Pico (+4.05‰) to Terceira samples (+3.87‰). The slight but systemic differences of mean δ7Li values between the Central Group islands correlate with other radiogenic isotopes (Fig. 7 and 8), including negative correlations between δ7Li and 206Pb/204Pb, 143Nd/144Nd and 176Hf/177Hf, and positive correlations with 87Sr/86Sr and 187Os/188Os. In contrast to the Central Group island samples, the δ7Li values of the São Miguel samples do not exhibit any systematic correlation with other radiogenic isotopes (Fig. 7 and 8). Although these São Miguel samples exhibit extreme variations in radiogenic isotopic signatures, their Li concentrations and isotopic values are within a narrow range. Based on the behavior of Li isotopes with respect to radiogenic isotopes, the samples were divided into two sample groups for further discussion: (1) the Central Group islands, whose δ7Li values correlate with radiogenic isotopes (Fig. 7 and Fig. 8) and which are characterized by sub-chondritic 187Os/188Os ratios (Fig. 9); and (2) São Miguel basalts, for which δ7Li values are not correlated with any radiogenic isotopes (Fig. 8) and which exhibit a large variation of 187Os/188Os (0.127-0.157) including very radiogenic signatures in samples with relatively high Os concentration (44-214 ppt; Fig. 9). These two sample groups have also been demonstrated to have distinct major and trace elements characteristics (Fig. 10), as described in Chapter 2.

5. Discussion 5.1 Weathering, seawater alteration and shallow-level contamination The results of this study do not reveal any anomalously low or elevated δ7Li compared to the range of global fresh MORB (+3‰ to +5‰). However, even though the Li isotope signatures of the Central Group islands have limited variation, they exhibit small but systematic differences between each island (within analytical uncertainty), that co-vary with radiogenic isotopes (Fig. 7 and Fig. 8). Before speculating whether this variation might reflect the signatures of the respective mantle source compositions, it is important to consider the potential effects of weathering, low temperature alteration, and assimilation on the isotopic composition of Li in the basalt samples. Low temperature seawater alteration would generate secondary minerals with seawater-derived high Li

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concentrations and high δ7Li (Chan et al., 1992; Chan et al., 2002), and assimilation of this material could increase Li concentrations and isotopic ratios in basalts. In contrast to seawater alteration, surficial weathering processes have been demonstrated to decrease δ7Li in silicate rocks. Based on the studies of Hawaiian soils produced by weathered lava flows (Huh et al., 2004), two saprolites from South Carolina (Rudnick et al., 2004) and natural basalt and granite weathering products (Pistiner and Henderson, 2003), it has been established that both Li concentrations and δ7Li values in the rocks will decrease with increasing weathering intensity (δ7Li to ~-20‰). All of the samples in this study are young and petrographically “fresh” with no evidence of secondary mineral phases, and all have very low loss on ignition values (LOIs from -0.51% to 0.04%) in which the LOI is the percentage of sample lost after heating at 950 ˚C for two hours. There is no correlation between LOI and measured δ7Li (Fig. 10c), indicating that the Li isotopic compositions of these samples are not controlled primarily by weathering. The Li concentrations and isotopic values of the Azores basalts plot within the OIB range, and the lack of any low δ7Li values (<0‰) also indicates that the samples were not significantly altered by weathering. However, it is critical to evaluate whether minor weathering has affected the Li isotopic compositions of the basalt samples. Elements that are commonly affected by weathering such as K and Rb exhibit well correlated trends with immobile elements such as Nb, despite the fact that they are highly mobile during basalt weathering, further suggesting that weathering has not significantly affected the Central Group island basalts (Fig. 11a, data from Yu et al., unpublished). On plots of K2O versus Nb or MgO (Fig. 11), Faial and Pico samples follow a very tight trend (R2=0.91), distinct from many of the Terceira samples that also follow relatively well defined trends of K2O versus Nb or MgO but generally at lower values of K2O. These relationships are consistent with K2O variability due to different source characteristics or magmatic fractionation paths rather than the result of weathering alteration. Beryllium is also a useful tracer for assessing effects of weathering or late-stage alteration processes. Like Nb, Be is a highly incompatible trace element that is immobile during weathering or hydrothermal processes, and variations in ratios of highly mobile elements (e.g. K2O, Rb) with Be could reflect the effect of weathering or hydrothermal exchange processes (Ryan, 2002; Vesely et al., 2002; Ryan and Kyle,

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2004). Fig. 11b shows that the K2O/Be ratios of the Central Group islands are positively correlated with Rb/Be ratios (R2=0.898), further supporting the contention that the Azores basalt samples in this study are not significantly affected by weathering or hydrothermal processes. Interaction of oceanic basalts with seawater could also potentially cause variability in δ7Li. Low temperature alteration of basalt by seawater would form secondary clay minerals that incorporate Li from seawater (Chan et al., 1992, 2002, 2009; Bouman et al., 2004), producing heavy isotopic compositions (up to +15‰) relative to fresh MORB, due to the heavier Li isotopic composition of seawater (δ7Li=~+31‰; Millot et al., 2004). The Li concentrations of our samples are in the range of unaltered MORB (Tomascak, 2004), indicating that low-temperature seawater alteration did not significantly affect these samples. However, within the very narrow range of Li isotopic compositions and Li concentrations, there is a positive trend between δ7Li and Li concentration (Fig. 10e). This correlation might be caused by minor seawater alteration or by assimilation of a seawater altered component during magma ascent to the surface. K/Rb ratio is sensitive to seawater alteration (Hart, 1969; Verma, 1992) because seawater exchange results in enrichments of K and Rb of 2 and 5 times respectively, and thus a decrease in K/Rb (<380) (Fig. 10f). However, lack of negative correlation between K/Rb ratios and either δ7Li or Li concentration in our sample suite argues against the effect of seawater alteration. Finally, olivine phenocrysts are substantially less susceptible to low-temperature alteration or weathering than the matrix of the whole rocks (Chan et al., 2009), thus the agreement in this study between the δ7Li signatures in olivine phenocrysts and their respective whole rocks further argue that neither low-temperature alteration nor weathering affected the Li isotopic signatures of the Azores basalts. Altered oceanic crust and pelagic sediment exhibit large Li isotopic variations due to seawater interaction (~-2‰ to ~+20‰, Fig. 1; Elliott et al., 2004). Minor shallow-level contamination could therefore also impact Li isotopic signatures of basaltic magmas. 7 Lack of correlation between δ Li and indices of fractionation (e.g. MgO and K2O) argue against an extensive role for crustal assimilation or assimilation-fractional crystallization (AFC) processes that might affect the δ7Li of Azores samples (Fig. 10). However, the positive correlation of δ7Li and Li concentration in the Central Group island basalts

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require that the potential effect of crustal assimilation be evaluated more thoroughly. The Os isotopic system is a highly sensitive tracer of crustal assimilation, and can provide additional constraints on the potential role of crustal assimilation in the petrogenesis of the Azores basalts. The Central Group island basalts have sub-chondritic 187Os/188Os with variable Os concentrations ranging from 23 to 140 ppt, with the exception of one sample that has supra-chondritic 187Os/188Os (0.134) but an Os concentration of only 5 ppt. Fig. 12 shows the possible mixing trends that would be imparted by shallow-level crustal contamination. The basalts in these models have a range of Os concentrations consistent with those observed in the Central Group island samples, including the lowest (5 ppt), intermediate (30 ppt) and the highest (140 ppt) Os concentrations. Fig. 12a shows that if ocean island basalts with Os concentrations of 30 ppt to 140 ppt experienced 2% shallow-level contamination by marine sediments with radiogenic 187Os/188Os (~1) and extremely elevated δ7Li (+15‰; Bottomley et al., 1999; Chan et al., 2006), the positive correlation between δ7Li and 187Os/188Os in the Central Group island samples could be produced. The Central Group island basalts also have slightly elevated 87Sr/86Sr relative to depleted mantle, and the positive correlation of δ7Li with 87Sr/86Sr (Fig. 7b) could also potentially be explained this way. However, Fig. 12b shows that shallow-level contamination produces mixing curves in Sr-Os space much steeper than that observed in the Central Group island samples, and with the exception of the one very low Os abundance sample with supra-chondritic Os, the Sr-Os isotope relationships cannot be attributed to sediment assimilation.

5.2 Partial melting and fractional crystallization Wide ranges of major and trace element concentrations and associated geochemical correlations amongst the Azores samples indicate an important role for fractional crystallization and possibly variable degrees of partial melting. Most of the Azores samples contain abundant olivine and clinopyroxene phenocrysts, and observed major element systematics suggest an important role for fractionation of these phases. Variable degrees of partial melting in the Azores have also been proposed, based on rare earth element systematics. Lower Dy/Yb (Fig. 10h), Ce/Yb ratios and TiO2 in the Central Group island basalts compared to São Miguel basalts have been suggested to reflect a

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higher degree of partial melting of mantle peridotite (3-5% vs. 1-2%) in the generation of the Central Group island basalts relative to São Miguel (Bourdon et al., 2005; Beier et al., 2008). It is therefore important to evaluate the potential of Li isotopic fractionation during partial melting and fractional crystallization. Previous studies of basalts from Kilauea Iki lava lake of Kilauea volcano, Hawaii showed that there is no measureable Li isotopic fractionation at magmatic temperatures (Tomascak et al., 1999). Therefore, we do not expect that variable degrees of partial melting or magma differentiation processes should affect the Li isotope composition of magma significantly. The Central Group island basalts, which exhibit correlations between δ7Li and radiogenic isotopes, do not show obvious co-variation between δ7Li and indicators of degree of partial melting such as Dy/Yb (Fig. 10h), suggesting that variable degrees of partial melting did not cause any systematic Li isotope fractionation in the Central Group islands. Additionally, the overlap in Li isotopic composition between olivine phenocrysts and their respective whole rocks (Fig. 5), and the lack of correlation of δ7Li with indices of fractionation such as MgO wt.%, SiO2 wt.% and K2O/TiO2, indicates that crystal fractionation did not induce fractionation of Li isotopes in the Azores basalts either (Tomascak et al., 2008). Therefore, we consider in the following sections the possibility that the minor variations in δ7Li observed in the Central Group island samples reflect heterogeneous mantle sources.

5.3 Origin of Li isotope variation in the source of the Central Group islands Among the Central Group islands, both intra- and inter-island radiogenic isotope heterogeneities are observed (Fig. 7, 8 and 9). The diverse signatures of radiogenic isotopes and major and trace elements amongst the Azores basalts indicate that there are several distinct components in their mantle source. Basalts from Terceira have radiogenic 206Pb/204Pb and display a HIMU/FOZO-like signature, Faial basalts have elevated 87Sr/86Sr characteristic of an EMI/EMII signature, and Pico basalts show a mixing trend between these HIMU/FOZO and EM components. Although the δ7Li variation of the Central Group island basalts is limited, co-variations of δ7Li with radiogenic isotopes (Sr, Nd, Pb, and Hf) are observed (Fig. 7 and 8) and may reflect Li isotope variations of their

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mantle sources. In the following discussion, we evaluate the implications of these Li isotopic signatures and their correlations with radiogenic isotopes regarding the nature of the mantle sources beneath the Azores.

