NOBLE GASES AND HALOGENS IN ICELANDIC BASALTS

Transcription

1 NOBLE GASES AND HALOGENS IN ICELANDIC BASALTS A thesis submitted to the University of Manchester for the degree of Doctor of Philosophy in the Faculty of Engineering and Physical Sciences by Bridget M. Weston School of Earth, Atmospheric & Environmental Sciences The University of Manchester 2012

2 Contents List of figures 6 List of tables 8 Abstract 9 Declaration 10 Copyright statement 10 Acknowledgements 11 1 Chapter One Introduction Introduction Aims of study Thesis outline and publication status Chapter Two Literature review Noble gases as geochemical tracers The global noble gas dataset for the mantle Helium Neon Argon Krypton and Xenon Elemental abundances Mantle reservoirs Halogens as geochemical tracers Disentangling mantle signatures Crustal contamination Atmospheric contamination Cosmogenic Component resolution (Total word count including references: 60,443) 2

3 2.5.5 Degassing Fractionation during melt generation Models of the mantle Volatile recycling Iceland Noble gas and halogen studies of Iceland to date Chapter Three Experimental methods Sample collection and field relationships Sample preparation Noble gas mass spectrometry Equipment overview Mass spectrometer details Noble gas extraction line Sample analysis Sample crushers Noble gas purification and separation procedure Isotopic analysis Extraction line blanks and calibrations CO 2 and 40 Ar corrections Halogen analysis Sample preparation and irradiation Spectrometry Chapter Four Disequilibrium degassing model constrains the 3 He concentration of the MORB and OIB mantle sources Abstract Introduction

4 4.3 Model formulation Model comparison Modeling MORB and OIB sample suites Number of degassing steps Open or closed system conditions Diffusivity Solubility Temperature East Pacific Rise (EPR) sample suite Loihi sample suite Iceland sample suite Model results Discussion Summary and conclusions Appendix A Appendix B Chapter Five Noble gases and volatile recycling in Iceland's source mantle Abstract Introduction Geological background DICE sample area Sample analysis Results Component analysis Nature of additional components for the DICE area Source mantle composition Ne/ 22 Ne Heavy isotopic ratios He/ 22 Ne Elemental ratios

5 5.10 Discussion Conclusions Appendix to noble gas paper Introduction Noble gas results Xenoliths in DICE area Iceland s mantle beyond the DICE area Chapter Six The halogen composition of Icelandic basalts Abstract Introduction Samples Experimental techniques Results Release pattern Helium and argon Cl, Br, I and K Discussion Halogen fractionation Halogen mantle ratios and concentrations Halogens and mantle models Conclusions Chapter Seven Summary and future work Constraints on mantle models Volatile recycling Heterogeneity in Iceland s mantle References 209 5

6 List of figures Figure 1 Volatile recycling in the mantle? 14 Figure 2 Noble gas abundance patterns 25 Figure 3 Halogen compositions for terrestrial reservoirs 30 Figure 4 Corrections for atmospheric contamination 34 Figure 5 Iceland 48 Figure 6 Map of Iceland showing sample locations 54 Figure 7 Schematic sketch of tuya formation 61 Figure 8 Hloðufell 62 Figure 9 Sample MID5 showing a gabbroic xenolith 63 Figure 10 Tuffaceous sediment, Snæfellsnes Peninsula 64 Figure 11 VG5400 spectrometer and gas extraction system 68 Figure 12 Regression example for 40 Ar 75 Figure 13 Correction factors for CO 2 78 Figure 14 Corrected 20 Ne/ 22 Ne ratios from calibration shots 79 Figure 15 Schematic sketch of degassing model formulations 92 Figure 16 3 He concentrations and 3 He/ 22 Ne ratios after multi-step degassing 94 Figure 17 Schematic sketch of three-stage degassing model 97 Figure 18 Melt oversaturation and CO 2 concentrations during degassing 100 Figure 19 Model sensitivity to diffusivity parameters 103 Figure 20 Model sensitivity to diffusivity parameters 105 Figure 21 Calculated results and measured composition for the EPR sample suite 111 Figure 22 Calculated results and measured composition for the Loihi sample suite 112 Figure 23 Calculated results and measured composition for the Iceland sample suite 113 Figure 24 Calculated results and measured composition for the Loihi and Iceland sample suites 116 Figure 25 3 He/ 4 He ratio across Iceland 130 6