5.3.1 The mantle source of HIMU/FOZO type Terceira basalts HIMU-type basalts have most commonly been attributed to a source containing ancient subducted oceanic crust that developed high U/Pb and Th/Pb ratios during subduction dewatering and loss of soluble Pb (Zindler et al., 1982; Weaver, 1991; Chauvel et al., 1992; Hofmann, 1997; Stracke et al., 2003). In previous Li isotope studies of OIB, HIMU type basalts have generally been found to have higher δ7Li values than MORB. For example, elevated δ7Li values were found on St. Helena island with δ7Li extending to +7.0‰ (Ryan and Kyle, 2004), and values up to +7.4‰ for basalts from Mangaia and Tubuai (Nishio et al., 2005). Rurutu Island in the Austral Chain has +5.4<δ7Li<+7.9‰ (Vlastélic et al., 2009), and olivine separates from Cook-Austral lavas have high δ7Li signatures up to +6.2‰ (Chan et al., 2009). Although Zack et al. (2003) proposed that the residue of recycled altered oceanic crust should have low δ7Li due to preferentially scavenging of 7Li with respect to 6Li by slab fluids during subduction dehydration, other studies have argued that the dewatered slab can retain high δ7Li imparted during seawater alteration, potentially explaining elevated δ7Li signatures in HIMU OIB that are believed to contain a component of recycled oceanic crust in their mantle source (Chan et al., 2009; Vlastélic et al., 2009). In contrast, our study did not reveal δ7Li signatures higher than MORB in the Terceira basalts, despite them sharing the relatively radiogenic 206Pb/204Pb characteristics of some of the other HIMU samples discussed above. However, Li isotope values within the range of MORB and similar to Terceira basalts have also been found in the Jan Mayen ankaramitic lavas (Magna et al., 2011) and the “non-HIMU” (19.11<206Pb/204Pb<20.45) lavas from Rurutu Island (Vlastélic et al., 2009), for which the mantle source origins remain unclear. Contrary to previous studies of HIMU basalts, which display positive correlations of δ7Li with 206Pb/204Pb, the Central Group island samples exhibit a negative correlation (Fig. 7a), and Terceira basalts have systematically lower δ7Li and higher 206Pb/204Pb than Faial and Pico basalts. One possibility is that the mantle source of the Terceira basalts

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might not contain recycled altered oceanic crust, a conclusion supported previously based on radiogenic isotopes (Sr, Pb, Nd, Hf and Os) and trace element systematic as discussed in Chapter 2. Most HIMU-type lavas with highly radiogenic Pb isotopes have supra-chondritic Os isotopes, indicating recycled oceanic crust and sediment in their mantle source (Hauri and Hart, 1993; Schiano et al., 2001; Lassiter et al., 2003). However, sub-chondritic 187Os/188Os signatures (Fig. 9) and lack of depletion of fluid mobile elements may argue against this model for the Terceira source. For example, Terceira basalts have higher K/U ratios (9000-15000, with the exception of WAT11, Fig. 11e), higher LILE/HFSE ratios (e.g. Rb/Nb and Ba/Nb) and higher LREE/HFSE ratios (e.g. La/Nb; Fig. 11f) than other HIMU lavas, and their K/U ratios do not correlate with Pb isotopes. These geochemical features suggest that the enriched component in the Terceira mantle source did not contain recycled crustal material that underwent significant subduction dehydration and associated loss of fluid mobile elements. 3 4 Terceira basalts have higher He/ He values (up to 13.5Ra) than MORB (8±1Ra) (Moreira et al., 1999; Jean-Baptiste et al., 2009), indicating the presence of a mantle plume beneath Terceira containing a relatively undegassed reservoir from the deep mantle. This is supported by elevated 20Ne/22Ne ratios (Madureira et al., 2005). It is possible that Terceira basalts reflect derivation from a FOZO plume with lower 206, 207, 208Pb/204Pb and higher 87Sr/86Sr compared to other HIMU sources (Fig. 2; Stracke et al., 2005). So far, no studies have discussed the Li isotope signature of the FOZO component, but some of the OIB from the literature for which Li isotope data exist might also be argued to have a FOZO source. For example, basalts from Rurutu island with 19.11<206Pb/204Pb<20.45 are FOZO-like and also have Li isotopic signatures (+2.9<δ7Li<4.8‰) similar to those of Terceira basalts.

5.3.2 The mantle source of EM type Faial basalts Faial samples have lower 206Pb/204Pb, 143Nd/144Nd, 176Hf/177Hf and higher 87Sr/86Sr than the other islands in the Central Group, and reflect an EM-type signature, although less extreme than the end-member EMI or EMII compositions. Similar to end-member EMI and EMII islands, but distinct from Terceira, the Faial samples also have 207Pb/204Pb-206Pb/204Pb and 208Pb/204Pb-206Pb/204Pb compositions that lie to the left of

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NHRL. Here, we evaluate the Li-isotope systematics of the Faial basalts in the context of potential enriched components that may contribute to their mantle source. In previous studies, some EM-type OIB have a MORB-like compositional range or slightly elevated δ7Li (+3.3 to +5.9‰ of Sikhote-Alin basalts and Rapa basalts; Chan et al., 2009). In this study, no anomalous δ7Li signatures were found in the Faial basalts; rather, the δ7Li signatures are slightly elevated compared to the Terceira basalts but within the range of fresh MORB. Nishio et al. (2004) proposed that the enriched mantle component of the EM-type OIB with δ7Li<-17‰ is the residue of subducted oceanic crust that preferentially lost 7Li during dehydration at low temperature, but models for the generation of OIB mantle sources for basalts with MORB-like or slightly elevated δ7Li signatures remain unclear (Nishio et al., 2004; Chan et al., 2009). The EM mantle end members have most commonly been attributed to incorporation of pelagic and terrigenous sediments subducted with recycled oceanic crust (Zindler and Hart, 1986 and reference therein; Weaver, 1991; Woodhead and Devey, 1993). However, seawater derived sediments have isotopically high δ7Li (Millot et al., 2004) and terrigenous sediments have low δ7Li caused by weathering (Rudnick et al., 2004). If pelagic sediment is considered to have a mean composition of [Li]=30 ppm and δ7Li=+8‰ (Chan et al., 2006) and terrigenous sediment a mean composition of [Li]=35 ppm and δ7Li=0‰ (the average composition of upper continental crust, Teng et al., 2004), addition of 2% sediment to a MORB source with [Li]=1.5 ppm and δ7Li=+3.2‰ (Jagoutz et al., 1979; Seitz et al., 2007) would increase δ7Li from +3.2‰ to +4.6‰ or decrease δ7Li from +3.2‰ to +2.2‰, respectively. Therefore, <2% addition of marine sediment to depleted Mid-Atlantic ridge-type mantle would not disturb the Li isotope signature in the mantle source, but 2% addition of terrigenous sediment would decrease the δ7Li signature in the mantle source to values lower than the range of fresh MORB. In addition, if Mid-Atlantic ridge depleted mantle and present-day pelagic and terrigenous sediments have [Sr]=10, 600, 300 ppm, 87Sr/86Sr=0.703, 0.709 and 0.715, [Pb]=0.02, 20, 10 ppm and 206Pb/204Pb=19.1, 18.9 and 18.9 respectively (Weaver, 1991; Hart et al., 1992; Plank and Langmuir, 1998; Widom et al., 1997), addition of <2% recycled pelagic sediment or terrigenous sediment would produce basalts with variable 87Sr/86Sr and 206Pb/204Pb. For example, mixing with 2% of 0.5Ga pelagic or terrigenous sediment (whose isotopic

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compositions can be calculated based on their current compositions, element fluid mobilities during subduction, and time since subduction), would produce basalts with 87Sr/86Sr of 0.703-0.707 and 0.703-0.708 and 206Pb/204Pb of 19.1-19.0 and 19.1-20.0. Because addition of pelagic sediment will increase 87Sr/86Sr but decrease 206Pb/204Pb, Faial basalts with 206Pb/204Pb up to 19.5 cannot be explained by pelagic sediment addition to the mantle source. A small amount of terrigenous sediment (~1%) added to the mantle source can produce basalts with variable 87Sr/86Sr and 206Pb/204Pb in the range of Faial basalts. However, the addition of terrigenous sediment with [Nd]= 12.64 ppm, 143Nd/144Nd= 0.51132, [Hf]= 2.81 ppm and 176Hf/177Hf= 0.281471 to depleted mantle with [Nd]= 7.45 ppm, 143Nd/144Nd= 0.5113088, [Hf]= 0.31 ppm and 176Hf/177Hf= 0.283011 (Vervoort et al., 1999; Stracke et al., 2003; Chauvel et al., 2008) would produce a trend in Nd-Hf space that is steeper than the slope exhibited by Faial basalts. For example, mixing 1% 1Ga terrigenous sediment will produce basalts with Nd of 8.2 and

Hf 6.7, which is below the global MORB-OIB array and the Central Group array (Chapter 2). Therefore, sediment addition to the Faial mantle source appears an unlikely explanation for the Li isotope systematics. Alternative potential mechanisms to produce EM type mantle sources with a MORB range of δ7Li include recycling or in situ assimilation of oceanic lithospheric mantle, or recycling of subduction-modified mantle wedge. Recycled oceanic lithospheric mantle (OLM) is a potential enriched source for Faial basalts. Hydrothermally altered OLM would be expected to have high Li concentrations and elevated δ7Li (up to +15‰) due to interaction with isotopically heavy seawater (Fig. 1; Decitre et al., 2002; Elliott et al., 2004). A small addition of this OLM could slightly increase δ7Li, but leave it within the MORB range. However, OLM typically has sub-chondritic Os isotopes due to long-term Re depletion (Schaefer et al., 2002), hence cannot explain the positive correlation between δ7Li and Os isotopes. Because Hf is a fluid-immobile element, reaction with seawater or metasomatic fluids (Weaver, 1991; Pearce et al., 1999; Jicha et al., 2004; John et al., 2004) would not significantly alter the Hf isotope signature of the OLM, thus there would be no expected correlation between δ7Li and 176Hf/177Hf. The negative correlation between δ7Li and 176Hf/177Hf in this study (Fig. 7d) therefore suggests that

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altered OLM is an unlikely source of Faial basalts. The in situ assimilation by OLM is not a likely cause of the EM signature of Faial basalts for the same reasons. Previous studies have suggested that subduction-modified sub-arc mantle might be recycled into the deeper mantle via viscous coupling to a subducted slab, and that this could be the “enriched” component in the source of EM-type basalts (Eiler et al., 1997; Lassiter et al., 2003). Since 7Li is preferentially transferred to the overlying mantle wedge with slab-derived fluids relative to 6Li, a metasomatized mantle wedge would be expected to have elevated δ7Li, with enriched trace element and radiogenic isotope signatures imparted by slab-derived fluid or melt (Chan and Kastner, 2000; Tomascak et al., 2002; Benton et al., 2004; Brooker et al., 2004; Simons et al., 2010a). Addition of 3% subduction-related mantle wedge with δ7Li=+15‰ and [Li]=6 ppm (data from olivine phenocrysts of arc lavas; Parkinson et al., 2007) to depleted mantle with δ7Li=+3.2‰ (Seitz et al., 2007) and [Li]=1.5 ppm (Jagoutz et al., 1979) could increase the δ7Li of the mixture from +3.2‰ to +4.5‰, which is within the range of the Faial samples. Additionally, by reacting with slab-derived fluid, the mantle wedge can acquire radiogenic Sr and potentially even supra-chondritic 187Os/188Os (Brandon et al., 1996; McInnes et al., 1999; Widom et al., 2003). Mixing 3% recycled mantle wedge with 187Os/188Os=0.157 and [Os]=8.2 ppb (Widom et al., 2003) into depleted mantle with [Os]=3.1 ppb (Morgan, 1986) would increase the 187Os/188Os slightly from 0.122 to 0.125. The addition of recycled metasomatized mantle wedge material can therefore explain the MORB-like range of δ7Li isotope range of in the Faial basalts as well as the observed positive correlations with Sr and Os isotopes (Fig. 9). Pico basalts exhibit mixing trends in radiogenic isotope systems between the enriched components in the mantle sources of both Faial and Terceira. The “Terceira” end-member of the Pico basalts is characterized by higher 206Pb/204Pb, 143Nd/144Nd and 176Hf/177Hf ratios, slightly elevated 87Sr/86Sr and negative Δ7/4Pb values, and slightly lower δ7Li values than the “Faial” end-member, which has higher 87Sr/86Sr ratios, lower 206Pb/204Pb, 143N/144Nd and 176Hf/177Hf ratios and slightly higher δ7Li values. Mixing of the deeply derived Terceira plume mantle with a relatively shallow and regional fragment of recycled metasomatized mantle wedge underlying the Pico-Faial lineament could explain these systematics.

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5.4 Heterogeneous components in the mantle source of São Miguel basalts The São Miguel basalts have average δ7Li values that are slightly lower than the Central Group island basalts, but also within the MORB range. Of particular note is the lack of any correlation between δ7Li and any radiogenic isotope system, in contrast to what is observed for the Central Group island basalts, and despite the extreme variations in the radiogenic isotopic compositions of these basalts. The São Miguel basalts exhibit evidence for two component mixing between a relatively depleted mantle source component similar to the most enriched APT, and a highly enriched mantle composition intermediate between end-member HIMU and EMII signatures (Fig. 3, 8 and 9; Yu et al., unpublished data). As discussed above, the EMII type component from a previous study (e.g. Bullenmerri xenoliths from Southeastern Australia; Nishio et al., 2004) was shown to have δ7Li signatures within the MORB range or slightly higher, and was attributed to a mantle source metasomatized by a “hydrous metasomatic agent” (Nishio et al., 2004), although the source of the “hydrous metasomatic agent” was not explained. The enriched São Miguel source, on the other hand, has been proposed to result from subducted terrigenous sediments (Hawkesworth et al., 1979; Turner et al., 1997), delaminated subcontinental lithosphere from the opening of the North Atlantic (Davies et al., 1989; Widom et al., 1997; Moreira et al., 1999), or ancient recycled oceanic basalts that were enriched either by underplating of E-MORB melts or evolved recycled crustal material (Beier et al., 2007; Elliott et al., 2007). One might speculate that recycled altered oceanic crust (variable δ7Li) mixed with terrigenous sediment (low δ7Li) could produce magma within the MORB range of δ7Li in the São Miguel basalts. However, a number of recent studies have argued that recycled terrigenous sediment is unlikely to explain the enrichment of the São Miguel mantle source (Widom and Farquahar, 2003; Beier et al., 2007; Elliott et al., 2007). Normal Nb/U and slightly elevated Ce/Pb ratios of São Miguel basalts compared to global MORB argue against a significant contribution of terrigenous sediment in the enriched mantle source, because mixing with more than 1-2% subducted sediments should produce anomalously low Ce/Pb and Nb/U (Hofmann et al., 1986). Furthermore, δ18O signatures of olivines from São Miguel basalts (4.92±0.03‰) are lower than typical mantle values

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(5.2‰) and do not exhibit a positive correlation with Sr or Pb isotopes, thus restricting any terrigenous sediments component (δ18O ~15‰) in the mantle source of São Miguel to <1% (Widom and Farquahar, 2003). Based on coupled enrichment of Rb/Sr, U/Pb and Th/Pb in the source of the São Miguel basalts, Beier et al. (2007) argued against the presence of sediments, suggesting that even less than 0.2% sediment would increase Pb concentrations and produce unradiogenic Pb isotope ratios with time. The Li isotopic signature of delaminated subcontinental lithospheric mantle is in the range of global fresh MORB (Jeffcoate et al., 2007; Ionov and Seitz, 2008). If delaminated subcontinental lithospheric mantle was contained in the mantle source of São Miguel, this could produce the Li isotopic signature found in the São Miguel basalts. However, Os isotopic data indicate radiogenic rather than sub-chondritic 187Os/188Os ratios in these basalts. 187Os/188Os ratios of up to 0.157 in basalts with relatively high Os concentration (214 ppt; Table 1) is unlikely to be produced by subcontinental lithospheric mantle, because ancient subcontinental lithospheric mantle generally has strongly sub-chondritic ratios (Walker et al., 1989; Pearson et al., 1995; Hassler et al., 1998). Therefore, it is unlikely that a component of delaminated subcontinental lithospheric mantle is the enriched component in the mantle source beneath the São Miguel, but rather an enriched component with radiogenic 187Os/188Os. Elliott et al. (2007) proposed that ancient (~3Ga) oceanic lithosphere containing underplated E-MORB basalts produced by modest-degree melting (~2%) of garnet peridotite could produce the enriched component in the São Miguel source. Recycling of lithosphere in which the enriched component is underplated could explain the lack of chemical fingerprints associated with near-surface alteration or subduction dehydration, and could explain unmodified, MORB-like δ7Li in the São Miguel basalts. Ancient recycled oceanic crust with an evolved seamount is an alternative possible component in the enriched São Miguel source (Beier et al., 2007), but modification of the δ7Li might be expected in this case. Based on previous studies, recycled oceanic crust may preserve elevated δ7Li signatures imparted during seafloor alteration (Chan et al., 2009; Vlastélic et al., 2009); if so, the MORB-like δ7Li in the São Miguel basalts would argue against this model.

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However, given the potential for rapid diffusion and homogenization of Li isotopes as discussed previously, it is important to consider the timescales over which we might expect subduction-related Li isotope signatures to survive in the mantle, whether imparted by hydrous slab-derived fluids or residual dewatered subducted oceanic crust.

5.5 Lithium isotope diffusion in time and space Lithium is a fast diffusing element compared to most others. Halama et al. (2008) modeled Li diffusion out of spheres with variable radii and over variable timescales, and found that for a sphere with a radius less than 2 km, a distinctive Li signature would be erased within 100Ma. In this modeling, Halama et al. (2008) assumed the recycled material had a homogeneous but anomalous Li isotopic composition at t=0, but with a constant Li concentration at the sphere surface equal to the average Li concentration of the mantle (Jagoutz et al., 1979). Based on modeling of Li concentration gradients, they suggested that recycled material with an anomalous Li isotopic signature would become homogenized with the surrounding mantle in less than 100Ma. If so, then it is unclear why some HIMU-type basalts appear to have inherited elevated δ7Li from their ancient mantle source. However, Halama et al. (2008) only modeled the Li concentration variations in the recycled material. Because the diffusion rates of 7Li and 6Li are different, it is necessary to consider the diffusion of the two isotopic species independently. Vlastélic et al. (2009) performed such modeling of Li isotope diffusion in subducted crust in a mantle environment. Their model parameters differed from that of Halama et al. (2008) in that they used a more complex boundary condition with variable Li concentrations instead of a constant Li concentration. They proposed that mantle convection would not homogenize Li effectively enough to maintain a constant Li concentration at the contact surface between the subducted slab and surrounding mantle. The Li concentration at the boundary depends on the distance to the center of the subducted slab and the length of time that the subducted slab has resided in the mantle, and is higher than the surrounding mantle at t>0. Therefore, the difference in Li concentration between the boundary and the subducted slab will be smaller than in the model of Halama et al. (2008), which will result in longer diffusion times if the other parameters in the two models are the same. However, in the Vlastélic et al. (2009) model,

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they used a one dimension model, which was only evaluated for the case in which the length and width of subducted slab were much greater than its thickness. Because the subducted oceanic crust could have a thickness of up to ~8 km, for a volume of mantle wedge with comparable dimensions (e.g. given coupled subducted wedge thicknesses of 5-20 km; Hacker et al., 2003), it might be more applicable to consider the case for diffusive behavior in a fragment of recycled material with subequal dimensions of length, width and thickness. Therefore, whether considering the case of recycling a fragment of oceanic crust or a volume of fluid-metasomatized mantle wedge, one dimensional modeling may not simulate Li diffusion accurately. In this study, we used the modeling approach of Halama et al. (2008), but considered both Li concentration and δ7Li variations during diffusion. To simplify the modeling, we used a constant Li concentration equal to that of the surrounding mantle as the boundary condition, and we considered several cases including a larger (8 km radius) fragment of subducted material than considered by Halama et al. (2008). The equation used to describe the Li diffusion is from Crank (1975):

(1)

The recycled material is represented as a sphere of anomalous composition that undergoes radial diffusion into the surrounding mantle. In the equation, a is the radius of the sphere, t is the time, r is the distance from the center of the sphere, C1 is the original

Li concentration in the sphere and C0 is the Li concentration at the sphere’s surface (and equal to the Li concentration of the surrounding mantle). D is the diffusion coefficient of Li from Coogan et al. (2005), and is based on experimentally determined diffusion coefficients for Li in clinopyroxene, which are related to temperature:

(2) where T is temperature in degrees K and R is the gas constant (8.31 J/mol K). The rate of Li diffusion in olivine is not well constrained, but most previous studies suggest it is slower than in clinopyroxene (Spandler and O’Neill, 2006; Jeffcoate et al., 2007; Parkinson et al., 2007; Rudnick and Ionov, 2007). Therefore, we chose to use the Li diffusivity in clinopyroxene for our modeling, since it will result in a shorter