7 Figure 26 Three-isotope plot of 'DICE' Ne data 134 Figure Xe/ 36 Ar vs. 40 Ar/ 36 Ar for the 'DICE' sample area 138 Figure 28 3 He/ 22 Ne vs. 20 Ne/ 22 Ne for the 'DICE' sample area 141 Figure 29 He/Ne degassing for the 'DICE' sample area 145 Figure 30 He/Ne and He/Xe degassing for the 'DICE' sample area 146 Figure 31 Three component mixing for the 'DICE' sample area 147 Figure Ar/ 22 Ne vs. 20 Ne/ 22 Ne for the 'DICE' sample area 150 Figure 33 Ar/Ne mixing models 151 Figure 34 Xe/Ne mixing models 152 Figure 35 Degassing corrections for 36 Ar/ 22 Ne 153 Figure 36 Key non-radiogenic elemental noble gas trends 157 Figure 37 MORB/Solar Ne mixing 158 Figure 38 3 He/ 130 Xe vs. 129 Xe/ 130 Xe for the 'DICE' sample area 162 Figure 39 Three-isotope plot of Iceland's Ne data 167 Figure Xe/ 36 Ar vs. 40 Ar/ 36 Ar for Icelandic samples 175 Figure 41 Elemental abundances for xenoliths from the 'DICE' sample area 177 Figure 42 He/Ne and He/Xe degassing across Iceland 180 Figure 43 Sample locations across Iceland 185 Figure 44 Halogen release pattern during step heating 189 Figure 45 Halogen and K concentrations across Iceland 192 Figure 46 Correlations between Br, I, K and Cl concentrations 193 Figure 47 Halogen composition of Icelandic samples 194 Figure 48 Halogen concentrations across Iceland 200 7

8 List of tables Table 1 Noble gas isotopes and major production pathways within the mantle 17 Table 2 Noble gas isotopic ratios for the atmosphere 20 Table 3 Sample locations and descriptions 55 Table 4 'Glove-box' sample preparation test 67 Table 5 Calibration bottle composition 71 Table 6 Irradiation details for halogen analyses 81 Table 7 Model parameters and default values 89 Table 8 Noble gas data for sample MID1 135 Table 9 Goodness of fit for multi component fits to the 'DICE' sample 142 Table 10 Calculated ranges for the source mantle of the 'DICE' sample area 155 Table 11 Noble gas results for Icelandic sample 168 Table 12 Noble gas composition of gabbroic xenoliths from the 'DICE' sample area 179 Table 13 Irradiation details for halogen analyses 187 Table 14 Halogen step-heating data 195 8

9 Abstract Noble gases and halogens in Icelandic basalts A thesis submitted to The University of Manchester for the degree of Doctor of Philosophy by Bridget M. Weston, December 2012 Noble gas and halogen data from a suite of Icelandic samples are presented. Iceland combines hotspot volcanism, a spreading ridge and abundant subglacially erupted samples. This combination allows for samples that erupted under high enough pressures to retain a measurable mantle volatile content, and also display signatures representing interaction between ocean island basalt (OIB) and mid-ocean ridge basalt (MORB) mantle sources. Erupted samples used to determine the mantle s halogen and noble gas content have undergone a degassing process that can alter their volatile composition. An existing disequilibrium degassing model is developed with the modified model taking into account the evolution of the major volatiles over a multi-stage process and the different conditions present during magma ascent and quenching. The modified model allows substantially lower elemental noble gas ratios to be reached under disequilibrium conditions than allowed by the original model. Initial CO 2 concentrations, pressure, diffusivity, ascent rate and degree of disequilibrium are shown to be critical parameters for this model. Final degassed noble gas concentrations are most affected by the surface quenching stage of an eruption, whereas noble gas elemental ratios can be primarily determined during magma ascent. In applying this model to MORB and OIB sample suites, the 3 He/ 22 Ne ratio of the MORB source mantle is constrained to be lower than 4.4, similar to estimates for the OIB source mantle. Additionally the most straightforward match between the degassing model and OIB helium and neon data suggest the OIB source mantle has 3 He concentrations similar to or lower than the MORB source mantle. This finding requires a model for the OIB source mantle in which a high 3 He/ 4 He component is added to a helium-poor protolith. Noble gas studies are hampered by the large, isotopically atmospheric component typically found in Icelandic subglacial samples, which can swamp other signatures. Detailed analysis of a volatile rich sample from SW Iceland shows evidence for more than one contaminant component and that two component fits used incorrectly can produce misleadingly precise source mantle noble gas ratios. Multi component best fits to noble gas elemental ratios find that four components are present in samples from this region. These components are unfractionated air, fractionated air and a mantle component which shows some variation due to degassing. Combining the disequilibrium degassing model with component resolution allows limits to be placed on the source mantle composition for this sample. The light noble gas source composition is compatible with mixing between a solar ( direct nebula ) component and a MORB-like component. This direct nebula signature is at odds with an implanted signature seen in both Ne and Kr for the convecting mantle, and shows that both accretionary volatile origins must have contributed during the Earth s formation. The heavy noble gases show an elemental abundance pattern which is distinct from air and solar patterns, and trends towards seawater. This confirms the presence of a recycled volatile signature in Iceland s mantle but it is not possible to further constrain the origin of this signature. The Icelandic halogen data shows no evidence for significant fractionation during degassing or melt generation. Source estimates for the Br/Cl and I/Cl ratios for Iceland s plume are found to be (1.56±0.03) x 10-3 and (3.1±0.3) x 10-5, compatible with estimates for the MORB source mantle. Halogen source concentrations in central Iceland are found to be approximately three times higher than estimates for the convecting mantle and correlate with the regions of Iceland that show high 3 He/ 4 He ratios and high source water contents. This may indicate a recycled halogen signature associated with Iceland s proposed mantle plume. 9

10 Declaration No portion of the work referred to in this thesis has been submitted in support of an application for another degree or qualification of this or any other university or other institute of learning. Copyright statement 1) The author of this thesis (including any appendices and/or schedules to this thesis) owns certain copyright or related rights in it (the Copyright ) and s/he has given The University of Manchester certain rights to use such Copyright, including for administrative purposes. 2) Copies of this thesis, either in full or in extracts and whether in hard or electronic copy, may be made only in accordance with the Copyright, Designs and Patents Act 1988 (as amended) and regulations issued under it or, where appropriate, in accordance with licensing agreements which the University has from time to time. This page must form part of any such copies made. 3) The ownership of certain Copyright, patents, designs, trade marks and other intellectual property (the Intellectual Property ) and any reproductions of copyright works in the thesis, for example graphs and tables ( Reproductions ), which may be described in this thesis, may not be owned by the author and may be owned by third parties. Such Intellectual Property and Reproductions cannot and must not be made available for use without the prior written permission of the owner(s) of the relevant Intellectual Property and/or Reproductions. 4) Further information on the conditions under which disclosure, publication and exploitation of this thesis, the Copyright and any Intellectual Property and/or Reproductions described in it may take place is available in the University IP Policy (see in any relevant Thesis restriction declarations deposited in the University Library, The University Library s regulations (see and in The University s policy on Presentation of Theses. 10

11 Acknowledgements Firstly I would like to thank my supervisors Professor Chris Ballentine and Dr Ray Burgess for their support, guidance and encouragement during the course of this study. For introducing me to Iceland and helping me find my way around the mysteries of sample collection, thanks to Angela Walker. For donating a set of Icelandic samples for noble gas analyses, thanks to Alex Nichols. Many thanks to John and Zoë Bowers, Chris and Nikki Goodhew, Paul and Diane Weston, Chris and Liz Moorsom, Rosie Arnfield and Lesley Mundy for childcare support that allowed me time to write-up. Especial thanks goes to James Weston for unwavering support and putting up with me at every stage of the study. 11

12 1 Chapter One Introduction 1.1 Introduction Tomographic and seismic evidence have made a convincing case for a mantle where the 670 km seismic discontinuity is not a barrier to convection (for example, [R. D. Van Der Hilst et al., 1997]). However, fitting geochemical observations into the picture of whole mantle convection is still a work in progress. A variety of mantle reservoirs have been proposed on the basis of trace element and noble gas geochemistry, but the details of the extent, interaction and origin of these reservoirs have yet to be worked out [Anderson, 1998a; C. J. Ballentine et al., 2002; Hofmann, 1997; Stracke, 2005]. In particular, models of whole mantle convection predict that material which is recycled into the mantle at subduction zones can be preferentially stored in the deep mantle [Brandenburg and van Keken, 2007a]. The subduction and recycling of oceanic crust has already been used to explain the enrichment in incompatible elements seen in ocean island volcanism, which is thought to originate in the deep mantle [Hofmann, 1997]. However, the role of the volatile recycling in models of the convecting mantle is still open for investigation (Figure 1). Preferential volatile recycling into the deep ocean island mantle source may explain some of the much-discussed differences between the volatile geochemistry of ocean islands and shallow-origin mid-ocean ridges; for example the comparatively low 40 Ar/ 36 Ar ratios in ocean island volcanism [Brandenburg et al., 2008]. H 2 O content would also be affected which has implications for mantle viscosity, conductivity and dynamics, melting processes and seismic studies [Richard et al., 2002]. The noble gases show distinct abundance patterns in terrestrial and solar reservoirs and, 12