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homogenization time than would other diffusion coefficient choices. This is also the diffusion coefficient used in the Halama et al. (2008) model. The diffusion rate of 7Li is slower than 6Li, and the correlation of their diffusion coefficients can be described as: (3) where β in equation (3) is 0.215 (Richter et al., 2003). Here the D in equation (2) represents D6, and D7 is calculated from equation (3). The initial compositions used in the modeling are those of the inferred recycled material with a Li concentration of 8 ppm (C1) and δ7Li of +10‰ (Chan et al., 2002). The surrounding mantle (and the composition of 7 the sphere surface) has a Li concentration of 1.5 ppm (C0, Jagoutz et al., 1979) and δ Li of +3.2‰ (Seitz et al., 2007). A range of radius sizes (0.5 km, 2 km and 8 km) were chosen for the recycled material in the modeling. If the Li concentration and δ7Li at the point where r/a=0.5 (r = the distance from the center of the subducted slab or mantle wedge to the edge, and a = the radius of the subducted slab or mantle wedge) at mantle temperatures (>1400 ˚C, Putirka, 2005) are within the range of MORB mantle values, we infer that Li was thoroughly homogenized by diffusion. In contrast, if δ7Li at that point is higher or lower than MORB mantle values (>+5‰ or <+3‰), we infer that the Li isotopic signature of the recycled material has not been homogenized by diffusion, but rather retains an anomalous Li isotopic signature. Fig. 13 shows Li concentrations and δ7Li profiles of recycled material with different sizes and different ages over a range of temperatures. Based on this modeling, both the Li concentrations and Li isotopic signatures of recycled material with a 0.5 km radius become homogenized at r/a=0.5 and 1450 ˚C within 25Ma (Fig. 13a and 13b). The peak δ7Li signature, which is even higher than the initial recycled material, is caused by the differential diffusion coefficients for 7Li and 6Li. Figure 13c and 13d show that recycled material with a 2 km radius becomes homogenized with the surrounding mantle in Li concentration and δ7Li at r/a=0.5 and 1450 ˚C after 250Ma. If the recycled material has a radius of 8 km, it can maintain a distinct isotopic signature beyond 2500Ma. In this case, at the point of r/a=0.5 and 1450 ˚C, δ7Li is +8.6‰ after 2500Ma of diffusion, even though the Li concentration has closely approached that of the surrounding mantle (1.54 ppm). As discussed previously,

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the diffusion timescales calculated in this model are likely faster than what would occur in the Earth’s mantle due to lower diffusion coefficients for Li in olivine relative to clinopyroxene, as well as likely variable Li concentrations at the boundary. Therefore, distinctive Li isotopic signatures imparted to the mantle by subduction processes should be retained over timescales even longer than those calculated in this model. These results imply that a fragment of ancient recycled oceanic crust with an evolved seamount would likely retain its elevated δ7Li over the ~3Ga recycling time proposed by Beier et al. (2007) to explain the enriched São Miguel source, unless it were substantially smaller than 8 km in radius. Alternatively, the São Miguel basalt source may be produced by underplating of enriched basalts (Elliott et al., 2007) that would already have MORB-like δ7Li at the time of recycling.

6. Conclusions The Azores basalts exhibit limited variations in δ7Li, which are within the range of global MORB despite large variations in radiogenic isotope systems. The δ7Li signatures of the Central Group island basalts correlate with radiogenic isotopes and are slightly higher than those of São Miguel. The slightly elevated δ7Li signatures in the Faial mantle source relative to Terceira and Pico are interpreted to represent mixing between a recycled mantle wedge component with slightly elevated subduction-related δ7Li and the FOZO-like Terceira source with a slightly lower but MORB-like δ7Li. The lack of anomalous δ7Li signatures or any correlation of δ7Li with radiogenic isotope signatures in the São Miguel basalts may reflect a source generated by recycling of underplated basalts rather than seawater altered and subduction dewatered crust, or a small (<<8 km) body of recycled seawater altered oceanic crust that was diffusively homogenized with the surrounding mantle over its ~3Ga residence time.

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Table 1. Compositions of Azores basalt samples including lithium concentrations and isotopic compositions Sample name WAF1a WAF 30 WAP8 WAP9 WAP10 WAT1 WAT3 WASM 1 WASM5 WASM17 WASM24 SiO2 46.38 46.65 46.84 47.63 48.58 46.74 47.75 45.54 47.31 44.40 45.86 TiO2 2.11 2.86 2.47 2.13 1.31 4.22 1.90 3.32 3.22 4.09 3.27 Al2O3 12.32 15.55 11.86 11.87 7.72 14.67 13.36 11.89 10.99 12.81 12.87 Total Fe 10.85 11.30 10.81 10.14 8.21 13.77 9.68 12.91 13.06 13.64 12.21 MnO 0.15 0.17 0.18 0.16 0.14 0.21 0.15 0.18 0.17 0.18 0.17 MgO 13.51 8.30 12.05 13.21 17.04 5.02 11.80 11.13 11.07 9.38 10.26 CaO 10.90 10.86 12.03 10.95 14.79 10.19 11.54 10.71 10.09 11.61 11.22 Na2O 2.54 2.90 2.44 2.47 1.50 3.45 2.55 2.37 2.26 2.47 2.55 K2O 0.89 0.95 0.86 0.98 0.44 1.21 0.98 1.52 1.38 0.98 1.07 P2O5 0.34 0.45 0.46 0.44 0.27 0.52 0.29 0.44 0.45 0.44 0.52 LOI -0.42 -0.27 -0.39 -0.15 -0.21 -0.32 -0.15 -0.51 -0.51 0.04 -0.16 Total 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 100.0 *Norm Factor 0.991 1.003 1.004 1.001 1.002 1.003 0.994 1.005 1.001 0.990 0.992

Li 4.8 5.6 4.7 5.2 3.0 5.7 4.3 5.1 6.8 5.1 5.1 Be 1.2 1.1 1.3 1.4 0.7 1.6 1.2 1.6 1.4 1.4 1.3 Sca 33 29 37 33 52 27 31 30 26 32 31 V 210 293 248 191 123 416 242 272 266 323 292 Cra 700 289 651 1029 2537 20.6 780 743 488 449 660 Co 59 47 57 56 58 45 54 67 68 59 60 Ni 339 124 255 335 387 32 324 393 321 164 290 Cu 63 52 58 58 29 36 77 170 52 86 111 Zn 81 100 92 83 62 131 69 111 109 119 97 Ga 15 20 16 15 10 22 15 19 19 21 19 As 0.53 0.19 0.95 0.68 1.00 1.08 0.65 0.39 0.57 0.83 0.78 Rb 18 20 17 20 9 25 23 37 34 36 29 Sr 410 555 417 422 230 540 381 605 456 600 600 Y 20 27 23 21 13 34 20 28 30 25 25 Zr 143 182 159 156 76 236 127 270 260 257 210 Nb 26 35 29 28 13 50 29 53 44 56 48 Cs 0.18 0.15 0.19 0.22 0.10 0.29 0.23 0.36 0.16 0.35 0.31 Baa 247 279 236 292 117 342 291 408 342 460 407 La 21 27 22 23 10 33 19 42 37 41 32 Ce 47 56 52 51 23 71 40 90 80 86 69 Pr 5.7 7.1 6.4 6.2 2.9 9.0 4.9 11.0 9.9 10.4 8.6

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Nd 24 29 27 25 13 37 20 45 41 41 35 Sm 5.6 6.7 6.4 5.8 3.2 8.5 4.7 9.4 9.1 8.8 7.8 Eu 1.7 2.1 2.0 1.8 1.0 2.7 1.5 2.7 2.6 2.6 2.4 Gd 4.9 6.7 5.5 5.0 2.9 7.5 4.1 7.7 7.6 7.5 6.5 Tb 0.77 0.93 0.86 0.77 0.48 1.21 0.66 1.12 1.15 1.10 0.95 Dy 4.3 5.1 4.8 4.3 2.7 6.6 3.8 5.8 6.1 5.7 5.1 Ho 0.84 1.01 0.92 0.82 0.53 1.30 0.77 1.09 1.18 1.04 0.96 Er 2.1 2.6 2.3 2.0 1.3 3.4 2.1 2.8 3.1 2.7 2.5 Tm 0.28 0.37 0.31 0.29 0.18 0.45 0.27 0.34 0.38 0.33 0.31 Yb 1.7 2.2 1.9 1.7 1.1 2.7 1.7 2.0 2.3 2.0 1.8 Lu 0.24 0.32 0.27 0.24 0.16 0.40 0.24 0.28 0.33 0.28 0.25 Hf 3.8 4.4 4.1 4.1 2.0 5.7 3.2 6.7 6.6 6.5 5.1 Ta 1.8 2.4 2.1 2.0 0.9 3.5 1.9 3.6 3.0 4.1 3.2 W 0.33 0.35 0.49 0.37 0.22 1.19 0.34 0.85 0.73 0.92 0.61 Pb 1.72 1.44 1.79 0.80 1.98 1.56 2.58 2.41 2.04 1.93 Th 2.1 2.7 2.2 2.3 0.9 3.0 2.0 5.0 4.5 4.6 3.6 U 0.63 0.73 0.74 0.67 0.29 1.08 0.61 1.31 1.15 1.16 1.02

δ7Li ‰ 4.4 4.7 4.0 4.5 3.7 4.0 3.7 3.7 3.3 3.8 3.1 Li (ppm)b 4.3 5.3 4.3 4.8 3.0 4.9 4.1 4.5 6.0 4.5 4.5

87Sr/ 86Sr 0.70388 0.70391 0.70363 0.70393 0.70378 0.70360 0.70344 0.70448 0.70575 0.70383 0.70352 206Pb/204Pb 19.579 19.299 19.977 19.519 19.887 19.982 19.768 19.892 20.095 19.642 19.533 207Pb/204Pb 15.628 15.618 15.598 15.588 15.571 15.611 15.605 15.730 15.785 15.636 15.604 208Pb/204Pb 39.262 39.087 39.335 39.093 39.068 39.278 39.272 40.106 40.345 39.595 39.353 143Nd/144Nd 0.512860 0.512873 0.512896 0.512839 0.512901 0.512950 0.512913 0.512737 0.512680 0.512844 0.512886 176Hf/177Hf 0.28296 0.28300 0.28303 0.28296 0.28300 0.28305 0.28304 0.28283 0.28275 0.28291 0.28298 187Os/188Os 0.1269 0.1336 0.1249 0.1255 0.1244 0.1252 0.1317 0.1567 0.1293 0.1269 Os ppt 124.9 5.0 22.5 139.6 94.1 0.8 111.9 153.7 213.9 43.9 108.4 The samples labeled as a were analyzed by DCP; b were analyzed by MC-ICPMS; other trace elements are analyzed by ICP-MS. The unit of measurement for trace elements is ppm. *Norm factor represents the normalizing factor used to correct the major element totals to 100%.

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Table 2. δ7Li values of four olivine – whole rock pairs. Sample Li (ppm) δ7Li ‰ WASM24w 3.1 3.5 WASM24o 1.5 3.5 WAP8w-1 6.3 2.1 WAP8w-1r 7.4 2.2 WAP8o-1 1.2 5.3 WAP8o-1r 1.4 4.8 WAP8w-2 5.0 3.5 WAP8w-2r 5.1 3.7 WAP8o-2 1.0 3.0 WAP8o-2r 1.1 3.0 WAP9w 4.1 4.7 WAP9o 1.4 4.9 WAP10w 2.4 4.4 WAP10o 1.2 4.4 “w” designates whole rock powder, “o” designates olivine crystal separates, and “r” designates replicated sample.

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Figure 1. Lithium isotopic compositions of Earth reservoirs. Modified after Elliott et al. (2004) and Tomascak (2004).