13 combined with their rare, inert nature, this makes them ideal geochemical tracers of mantle processes from the Earth s formation to the present day. The noble gas composition of mantle reservoirs also provides key constraints on mantle models [C. J. Ballentine et al., 2002]. Halogen studies of the mantle have shown their potential to trace mantle processes and combined noble gas and halogen analyses can provide valuable insights into the nature of the mantle [R. Burgess et al., 2002]. Despite the broad implications, the possibility of any noble gas recycling into the mantle has long been ignored as research indicated that a significant amount of these volatiles never got further than the subduction zone and its associated volcanism [Staudacher and Allegre, 1988]. This assumption was called into question by more recent studies of magmatic well gases in which the heavy noble gases have shown evidence for a seawater signature which is of mantle origin [Holland and Ballentine, 2006]. This could indicate that significant amounts of noble gases, and potentially other volatiles, are recycled into the mantle after all. Significant amounts of volatile recycling would alter the mantle s inventory of noble gases and halogens. The original volatile signature of the subglacial and submarine samples used to investigate the mantle s volatile composition can be altered by processes such as degassing, as well as atmospheric and crustal contamination. Taking full account of these processes is necessary if the original noble gas and halogen signatures of the source mantle are to be determined. Progress has been made in developing full disequilibrium degassing models for the noble gases, but the details of applying these models to specific sample suites requires further investigation [Gonnermann and Mukhopadhyay, 2007]. The combination of mantle phenomena involved in Iceland s volcanism make it an ideal place to study volatile recycling and the interaction of different mantle reservoirs: Iceland is situated on a hotspot which is thought to be due to a mantle plume which samples the deep mantle. It is also situated on the spreading mid-atlantic ridge which samples the depleted (convecting) mantle. It has the added advantage of prolific subglacial samples, which are an obvious choice for volatile studies as ice thicknesses 13

14 up to 1 km will provide higher than atmospheric eruption pressures, leaving samples relatively undegassed. The noble gas picture for Iceland can be complimented by analyses of the halogen composition, which is little studied and has not been constrained for central Iceland. 33 km Volatile Recycling? CORE CORE 2900 km Figure 1: In the whole mantle convection model, recycled material ends up concentrated in the deep mantle; this is then the source mantle for ocean island hotspot volcanism. 1.2 Aims of study The aim of this study is to collect a new suite of subglacial samples from across Iceland and determine the noble gas and halogen composition of these samples using noble gas mass spectrometry techniques. The study then aims to use this data to constrain the noble gas and halogen composition of Iceland s source mantle through investigation and quantification of syn-eruptive and post-eruptive processes such as degassing and atmospheric contamination. Constraints on the volatile composition of Iceland s mantle can then be related to existing models of the mantle. In particular, this 14

15 study aims to test the prediction that a recycled signature from subducted slabs should play a key role in the deep mantle reservoir thought to provide the source of Iceland s proposed mantle plume. 1.3 Thesis outline and publication status Chapter two consists of a review of the literature on noble gas and halogen studies of the mantle. It also includes a brief geological and geochemical background for Iceland. Chapter three details the sample collection undertaken in Iceland and the experimental methods used to determine the noble gas and halogen composition of these samples. Chapter four presents a modified noble gas disequilibrium degassing model and explores the consequences of this model for noble gas mantle models. Chapter four is prepared for submission to the Journal of Geophysical Research (Solid Earth). Chapter five presents and discusses the noble gas data from Icelandic samples, with a particular focus on a region in SW Iceland which produces unusually volatile-rich samples. Chapter six presents and discusses the halogen composition of Icelandic samples. Chapters five and six are prepared for submission to Chemical Geology or Earth and Planetary Science Letters. Chapter five includes an appendix not intended for publication. A summary of the main conclusions of this thesis and suggestions for future work are given in Chapter seven. 15

16 2 Chapter Two Literature review 2.1 Noble gases as geochemical tracers Along with geophysical tools such as seismic studies, geochemical studies of oceanic basalts provide a key insight into the dynamics of the mantle. In such studies the noble gases are invaluable geochemical tracers. Their rarity means mantle signatures are less easily confused by surface contamination; their inert nature makes them ideal tracers of distinct noble gas reservoirs and allows for the identification of ancient signatures potentially unchanged since the Earth s formation. The noble gases include a number of radiogenic isotopes with known production pathways (see Table 1). These can be used as dating tools [Merrihue and Turner, 1966], tracers of parent isotopes and to constrain the timing of events during the Earth s history [Kunz et al., 1998; G. Turner, 1965]. The non-radiogenic noble gas isotopes are not produced within the mantle and so have tracked mantle processes since the Earth s formation [Claude J. Allegre and Turcotte, 1986; C. J. Allegre et al., 1983; Chris J. Ballentine and Barfod, 2000; Chris J. Ballentine et al., 2005b; E T Dixon et al., 2000; D. Harrison et al., 1999; M. D. Kurz et al., 1982; M. Moreira et al., 2001; Samuel Niedermann et al., 1997; Porcelli and Ballentine, 2002; Sarda et al., 2000]. 16

17 Table 1: Noble gas isotopes and major production pathways within the mantle [D W Graham, 2002] Isotope Significant mantle production pathways 3 He None 4 He Radiogenic production from 235 & 238 U and 232 Th decay 20 Ne LIGHT None 21 Ne Nucleogenic production from particle reactions with 18 O, 24 Mg 22 Ne None 36 Ar None 38 Ar None 40 Ar Radiogenic production from 40 K 78 Kr None 80 Kr None 82 Kr None 83 Kr Minor production from 238 U fission 84 Kr Minor production from 238 U fission 86 Kr Minor production from 238 U fission HEAVY 124 Xe None 126 Xe None 128 Xe None 129 Xe Produced from 129 I decay ( 129 I is a short-lived, extinct isotope) 130 Xe None 131 Xe Produced from 244 Pu (extinct) and 238 U fission 132 Xe Produced from 244 Pu (extinct) and 238 U fission 134 Xe Produced from 244 Pu (extinct) and 238 U fission 136 Xe Produced from 244 Pu (extinct) and 238 U fission 17

18 2.2 The global noble gas dataset for the mantle Solar and meteoritic noble gas compositions provide potential starting points for the Earth s noble gas reservoirs (key components are shown in Table 2. Comparing these noble gas abundances and isotopic ratios with those found in the atmosphere and mantle reservoirs allows the story of the mantle from formation, through atmosphere formation to the present day to be gradually unravelled [C. J. Ballentine and Holland, 2008; Chris J. Ballentine et al., 2005a; Farley and Neroda, 1998; Hiyagon et al., 1992; G. Holland et al., 2009a; Bernard Marty, 1989; 2012; Bernard Marty and Alle, 1994; Bernard Marty and Dauphas, 2002; Bernard Marty and Meibom, 2007; Pepin and Porcelli, 2002; Pepin and Porcelli, 2006; Pujol et al., 2011; Shaw et al., 2001; Stuart, 1994; Yokochi and Marty, 2005; Youxue and Zindler, 1989]. The noble gas composition of the atmosphere is readily available for study and has been determined to high precision [J-Y Lee et al., 2006; Porcelli et al., 2002]. Solar and planetary noble gas compositions are studied through meteorites, analysis of the solar wind (direct or via solar wind irradiated meteorites), planetary probes such as Galileo and lunar samples [Anders and Grevesse, 1989; Heber et al., 2009; Porcelli et al., 2002; Rainer Wieler, 2002; R. Wieler et al., 2006]. Noble gas studies of the mantle mainly focus on oceanic or subglacial basalt glasses as the high eruption pressures in these environments lead to samples which have retained a measurable mantle volatile content. The fast-quenched glassy rims of pillow basalts are ideal as they are undegassed relative to the crystallised basalt and the glass structure is also efficient at trapping vesicles containing mantle volatiles. Mantle noble gases can also be trapped in fluid inclusions contained in minerals such as olivine [E T Dixon, 2003; Jens Hopp and Ionov, 2011]. Volatile concentrations are often lower, making detection more challenging. Possible fractionation during the formation of these inclusions must also be taken into account [Baker, 2008]. Coated diamonds have been shown to include fluid inclusions which contain mantle volatiles in detectable quantities [Ray Burgess et al., 2009; C. Gautheron et al., 2005; Johnson et al., 2000; Ozima and Zashu, 1991]. Magmatic CO 2 natural gas fields, due to larger sample volume availability, are a 18

19 sample resource that has recently provided an unprecedented improvement in the accuracy and detail of the noble gas composition of the upper mantle [C. J. Ballentine and Holland, 2008; Chris J. Ballentine et al., 2005; Caffee, 1999; Holland and Ballentine, 2006; Greg Holland et al., 2009b]. 19