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Figure 1 129

Figure 2. Map of Azores (modified after Moreira et al., 1999). The Azores archipelago is located between ~37-40°N, in the vicinity of the Mid-Atlantic Ridge (MAR) and the triple junction between the North American, African and Eurasian plates. The archipelago comprises nine islands situated on both sides of the mid-Atlantic ridge. The basalt samples in this study are from the islands of São Miguel, Terceira, Faial and Pico. The islands belong to the Central Group, except for São Miguel, which is an eastern group island.

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Figure 2

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Figure 3. 87Sr/86Sr versus 206Pb/204Pb for basalts of the Azores archipelago (this study, Yu and Widom, 2010 and Yu et al., unpublished data). Azores basalts exhibit extreme heterogeneity on an intra-island and archipelago-wide scale, ranging from relatively depleted signatures in the Azores Platform MORB to relatively strong HIMU, EMI and EMII signatures in the Azores island basalts. The symbols with white borders are the samples analyzed for Li isotopes. APT is Azores platform tholeiites (data from White et al., 1976; White and Schilling, 1978; Dupré and Allègre, 1980; Hamelin et al., 1984; Ito et al., 1987; Dosso et al., 1996). FOZO mantle component is based on data from Stracke et al. (2005). Black symbols show the data from literature (White et al., 1979; Davies et al., 1989; Turner et al., 1997; Moreira et al., 1999; Millet et al., 2009), and for the samples from a given island, the same symbol shape is used. Our isotopic data show similar variations to previous studies.

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Figure 3

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Figure 4. Deviation of individual analyses from the mean of replicate analyses of all standards (L-SVEC 50ppb). The first two and the last two standards exhibit a relatively larger deviation from the mean (>±0.3‰), but only the first standard is outside of ±0.6‰, which we used as the analytical error in this study. All samples were analyzed after the third standard and before the sixteenth, during which time the behavior was very stable and the standard exhibited variations of <±0.3‰ about the mean.

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Figure 4

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Figure 5. Comparison of δ7Li values between whole rocks and olivine separates. Except for the first analysis of the WAP8 pair (WAP8-1 and WAP8-1r), the olivine-whole rock pairs plot on or near the 1:1 correlation line within analytical error.

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Figure 5

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Figure 6. Histogram of Li isotopic composition of the Azores ocean island basalts (a) and MORB (b). East Pacific Rise (EPR) MORB data are from Chan et al. (1992), Elliott et al. (2006) and Tomascak et al. (2008); Mid-Atlantic Range (MAR) MORB data are from Chan et al. (1992) and Tomascak et al. (2008); Indian Ridges (IR) MORB data are from Nishio et al. (2007) and Tomascak et al. (2008); North Atlantic Ridge data are from Nishio et al. (2007); and other MORB data are from Moriguti and Nakamura (1998b) and Tomascak et al. (2008).

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Figure 6

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Figure 7. δ7Li values of whole rock powders from the Central Group islands versus (a) 206Pb/204Pb, (b) 87Sr/88Sr, (c) 143Nd/144Nd, and (d) 176Hf/177Hf. δ7Li signatures in São Miguel basalts do not correlate with radiogenic isotopes (not shown). For the Central Group island samples, δ7Li correlates positively with 87Sr/86Sr and negatively with 206Pb/204Pb, 143Nd/144Nd and 176Hf/177Hf. The highest δ7Li signatures are found in samples from Pico and Faial.

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Figure 7

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Figure 8. Modified from Ryan and Kyle (2004), Nishio et al. (2007), Simons (2008), Simons et al. (2010b), Chan et al. (2009) and references therein. The δ7Li values for all of the Azores basalts fall within the range of global MORB, and none has anomalously high δ7Li as reported in some other OIB. Nevertheless, there are two apparent trends in the Azores samples: a near-horizontal trend for the São Miguel samples, and a positive trend for the Central Group island samples, which are on average slightly higher in δ7Li than the São Miguel samples. Both of these two trends share the same end-member of Azores platform tholeiites (APT)

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Figure 8

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Figure 9. Osmium concentrations and isotopic compositions of the Central Group island basalts (this study and Yu et al., in prep.). (a) 187Os/188Os of the Central Group islands fall in the range of abyssal peridotites (Alard et al., 2005) and green line is the peak of the Os isotope variation; (b), (c), (d) and (e) show the 187Os/188Os signatures of the Azores basalts versus 206Pb/204Pb, 87Sr/86Sr, 143Nd/144Nd and δ7Li respectively.

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Figure 9

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Figure 10. Major and trace elements versus δ7Li for Azores basalts.

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Figure 10 147

Figure 11. Plots of isotopic ratios, and major and trace element compositions of the

Azores basalts (this study, Yu et al., in prep.; Yu et al., unpublished data). (a) K2O vs. Nb (ppm) of the Azores basalts. There are two trends in the Central Group island samples: most Terceira samples have lower K2O values than the samples from the other islands at a given Nb values. (b) Rb/Be vs. K/Be of the Azores basalts. All of the Central Group island samples plot on a positive trend. (c) K2O values and (d) Nb/Zr vs. MgO of the Azores basalts. (e) 206Pb/204Pb vs. K/U of the Azores basalts. K/U of the extreme HIMU components are from Sun and McDonough (1989). (f) La/Nb vs. Ba/Nb of the Azores basalts. Ba/Nb and La/Nb of the extreme HIMU components are from Weaver (1991).

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Figure 11

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Figure 12. (a) Lithium isotope and Os isotope mixing trends for mixing of magmas with variable Os concentrations and pelagic sediment (PS) and (b) Sr isotope and Os isotope mixing trends for mixing of magma and marine sediment. AM is Azores magma, with [Li]=4 ppm, δ7Li=+3‰, [Os]=5 ppt, 30 ppt and 140 ppt (the variation in this study), 187Os/188Os=0.122 (equal to the depleted mantle, Snow and Reisberg, 1995), [Sr]=90 (Sun and McDonough, 1989) and 87Sr/86Sr=0.7025, which reflects the isotopic composition of enriched mantle plume (Hart et al., 1992) and is in the range of Azores platform tholeiites. PS is pelagic sediment with [Li]=30 ppm (You and Chan, 1996), δ7Li=+15‰ (Bottomley et al., 1999; Chan et al., 2006), [Os]=20 ppt (Peucker-Ehrenbrink et al., 2003), 187Os/188Os=1 (Levasseur et al., 1998), [Sr]=115 (Becker et al., 2000) and 87Sr/86Sr=0.710 (Weaver et al., 1991).

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Figure 12

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Figure 13. Diffusion models for Li concentrations and δ7Li of sphere shaped recycled material with variable sizes and mantle residence times. Axes show the variation of temperature (x-axis), the ratio of the distance to the center and radius (y-axis), and Li concentration and δ7Li (z-axes). (a) and (b) show Li concentration and δ7Li of recycled material with 0.5 km radius at 25Ma after subduction. (c) and (d) show Li concentration and δ7Li of recycled material with 2 km radius at 250Ma after subduction. (e) and (f) show Li concentration and δ7Li of recycled material with 8 km radius at 2500Ma after subduction. The initial Li concentration of the recycled material is 8 ppm and δ7Li is +10‰ (Chan et al., 2002), and the Li composition of the surface, which is equal to the surrounding mantle, is [Li]=1.5 ppm (Jagoutz et al., 1979) and δ7Li=+3.2‰ (Seitz et al., 2007). r is the distance to the center of the subduction-related material, and a is the radius of the subduction-related material. The black dots show the position of r/a=0.5 at a mantle temperature of 1450 ˚C. The Li concentrations and δ7Li of each black dot are (a) [Li]=1.50 ppm; (b) δ7Li=+3.2‰; (c) [Li]=1.50 ppm; (d) δ7Li=+3.6‰; (e) [Li]=1.54 ppm; (f) δ7Li=+8.6‰.

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Figure 13

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CHAPTER 4

Microwave Digestion Method for Os analysis

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Abstract We have developed two microwave digestion methods for Os isotopic analysis, with efficient recoveries of Os and low blank levels. One method involves a standard reverse aqua regia leaching procedure, and the other achieves complete dissolution in HF-HCl-EtOH. Osmium recoveries for both methods are > 90% with total procedural Os blanks <0.6 pg. Compared to previous digestion methods, microwave digestion is substantially faster (4 hours vs. 12-48 hours) and does not require expensive and custom-made consumables (e.g. Carius tubes) or equipment that is incompatible with a clean lab (e.g. gas cylinders and torches). Furthermore, in contrast to Carius tubes and high pressure ashers, the Teflon microwave vessels allow for use of HF to achieve complete dissolution of silicate samples. In addition to being time and cost efficient, this digestion method is safer than previous methods, because the pressure in the microwave can be monitored and controlled, so that samples will not burst or explode during digestion. We have also modified the “macro” distillation method developed by Nägler and Frei (1997) for separation and purification of single Os samples. Our method allows for the simultaneous distillation for 8 (or more) samples under controlled conditions that produce yields of >70%, similar to that of Nägler and Frei (1997). In this method, samples dissolved by microwave digestion are transferred to the distillation system, which extracts Os from reverse aqua regia and traps it in HBr. Following the macro-distillation, Os is further purified by a micro-distillation step (Roy-Barman, 1993; Roy-Barman and Allègre, 1995) prior to isotopic analysis. This technique can be easily employed and modified to accommodate any number of samples for a variety of geological and cosmochemical applications.

Keywords: geochemistr; microwave digestion; distillation; osmium separation

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1. Introduction With the increasing interest in the Re-Os isotope system as an important tool for geochemical, cosmochemical and geochronological applications, multiple analytical methods for Re-Os separation and mass-spectrometric techniques have been explored and developed to maximize yield and ionization efficiency, and minimize processing blanks. For most geological samples, Os abundances are in the ppb to ppt range, thus relatively large amounts of sample material must be processed (≥1g), and maintaining low blank levels (<1 pg) is critical. In addition, to obtain an accurate and precise Os concentration and isotopic composition, a high extraction yield and full equilibration between the spike and the sample must be achieved. The digestion methods used for Os analysis have evolved through time, with most of the earlier work utilizing Teflon bomb digestion methods (Walker, 1988; Roy-Barman and Allègre, 1995). Carius tube digestion methods were developed in the mid-1990’s (Shirey and Walker, 1995; Nägler and Frei, 1997) and are currently utilized in many labs, although more recently some labs have employed high pressure asher (HPA-S) digestion methods (Meisel et al., 2001a). Teflon bomb digestion involves sealing the sample powder and spike with HF and reducing solutions (HBr or a mixture of HCl-EtOH) in Teflon digestion vessels that are placed in an oven at 140 ˚C for 8 hours, in order to digest the sample and equilibrate the sample and spike. This digestion method can dissolve silicate rock samples by breaking the Si-O bonds and releasing silicon as volatile silicon tetra-fluoride. However, previous studies have shown that some samples are ineffectively dissolved and equilibrated using the Teflon bomb. In such cases, samples that produce isochronous behavior when dissolved by Carius tube, show significant scatter about an isochron when dissolved by the Teflon bomb method, indicating that the Teflon bomb method may incompletely extract Os from samples, fail to equilibrate sample and spike, or result in loss of sample or spike Os (Shirey and Walker, 1995). Carius tube digestion involves sealing sample and spike with reverse aqua regia or other oxidizing reagents (e.g. CrO3 and H2SO4 solution) in a Carius tube, and heating in an oven at 220-260 ˚C for at least 12 hours (and more commonly 24-48 hours) to extract Os and equilibrate sample and spike. This digestion method generally extracts Os sufficiently for sulfides and some silicates (such as mantle peridotites), and results in a