20 Table 2: Noble gas isotopic ratios for the atmosphere [J-Y Lee et al., 2006; Porcelli et al., 2002], solar wind [Heber et al., 2009] and chondritic reservoirs.the He, Ne and Ar chondritic isotopic ratios are for the Ne-B component of carbonaceous meteorites [Black, 1972]; Kr and Xe are the average carbonaceous chondrite composition [Greg Holland et al., 2009b; Mazor et al., 1970; Pepin, 2003]. The value for 129 Xe/ 130 Xe is an average for CI and CM carbonaceous chondrites, the approximate range is [Mazor et al., 1970; Pepin, 2003]. Isotope Ratio Atmosphere Solar Wind Chondritic/Ne-B 3 He/ 4 He (1.399 ± 0.013) (4.64 ± 0.09) (3.9 ± 0.3) x 10-6 x 10-4 x Ne/ 22 Ne 9.80 ± ± ± Ne/ 22 Ne ± ± ± Ar/ 36 Ar ± ± ± Ar/ 36 Ar ± Kr/ 84 Kr ± ± ± Kr/ 84 Kr ± ± ± Kr/ 84 Kr ± ± ± Xe/ 130 Xe ± ± Xe/ 130 Xe ± ± ± Xe/ 130 Xe ± ± ± Xe/ 130 Xe ± ± ± Xe/ 130 Xe ± ± ±

21 2.2.1 Helium Although small amounts of 3 He can be produced on the terrestrial surface through cosmic ray spallation, there are no significant production pathways for 3 He within the mantle. 3 He/ 4 He ratios are considerably higher in the mantle than in the atmosphere. The 3 He/ 4 He atmospheric ratio is 1.399x10-5. Ratios between ~ 4 and 50 times higher than this characterise the mantle [D W Graham, 2002; Matsumoto et al., 2008; Stuart et al., 2003]. As helium has a residence time in the atmosphere of 1 to 10 million years, and is not recycled into the mantle, the high mantle 3 He/ 4 He ratios must be due to primordial 3 He that has been trapped within the mantle since the Earth s formation [Craig and Lupton, 1976; Lupton and Craig, 1975; Mamyrin et al., 1969; I N Tolstikhin, 1975]. This makes 3 He an important tracer of mantle involvement as mantle degassing is the most significant source of 3 He on Earth. The origin and location of this trapped 3 He component in the mantle remains controversial [C-J Allegre et al., 1993; C. J. Allegre et al., 1994; Anderson, 1998a; b; C. J. Ballentine and Holland, 2008; Porcelli and Ballentine, 2002; Stuart, 1994; Peter E. van Keken and Ballentine, 1998] see Mantle models section for a more detailed discussion Neon The atmospheric 20 Ne/ 22 Ne ratio is 9.80 ± 0.08 whereas mantle ratios range from atmospheric to considerably higher values, often correlating with 21 Ne/ 22 Ne ratios that vary from close to atmospheric (0.029 ± ) to higher, more radiogenic values [Porcelli et al., 2002; Sarda et al., 1988]. This is generally interpreted as mixing between a high 20 Ne/ 22 Ne and 21 Ne/ 22 Ne mantle source and an atmospheric component [Honda et al., 1993a]. The difference between these two components has implications for the source of the Earth s volatiles and the relationship between mantle volatiles and the atmosphere [Porcelli and Ballentine, 2002]. As neon recycling into the mantle is not expected to be significant [Holland and Ballentine, 2006], the observed atmospheric component seen in mantle samples is most likely due to contamination added during sample preparation and ratios measured in oceanic basalts can therefore 21