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low blank (~0.2 pg) and high Os yield (>80% Shirey and Walker, 1995). However, for silicate samples in which Os is structurally bound or occurs as inclusions within the silicate minerals, Carius tube digestion cannot extract Os completely due to the lack of HF to break down the silicate structure. The high pressure asher digestion method involves adding up to 2g sample powder to a quartz glass vessel with chilled, concentrated HCl (3ml) and HNO3 (5ml), followed by the appropriate amount of spike. The vessel is then sealed and pressurized with N2 to 100 bar and heated to 300 ˚C for 4 hours. This method is similar to the Carius tube digestion method in that the glass vessel prevents the use of HF in the digestion procedure, hence does not achieve complete dissolution of silicates. Microwave digestion has not been widely adopted for Os analysis, due to potential issues with memory effects in Teflon (Roy-Barman and Allègre, 1995; Schoenberg et al., 2000) and concerns that the volatile oxide species may stick to, or be lost by diffusion through, Teflon vessel walls during digestion. Suzuki et al. (1992) reported several tests of microwave digestion techniques for analysis of Os isotopes and concentrations in molybdenite by a CEM MDS-81D microwave digestion system. They concluded that the most reliable procedure, which achieves both complete decomposition of the sample and complete oxidation of the Os, is to dissolve the molybdenite in HNO3 and H2SO4 in Teflon (PTFE) microwave vessels first, and then after cooling, to add a strong oxidizing agent such as K2Cr2O7 to the vessels, and return the solution to the microwave. However, this study focused only on molybdenite samples with very high Os concentration (>200 ppb), and did not investigate the procedure for effectiveness in silicate materials that generally have Os only in the low ppb to ppt range. Here, we report a new microwave digestion method that can combine both HCl-HNO3 leaching and HF digestion methods for Os analysis of environmental and geological materials including silicates with ppt range Os concentrations. Compared with other methods, this method is faster, safer and easier.

2. Analytical methods and experiments 2.1 Microwave digestion and Carius tube digestion

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The microwave has been used widely to digest a variety of environmental and geological samples for analysis by graphite furnace atomic absorption spectrometry (GFAA), inductively coupled plasma optical emission spectrometry (ICP-OES), or inductively coupled plasma mass spectrometry (ICP-MS). So far, few experiments have been done to test microwave digestion techniques for Os analysis (Suzuki et al., 1992; Hassler et al., 2000). As discussed previously, Os isotopic analysis of geological samples with ppt to ppb level Os requires low procedural blanks, high yields during extraction and purification of sample Os, and complete equilibration of sample and spike. Therefore, in order to establish the utility of microwave digestion for analysis of Os in geological samples, it is necessary to evaluate the microwave digestion procedures to confirm that high yield extraction of Os is achieved, that sample and spike Os are equilibrated, and that it produces a sufficiently low procedural blank level. In this study, we first tested the feasibility of microwave digestion of silicate rock samples for Os analysis by digesting a powdered basalt rock sample (WAT13 from Terceira, Azores) by both microwave and Carius tube digestion methods to compare results. After confirming that similar isotopic compositions and Os concentrations were obtained from these two methods, indicating that the microwave digestion method produced adequate spike-sample equilibration and that Os was not lost entirely, we then performed additional experiments to determine the extraction yield for Os and the total procedural blank level for the microwave digestion method, to fully assess the utility of this method for Os isotopic analysis of environmental and geologic samples with low (ppt level) amounts of Os. To compare the Os concentrations and isotopic ratios determined for the WAT13 basalt sample by Carius tube and microwave digestion, two sample aliquots were digested by Carius tube, and four sample aliquots were digested by microwave. The Carius tube digestion followed the methods described by Shirey and Walker (1995). One gram of whole rock powder and an appropriate amount of 190Os spike were weighed and transferred to a Carius tube chilled by a dry ice-EtOH (~36% EtOH) mixture, and 2ml concentrated HCl and 4ml concentrated HNO3 were then added to the sample. The tubes were then immediately sealed, warmed to room temperature, and placed in stainless-steel jackets. The tubes were then heated to 240 ˚C for 48 hours to extract Os from the rock

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and achieve spike-sample equilibration. After cooling to room temperature, the Carius tubes were frozen in an EtOH-dry ice slush and the tubes were opened. Once the solutions were sufficiently defrosted but still chilled, they were transferred directly to the distillation system for separation of Os. The microwave digestions were performed in a Milestone Ethos EZ microwave system equipped with both high-temperature and high pressure rotors and vessels. The recommended temperature, time and pressure limits for the high pressure vessels are 225-230 ˚C for less than 25 minutes, with pressures ≤70 bars. For the high temperature vessels, the recommended temperature and pressure limits are 260 ˚C for less than 2 hours or 270 ˚C for less than 20 minutes, with pressures ≤60 bars. Although ultimately the higher temperatures and longer heating times allowed by the high temperature rotor are preferable for sample digestion and spike-sample equilibration, initial experiments were performed in the high pressure rotor, which is equipped with a pressure sensor. The pressure sensor measures pressure during heating in a reference vessel that contains the same amount and type of sample and reagents as the other sample vessels. Before testing any new digestion method in the high temperature vessels, experiments were first conducted in the high pressure vessels to monitor the pressure change during heating and reaction of the samples and reagents with increasing temperature, up to 230 ˚C. For experiments using ~1 g sample and 10.5ml concentrated HCl (3.5ml) and HNO3 (7ml), the maximum pressure produced was ≤32 bars (Fig. 1), which is well within the range that can be accommodated in the high temperature vessels. Because the high temperature vessels allow sustained operation at temperatures ≥230 ˚C, which should aid sample-reagent reaction and spike-sample equilibration, we subsequently used high temperature vessels to digest the samples for Os analysis. The first set of microwave digestion experiments (WAT13_3 to WAT13_6, Table 1) was performed on 1 gram of basalt sample WAT13 with added 190Os spike, in order to compare the measured Os concentration and Os isotope ratio in the sample resulting from the two different methods. The reagents used in the first microwave digestion method that we evaluated are the same as those used in the Carius tube digestion method, both reverse aqua regia with a 2:1 ratio of concentrated HNO3: HCl. However, the acid volumes used in microwave digestion (10.5ml) and Carius tube digestion (6ml) are different, because

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the microwave requires a minimum of 10ml of solution volume, whereas there is concern about Carius tube failure if the tube is filled to >2/3 capacity (Shirey and walker, 1995). For the microwave digestion experiments, samples were weighed and transferred into the high temperature microwave vessels, and spike was then added to the vessels, followed immediately by the addition of chilled, concentrated HCl (3.5ml) and concentrated HNO3 (7ml). The vessels were then immediately sealed and assembled on the rotor, and placed into the microwave. The vessels were heated from room temperature to 230 ˚C in 40 minutes, and then held at 230 ˚C for two hours, for a total digestion period of 3 hours and 40 minutes prior to a one hour cool down. After cooling to room temperature, the microwave rotor with sample vessels was placed in a freezer overnight, in order to chill the sample solutions prior to transferring to the distillation system. In addition to the experiments designed to compare the sample Os concentration and isotope ratio using microwave versus Carius tube digestion methods, the Os recovery in the microwave digestion procedure was separately tested in this study. For these experiments, ~1 ng of Os standard with known concentration was added to the microwave vessel with the basalt sample and reverse aqua regia. Based on the preliminary work, we knew that the Os concentration of the sample (WAT13) is ~10 ppt, so the Os contributed by 1 gram Os sample (~10 pg) is insignificant relative to the 1 ng of Os contributed by the standard. In these experiments, the spike was added to the mixture of sample and standard solution only after running the microwave digestion procedure and chilling the vessels as described previously. In order to ensure spike-sample equilibration at this point, the spiked samples were reheated either by running the same microwave procedure again, or for one sample (MY-3), by transferring the solution to Carius tubes and re-heating using the standard Carius tube method. For sample MY-3, 5ml of solution from the microwave vessel was transferred to the Carius tube. After adding spike to the solution, the Carius tube was sealed and heated to 240 ˚C for 48 hours to equilibrate the sample and spike. For all of the other samples, the spike was added to the microwave vessels that contained the sample and standard solutions, and the vessels were put back into the microwave to equilibrate the sample and spike by repeating the same high temperature program used for the original digestion. The total amount of Os in the sample solutions following the first microwave digestion was the determined by

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isotope dilution calculations following mass spectrometry. The fraction of Os remaining relative to the total initial Os added in the experiments (~1 ng) allows precise calculation of the procedural yield. In this study, we also tested the yield for a microwave procedure using HF with EtOH and HCl to digest silicate samples. Because of the unknown effects of volatile

EtOH and the SiF4 gas produced by sample-HF reaction, these experiments were also first performed in the high pressure vessels to monitor the pressure change with temperature. Sample aliquots (~1 g of WAT13) and Os standard were weighed and transferred into the high pressure microwave vessels. Following addition of 5ml HCl, 1ml EtOH and 5ml HF, the solutions were heated from room temperature to 210 ˚C over 40 minutes, held at 210 ˚C for 20 minutes, and cooled down for 1 hour. The temperature and pressure changes during the run were monitored, and pressures did not exceed 22 bars (Fig. 2), again within the allowable range for the high temperature vessels. After drying, the sample residues were reacted with reverse aqua regia in high temperature microwave vessels following the procedure described previously. Spike was then added to the sample-standard solutions and heated again using the high temperature microwave procedure to equilibrate sample and spike.

2.2 Distillation procedure The distillation system builds on the method of Nägler and Frei (1997) that was developed for single sample distillation. The apparatus developed here allows for eight samples to be distilled simultaneously, and could be easily expanded to increase sample capacity (Fig. 3). The distillation system includes a Teflon coated aluminum heating block (Fig. 3a) with a heater and controller (Fig. 3b) that maintain temperature to ±1 ˚C of the set point, which for this application is 115 ˚C. Eight sample wells in the heating block accept 15ml Savillex vials (Fig. 3c), which are fitted with dual-port caps. One port is an inlet port that is attached via Teflon tubing (1.6 mm inner diameter and 3.2 mm outer diameter) to a manifold that is connected to a small fish-tank air pump fitted with a filter to provide a clean carrier gas for the distillation (Fig. 3d). The manifold serves to separate and feed the filtered air from the fish-tank pump into the 8 Teflon inlet tubes, and allows the air flow rate to be controlled. The other port of the 15ml vial cap serves as

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an outlet port that is connected via Teflon tubing (same diameter as inlet) to 30ml Savillex vials that are set in an ice water bath and serve as the distillate traps. Those vials also are outfitted with dual-port caps, one port connected to the Teflon tubing from the distillation vial, and one serving as a vent. The distillation apparatus was set up in a laminar flow hood with HEPA filtered air. After digestion, sample solutions were transferred from the Carius tubes or microwave vessels into the 15ml distillation vials. These vials were then inserted into the heating block that had been pre-heated to 115 ˚C, and the vials covered with the dual-port caps. A short Teflon FEP tube was inserted through the top port and extending to the bottom of the vial, with the other side of the tube connected to the manifold which is adjusted to produce an airflow resulting in ~2 bubbles per second in the sample vials. A long Teflon FEP tube was connected from the side-port of the sample vial cap through the top port and into the bottom of the trapping vial, which contained 10ml concentrated HBr, chilled in the ice bath. Samples were then distilled at 115 ˚C for 150 minutes to ensure complete transfer of the oxidized Os (as OsO4) from the distillation vials with reverse aqua regia, 2- into the trapping vials where the Os reacts with the HBr and is reduced to stable OsBr6 . Following distillation, the trapping vials were removed from the chilled water bath, covered with portless caps, and placed in the heating block at 90 ˚C for two hours to 2- ensure complete reduction of Os to the stable OsBr6 . The Os-bearing HBr solutions were then evaporated and further purified by a micro-distillation procedure following the method of Roy-Barman (1993) and Roy-Barman and Allègre (1995). The extracted Os, after drying down the HBr trapping solution, was re-dissolved in 20 μl concentrated HBr, and transferred into the cap of a small conical Teflon vial and dried again. 20 μg concentrated HBr was placed in the conical tip of the beaker, to serve as a trap for the microdistillation. After adding 30 μl concentrated CrO3 (dissolved in 6 mol/l H2SO4) to the dried sample in the vial cap, the conical vial was turned upside-down and closed with the lid on the bottom, and covered in aluminum foil except for the conical tip. The vial was then heated upside-down on a hot plate at 80 ˚C for 3 hours allowing the sample Os to distill and be trapped in the HBr. The HBr trapping solution with the sample Os was then evaporated to a volume ~3 μl for loading. The reverse aqua regia sample solutions,

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essentially devoid of Os after the original distillation, were evaporated and prepared for Re separation via conventional anion exchange procedures (Morgan et al., 1991). Following micro-distillation, the Os samples were loaded onto high purity Pt filaments. After slowly drying the loaded sample solution at 0.8A, a small amount of

Ba(OH)2 (≤0.1 μl) was spread over the dried sample and evaporated. The Os isotopic - compositions were analyzed as OsO3 by negative thermal ionization mass spectrometry (TIMS) on a Thermo-Finnigan Triton following the methods of Creaser et al. (1991) and Volkening et al. (1991), and corrected for oxygen isotopic composition and mass fractionation. The oxygen correction assumed oxygen isotopic abundances of Nier (1950), and the mass fractionation correction used a ratio of 192Os/188Os=3.0826.

3. Results and discussion 3.1 Comparison of microwave and Carius tube digestion The results of the experiments from this study are presented in Tables 1 and 2. Table 1 shows the Os concentrations and isotopic compositions of six experiments using basalt sample WAT13 (WAT13_1 to WAT13_6), which compare the results of the microwave digestions with those of Carius tube digestions. In these experiments, WAT13_1 and WAT13_2 were digested by Carius tube, and WAT13_3 to WAT13_6 were digested using the microwave procedure with the high temperature vessels. With the exception of sample WAT13_6, which has a slightly lower Os concentration (8.9 ppt), the other samples have a very limited variation of Os concentrations (9.4 ppt to 10.5 ppt), and all of these data are within 2SD (±1.1 ppt) and ~10% of the mean (Fig. 4a). The internal measurement errors of Os concentrations in these samples are <±0.04 ppt (smaller than the symbols in Fig. 4a) and therefore do not contribute significantly to the variability in measured concentration. The variability in measured Os concentration may be caused by the nugget effect, due to platinum group elements (PGE) being hosted primarily in ultra-trace phases such as sulfides and/or PGE metal alloys (Meisel et al., 2001b) that may not be homogenized even in a finely powdered sample. WAT13_6, which resulted in the lowest measured Os concentration, also has slightly lower 187Os/188Os ratio (0.1300) then the other 5 samples, all of which have relatively limited variation in 187Os/188Os ratios (0.1314 to 0.1328). Nevertheless, all six of the samples produced measured

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187Os/188Os within 2SD (±0.002) of the mean (Fig. 4b). The measured 187Os/188Os ratios and Os concentrations show no systematic difference between the Carius tube and microwave digestion methods (Figure 4), indicating that microwave digestion is effective and comparable to the Carius tube digestion method in the extraction of sample Os and equilibration of sample and spike.

3.2 Blank level Total procedural blanks were determined by combining 190Os spike with the reagents but no sample, and running the blanks through the same digestion and purification procedures as samples. Eight total procedural blanks were performed along with the samples to test the procedural blank for the various digestion methods. Total procedural blanks involving the microwave digestion procedure are reported in Table 1, and ranged from 0.1 to 0.6 pg Os (average = 0.3 pg). These blanks are similar to the typical total procedural blanks obtained with Carius tube digestion (~0.2 pg), implying that there is not a significant blank contribution from the microwave system. Rather, as in the case for Carius tube digestion, the procedural blank comes primarily from the reagents used for sample digestion and purification of Os. The HF and EtOH used in the experiments were purified by sub-boiling double distillation, and the HCl was made by bubbling high-purity Cl2 into 18.0 mega-Ω E-pure water. Prior experiments have demonstrated that most of the blank contribution comes from the HNO3 if sub-boiling double distilled

HNO3 is used directly. Therefore, the double distilled HNO3 was further purified by sparging with triple filtered air for three to five days to remove volatile, oxidized Os. In previous studies, it was found that Teflon bombs could occasionally produce memory effects for OsO4 under high temperatures and pressures (Roy-Barman and Allègre, 1995; Schoenberg et al., 2000). In our experiments, the Teflon vials and Teflon tubes for distillation were cleaned by boiling in 8N HNO3 for 3 hours after each use. The Teflon microwave vessels were cleaned by heating the vessels in the microwave with

20ml of 8N HNO3 at 240 ˚C for 20 minutes. After cleaning, the Teflon vials, tubes and microwave vessels were re-used for both samples and blanks. Given the relatively low procedural blanks, and the similarity between the blanks obtained using microwave and Carius tube digestion, there is no evidence for significant cross contamination despite

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re-use of the Teflonware and the very high amounts of Os (1 ng) run in the experiments. Nevertheless, it is advisable to designate separate Teflonware for use with samples that have dramatically different Os concentrations, and especially for very low level samples, to use new Teflon tubes for each analysis.

3.3 Chemical yields The chemical yields for the two different microwave digestion methods are reported in Table 2. The yield experiments for all samples except MY-4 and MY-5 involved digestion of ~1g WAT13 sample powder and ~1ng Os standard, using the previously described method with 10.5ml reverse aqua regia heated in the high-temperature microwave vessels for 3 hours and 40 minutes. With the exception of MY-3, these sample solutions were then spiked and heated again by microwave with the same high-temperature digestion program to equilibrate sample and spike. For sample MY-3, an aliquot of ~5ml of the reverse aqua regia was transferred from the microwave vessel to a Carius tube and spiked, and the Carius tube heated in an oven at 240 ˚C for 48 hours to equilibrate sample and spike. The chemical yields for the microwave digestion procedure for all three experiments were better than 80%. The samples (MY-1 and MY-2) that were digested and equilibrated after spiking by microwave have yields higher than 90% (average = 95%), which we interpret to be the true yields for the microwave digestion procedure, as the transfer of solution back to Carius tube for sample MY-3 was not precisely quantitative. Samples MY-4 to MY-5 were digested in Ethanol + HF + HCl, again using the high temperature microwave vessels run with the same time and temperature program as for the reserve aqua regia experiments. The sample solutions were then evaporated to dryness, spiked, and reverse aqua regia added prior to reheating in the microwave to equilibrate spike and sample. These samples had yields of 96% and 98% respectively (average = 97%), similar to that obtained with the reverse aqua regia digestion experiments. The high yields of MY-4 and MY-5 indicate that there is no significant Os lost during the digestion in HF-HCl-EtOH, or the dry-down that is required to drive off the SiF4, suggesting that the samples remained sufficiently reduced to prevent formation of volatile OsO4. 165

3.4 Advantages and potential disadvantages of microwave digestion for Os analysis Microwave digestion is a relatively simple method for sample dissolution and spike-sample equilibration for Os isotope analysis. Other than the microwave system itself, no extra costly consumables such as Carius tubes or dedicated equipment such as compressed gas and torches are required. This method can effectively extract Os from geological samples including silicates, and effectively equilibrates sample and spike Os in a relatively short time (<4 hours) compared to Carius tube digestion, and similarly low Os procedural blank levels are obtained. Microwave digestion of silicate rocks using HF is advantageous compared to leaching in reverse aqua regia alone, in that complete sample dissolution can be achieved. For some silicate rock samples, especially those that contain Os in structurally bound sites or as inclusions in silicate minerals, breakdown of the silicate phases is particularly important. The microwave technique can be successfully used to digest these sample with HF, which is not possible with either Carius tube or high pressure asher digestions. Additionally, for rock samples that effervesce with acid and produce high gas pressures, microwave digestion is safer than Carius tube digestion. Caution is required when digesting such samples by Carius tube, because during heating or opening of the tube, high internal pressures may exceed the strength of the tube. During microwave digestion, high pressure vessels can be used to monitor the pressure of the digestion process using the pressure sensor in the reference vessel. If the internal pressure exceeds a pre-defined pressure limit below the maximum for vessel operation, the heating system will be turned off, and then restarted only after the pressure decreases. In some situations, excess gas will be vented. This may result in loss of sample Os, but prevents explosions that might otherwise occur in a Carius tube or high pressure asher. Although microwave digestion has many advantages over other sample digestion methods, there are nevertheless some potential disadvantages of this method. During the period of the experiments, there were two occasions for which there was evidence that sample solutions leaked during or after microwave digestion. The high temperature rotor has a leak detector that measures acid vapor levels in the microwave chamber, but the leak detector did not indicate any significant acid leak during any of the microwave runs,

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which indicates that the leaks may have occurred after the rotor was removed from the microwave, when the rotor was placed in the freezer. This issue may be resolved by simply chilling the rotor in a refrigerator, rather than the freezer after digestion. It was originally intended that this study would evaluate the yields for Re in addition to Os by microwave digestion. However, it was immediately determined that the Re blanks are substantially (~10 times) higher than typical blanks by Carius tube digestion, and such high blanks are clearly prohibitive for many applications. Re procedural blanks by Carius tube digestion are typically ~2 pg, whereas the blanks for the initial microwave digestion experiments ranged from 20 pg to 50 pg. The high Re blanks are not due to the reagents, because low Re blanks (~2 pg) for Carius tube digestion were determined at the same time. Rather, it appears that the high blanks are contributed from the microwave digestion process itself. Several methods were attempted to clean the microwave vessels before digesting rock samples, including boiling them in 8N HNO3 or in aqua regia for 3 hours. However, these methods did not decrease the Re blank level below that obtained when the vessels were simply cleaned in situ with 8 N HNO3 heated in the microwave. It is possible that the Teflon used in the Milestone high temperature vessels has, for some reason, high levels of leachable Re that cannot be effectively removed. Until this issue is resolved, we would not currently suggest using microwave digestion to analyze samples for Re unless the abundances are sufficiently high (>1 ng) to overwhelm the blank.