22 only give a minimum value for the mantle. Although production pathways for 20 Ne and 22 Ne do exist, neither isotope is produced in significant quantities within the mantle and so mantle ratios should reflect the Earth s primordial composition [C-J Allegre et al., 1993; C. J. Allegre et al., 1994; Chris J. Ballentine et al., 2005a]. Identifying the origin of this primordial neon is important not only for models of the early Earth but also to enable corrections for atmospheric contamination added during sample preparation to be made for the heavy noble gases (see Atmospheric Contamination section). Two possible origins for the high 20 Ne/ 22 Ne source are considered: Solar wind measurements give a ratio close to 13.8 [Heber et al., 2009; Rainer Wieler, 2002] which could have been incorporated into the Earth directly from a planetary nebula atmosphere [Honda et al., 1991; Honda et al., 1993b]; the alternative is an implanted origin [Chris J. Ballentine et al., 2005a; Podosek et al., 2000; Mario Trieloff et al., 2000] which may explain ratios closely matching those found in solar-wind irradiated meteorites and the lunar regolith ( 20 Ne/ 22 Ne=12.5, termed Ne-B). Since ratios as high as ± 0.32 have been found in some Icelandic samples [D. Harrison et al., 1999; Mukhopadhyay, 2012], a direct nebula origin for all mantle neon was proposed. Only a few analyses gave results higher than 12.5 and it was originally suggested that these results represent the statistical extremes of a normal distribution centred closer to 12.5 (Trieloff s reply to [C. J. Ballentine et al., 2001]). It is also possible that the neon analysis that found the highest value of ± 0.32 was compromised by the exposure of the spectrometer to spiked neon [D. Harrison, 2003]. Other step-wise crushing analyses of samples from the same area of Iceland have found ratios approaching a maximum of 12.5 [Parai et al., 2009; Mario Trieloff et al., 2000]. However the latest high quality results from this sample show ratios consistently higher than 12.5 (three separate measurements gave 20 Ne/ 22 Ne = (±0.03), (±0.01) and (±0.06) [Mukhopadhyay, 2012]. Measurements up to 13.0 (±0.2) have also been found in carbonatites from the Kola peninsula plume [Yokochi and Marty, 2005]. These ratios indicate a direct nebula origin for neon for at least some mantle reservoirs. Results of neon analyses of basalt samples are often obscured by large amounts of 22

23 atmospheric contamination: Studying the noble gas composition of magmatic CO 2 from wells in New Mexico avoids this problem and a mantle 20 Ne/ 22 Ne ratio of (±0.04) was found [Holland and Ballentine, 2006]; enhanced 22 Ne production, solar-air mixing and mass-fractionation are ruled out as possible mechanisms for reducing a direct nebulacomponent to 12.5 and an implanted origin for these primordial mantle gases is considered the most likely explanation [Chris J. Ballentine et al., 2005a] Argon 40 Ar/ 36 Ar ratios in the mantle are more radiogenic than in the atmospheric value of [M. Moreira, 1998] due to 36 Ar outgassing and the production of 40 Ar from 40 K over time in the mantle [A. Jambon et al., 1985], with values up to 40,000 found in individual vesicles of a basaltic sample from the mid-atlantic Ridge [P Burnard et al., 1997]. Finding evidence for a solar signature in the non-radiogenic 38 Ar/ 36 Ar ratio has proved problematic due to the complications of air contamination and low 38 Ar abundances. However, a few studies have found evidence for 38 Ar/ 36 Ar ratios that are lower (i.e. closer to solar) than the atmospheric value of [Pepin, 1998] although the conclusiveness of the data is debated [Kunz, 1999] and [Raquin et al., 2008] conclusively showed that mantle 38 Ar/ 36 Ar ratios were not distinct from atmospheric values. 38 Ar/ 36 Ar ratios distinct from atmospheric values were also not found in a high precision study of mantle well gases [C. J. Ballentine and Holland, 2008] Krypton and Xenon Isotopes of Kr generally show indistinguishable isotopic ratios in the mantle and the atmosphere, although resolving any small differences is hard to achieve due to the low abundances of these isotopes. Nonradiogenic Xe isotopes pose a similar resolution problem due to small abundances, although small excesses in Xe/ 130 Xe have been found in some magmatic CO 2 samples [Caffee, 1999] that could indicate the presence of solar Xe. Recently, the high volatile abundances in CO 2 well gases have allowed 23