4. Further studies The work completed in this study serves as a proof-of-concept for successful microwave digestion of geologic samples. Additional follow-up experiments that build on these results may further validate and improve the microwave digestion method. One important follow-up experiment includes comparison of the Os concentrations and 187Os/188Os ratios resulting from HF digestion and reverse aqua regia leaching alone to confirm the effectiveness of Os extraction from silicate samples in the microwave, for samples which may have significant Os bound in silicate phases or inclusion therein. In addition, to address the apparent issue of occasional venting of vessels in the freezer, it will be important to test the Os recovery after cooling the microwave rotor in a refrigerator only. This should reduce potential effects of differential contraction upon

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cooling, and minimize potential freeze-thaw deterioration of the vessels due to those related stress effects. Finally, additional experiments to assess the minimum time and temperature required for complete extraction of Os and spike-sample equilibration will be important in order to establish the most efficient and safest digestion method for high sample throughput situations.

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References Creaser, R.A., Papanastassiou, D.A., Wasserburg, G.J., 1991. Negative thermal ion mass spectrometry of osmium, rhenium, and iridium. Geochim. Cosmochim. Acta 55, 397-401. Frei, R., Villa, I.M., Nägler, T.F., Kramers, J.D., Przybylowicz, W.J., Prozesky, V.M., Hofmann, B.A., Kamber, B.S., 1997. Single mineral dating by the Pb-Pb step-leaching method: assessing the mechanisms. Geochim. Cosmochim. Acta 61, 393-414. Hassler, D.R., Peucker-Ehrenbrink, B., Ravizza, G.E., 2000. Rapid determination of Os isotopic composition by sparging OsO4 into a magnetic-sector ICP-MS. Chemical Geology 166, 1-14. Meisel, T., Moser, J., Fellner, N., Wegscheider, W., Schoenberg, R., 2001a. Simplified method for the determination of Ru, Pd, Re, Os, Ir and Pt in chromitites and other geological materials by isotope dilution ICP-MS and acid digestion. The Analyst 126 (3), 322-328. Meisel, T., Moser, J., Wegscheider, W., 2001b. Recognizing heterogeneous distribution of platinum group elements (PGE) in geological materials by means of the Re-Os isotope system. Fresenius Journal of Analytical Chemistry 370 (5), 566-572. Morgan, J.W., Golightly, D.W., Dorrzapf, A.F., 1991. Methods for the separation of rhenium, osmium and molybdenum applicable to isotope geochemistry. Talanta 38, 259-265. Nägler, T.F., Frei, R., 1997. “Plug in” Os distillation. Schweiz. Mineral. Petrogr. Mitt. 77, 123-127. Roy-Barman, M., 1993. Mesure du rapport 187Os/186Os dans les basalts et peridotites: contribution a la systematique 187Re/186Os dans le manteau. Thesis, Univ. Paris 7. Roy-Barman, M., Allègre, C.J., 1995. 187Os/186Os in ocean island basalts: tracing oceanic crust recycling in the mantle. Earth Planet. Sci. Lett. 129 (1-4), 145– 161. Schoenberg, R., Nägler, T.F., Kramers, J.D., 2000. Precise Os isotope ratio and Re-Os isotope dilution measurements down to the picogram level using multicollector inductively coupled plasma mass spectrometry. International Journal of Mass Spectrometry 197 (1-3), 85-94. Shirey S.B., Walker R.J., 1995. Carius tube digestion for low-blank rhenium-osmium analysis. Anal. Chem. 67 (13), 2136–2141. Suzuki K., Lu Q., Shimizu H., Masuda A., 1992. Determination of osmium abundance in molybdenite mineral by isotope dilution mass spectrometry with microwave digestion using potassium dichromate as oxidizing agent. Analyst 117, 1151-1156. Volkening, J., Walczyk, T., Heumann, K., 1991. Osmium isotope ratio determinations by negative thermal ionization mass spectrometry. Int. J. Mass Spectr. Ion Proc. 105, 147-159.

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Table 1: Comparison of microwave and Carius tube digestions Method 187Os/188Os Os (ppt) WAT13_1 CT 0.1325±0.0003 10.0±0.02 WAT13_2 CT 0.1321±0.0003 9.9±0.02 Average_1 0.1323±0.0006 10.0±0.07 WAT13_3 MW 0.1314±0.0003 10.5±0.03 WAT13_4 MW 0.1316±0.0005 9.4±0.04 WAT13_5 MW 0.1328±0.0003 9.7±0.02 WAT13_6 MW 0.1300±0.0004 8.9±0.03 Average_2 0.1314±0.0023 9.6±0.67 Blanks MB-1 MW 0.3pg MB-2 MW 0.3pg MB-4 MW 0.2pg MB-5 MW 0.3pg MB-6 MW 0.4pg MB-7 MW 0.6pg MB-8 MW 0.1pg MW designates microwave digestion, and CT Carius tube digestion. The 187Os/188Os ratios and Os concentrations of individual sample are reported with internal measurement errors (2ζ). Average_1 and Average_2 are mean values of data from Carius tube digestion and microwave digestion respectively, indicating ±2SD for each group data.

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Table 2: Osmium yields of microwave digestion Method % Os recovered MY-1 MW; spike back to MW 91% MY-2 MW; spike back to MW 95% MY-3 MW; spike back to CT 82% MY-4 MW (Ethanol + HF + HCl); MW; spike back to MW 98% MY-5 MW (Ethanol + HF + HCl); MW; spike back to MW 96% MY-6 MW; spike back to MW 97% MY-7 MW; spike back to MW 96% MW (Ethanol + HF + HCl): High pressure microwave program MW: High temperature microwave program CT: Carius tube

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Figure 1. The temperature and pressure change during heating of aqua regia in the microwave.

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Figure 1

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Figure 2. The temperature and pressure change during heating of HF + EtOH + HCl in the microwave.

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Figure 2

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Figure 3. Distillation system for Os separation. 3a is the Teflon coated aluminum heating block with sample wells, 3b is the heater and controller, 3c shows the sample vials connected with dual-port caps, 3d is the manifold, fish-tank air pump and filter, and 3e is the complete assembled system.

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Figure 3

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Figure 4. Osmium concentrations (4a) and 187Os/188Os ratios (4b) of sample aliquots digested by Carius tube (star symbols) and microwave (cycle symbols). The black solid line shows the average of the six samples and the dashed line shows the 2SD variation about the mean of these six samples.

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Figure 4 179

CHAPTER 5

Summary

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Azores basalts exhibit variable trace element and isotopic signatures that indicate derivation from distinct and heterogeneous mantle sources. Most of previous studies have focused on Sr-Nd-Pb isotopes which have limited ability to fully constrain the processes that might produce the heterogeneous mantle sources. The Os, Hf and Li isotope systematics of the Azores basalts in this dissertation have provided additional insights into the mantle sources of the basalts from the Azores archipelago. 187 188 Sub-chondritic Os/ Os and ΔεHf ≈0 of São Jorge and Terceira basalts, combined with large variations of 206Pb/204Pb and slightly elevated 87Sr/86Sr, indicates that these basalts reflect the FOZO signature characteristic of a deep-rooted mantle plume. 187 188 Sub-chondritic Os/ Os and negative ΔεHf (-1 ~ -2) of Faial basalts, combined with slightly elevated 206Pb/204Pb and high 87Sr/86Sr (up to 0.704), are interpreted to reflect a component in the Faial mantle source of recycled metasomatized mantle wedge with an EM signature. Stabilization and subsequent melting of an amphibole-bearing metasomatized mantle wedge can explain the relatively K2O-rich and low La/Nb signatures of the Faial basalts relative to those of Terceira and São Jorge. Pico basalts are likely produced by mixing between the EM component beneath Faial and the FOZO plume beneath São Jorge and Terceira. The Li isotope data for the Azores basalts combined with the new diffusion modeling in this study provide new constraints and perspective on the applicability of Li isotopes as a tracer of recycled material in the mantle. Despite large variations in radiogenic isotope systems, the Azores basalts exhibit limited variation in δ7Li that are within the range of global MORB. Although the δ7Li variation of the Central Group island basalts is limited, there are correlations between δ7Li and radiogenic isotopes that can help constrain the origin of their mantle sources. The δ7Li signatures of the Central Group island basalts support the conclusions based on the Os and Hf isotope systematics combined with Sr-Nd-Pb isotopes and trace elements. The slightly elevated δ7Li signatures in the Faial mantle source relative to Terceira and Pico are interpreted to represent mixing between a recycled mantle wedge component with slightly elevated subduction-related δ7Li and the FOZO-like Terceira source with a slightly lower but MORB-like δ7Li. The δ7Li signatures of the São Miguel basalts are slightly lower than those of the Central Group islands. The lack of anomalous δ7Li signatures or any correlation of δ7Li with radiogenic

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isotope signatures in the São Miguel basalts may reflect a source generated by recycling of underplated basalts rather than seawater altered and subduction dewatered crust, or a small (<<8 km) body of recycled seawater altered oceanic crust that was diffusively homogenized with the surrounding mantle over its 3Ga residence time. The new microwave digestion method for Os isotopic analysis developed in this dissertation has high recoveries of Os and low blank levels. Compared with previous digestion methods including Teflon bomb digestion, Carius tube digestion and high pressure asher (HPA-S) digestion, this new method requires less time (~4 hours vs. 48 hours), no costly consumables (e.g. Carius tubes), and no specialized equipment that is incompatible with the clean lab environment (e.g. gas cylinders and torches). In addition, this new method allows the use of HF to achieve complete digestion of silicate rock samples, which is prevented in the Carius tube and high pressure asher digestion methods due to the glass vessels employed in those digestion procedures. This project also involved the development of a modified “macro” distillation method for Os purification, based on that developed by Nägler and Frei (1997), which enables simultaneous purification of multiple (8 or more) samples. Under controlled conditions, the “macro” distillation method produces yields of >70%, and is easily employed and modified for a variety of geological and cosmochemical applications. In the future, additional follow-up experiments will be useful, including (1) comparison of the Os concentrations and 187Os/188Os ratios resulting from HF digestion versus reserve aqua regia leaching to confirm the effectiveness of Os extraction from silicate samples in the microwave, (2) experiments to address the apparent issue of occasional venting of vessels in the freezer by testing the Os recovery after cooling the microwave rotor in a refrigerator, and (3) experiments to assess the minimum temperature and time required for complete extraction of Os and spike-sample equilibration, to minimize internal vessel pressure and to enable maximum sample throughput.

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Reference Nägler, T.F., Frei, R., 1997. “Plug in” Os distillation. Schweiz. Mineral. Petrogr. Mitt. 77, 123-127.

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