24 high precision Xe and Kr ratios to be determined [G. Holland et al., 2009b]. Holland et al. (2009) show that even using these high resolution results it is not possible to distinguish between direct nebula and implanted signatures from Xe ratios. In contrast the Kr ratios showed a deviation from atmospheric ratios and are consistent with an implanted origin for the upper mantle s krypton. Radiogenic 129 Xe/ 130 Xe and 136 Xe/ 130 Xe ratios are generally higher in the mantle than the atmosphere [M. Moreira, 1998]. 129 I and 244 Pu are extinct isotopes which contributed daughter Xe isotopes to the mantle ( 129 Xe from 129 I and 131,132,134,136 Xe from 244 Pu). The abundance patterns of these daughter isotopes allow the closure age of the Earth to xenon loss to be determined [Porcelli and Ballentine, 2002; Wetherill, 1975] Elemental abundances Noble gas elemental trends also differ between terrestrial, solar and meteoritic reservoirs, with the Earth showing a general depletion relative to solar wind abundances (see Figure 2) [Pepin and Porcelli, 2002]. In principle, elemental noble gas ratios can be used in conjunction with isotopic ratios to trace signatures within the mantle. However, mantle elemental ratios are not easy to determine from surface samples as they are much more susceptible to fractionation during processes such as melting, degassing and crystallisation than isotopic ratios. Some progress can be made by starting from the mantle production ratios for 4 He/ 21 Ne. 4 He is produced via 235/238 U and 232 Th decay and 21 Ne is produced via ( 4 He) particle reactions with 18 O and 24 Mg. Hence these two isotopes have strongly correlated production pathways and the mantle ratio should tend to the production ratio of 21 Ne/ 4 He = 4.5x10-8 [Porcelli and Ballentine, 2002]. 40 Ar is produced from 40 K and 136 Xe from 238 U allowing mantle production ratios for 4 He/ 40 Ar (1-5) and 4 He/ 136 Xe (4.5 x 10 8 ) to be determined from estimates of the parent isotope abundances within the mantle [D W Graham, 2002]. Comparing measured ratios with these production ratios can allow the degree of fractionation a particular sample has undergone to be gauged. Calculations based on these production ratios have led to various estimates for the 3 He/ 22 Ne, 22 Ne/ 36 Ar and 24

25 130 Xe/ 36 Ar ratios within the mantle, although such calculations rely on a range of assumptions such as the number of components present in a sample [E T Dixon, 2003; D W Graham, 2002; Honda and McDougall, 1997; Bernard Marty et al., 1998]. A few samples fortuitously show measured ratios close to mantle production ratios which may indicate they are relatively unfractionated. These include popping rock (a rare dredged sample from the mid-atlantic ridge with such high volatile contents that vesicles could be heard to pop when the sample was exposed to atmospheric pressures) [Javoy and Pineau, 1991; Sarda and Graham, 1990]; some mantle CO 2 well gases [C. J. Ballentine, 1997] and a glassy pillow basalt from SW Iceland [D. Harrison et al., 1999]. A detailed understanding of these rare samples origin and formation is key as they are often used as representative of their particular mantle reservoirs. Seawater Convecting Mantle (Bravo Dome) CI Chondrites Solar Wind AIR Figure 2: Non-radiogenic noble gas abundance patterns ( 3 He, 22 Ne, 36 Ar, 84 Kr and 130 Xe) for air, seawater, the solar wind, CI chondrites and an estimate for the convecting mantle from Bravo Dome mantle CO 2 data [Holland and Ballentine, 2006; Ott, 2002; Pepin and Porcelli, 2002]. 25

26 2.3 Mantle reservoirs In the history of applying mantle noble gas data to mantle models, the differences in the observed compositions of ocean island basalts (OIBs - e.g. Hawaii) and mid-ocean ridge (MORBs) have been a key tool. The MORB data set is relatively homogenous when compared to the OIB set and is thought to predominantly sample the upper mantle, although other components are also detectable. OIBs generally appear to be a more complex mix of the different components and display a wider range of noble gas compositions [D W Graham, 2002; Mourão et al., 2012]. It has been argued that the homogeneity of MORBs is due to the greater abundance of MORB studies, meaning that variations have been averaged out over the larger data set [Anderson, 2007]. However, the distinctive features of OIBs have also been explained as indicative of a different mantle source, often associated with potential mantle plumes (see review by [Hawkesworth and Scherstén, 2007]). The OIB picture is complicated by the possibility that different ocean island locations may represent different mantle phenomena. The major differences between these two datasets are summarised below. MORBs show 3 He/ 4 He ratios averaging 8R A (R A is the atmospheric 3 He/ 4 He ratio of 1.399x10-6 ; ratios between ~ 4 and 50 times higher than this characterise the mantle): For OIBs, values are more variable [M. D. Kurz et al., 1982] with ratio ranges from 5.5R A (Canary Islands) up to the highest values found in Baffin Island basalts of 50R A [D W Graham, 2002; Stuart et al., 2003] with a general trend towards 3 He/ 4 He ratios higher than found in MORBs [C. J. Allegre et al., 1983; Macpherson et al., 1998]. Both MORBs and OIBs show 20 Ne/ 22 Ne ratios higher than the atmospheric value. No samples with 20 Ne/ 22 Ne > 12.5 have yet been found for the MORB source mantle, whereas a few OIBs have shown 20 Ne/ 22 Ne > 12.5 [D. Harrison et al., 1999; Mukhopadhyay, 2012; Yokochi and Marty, 2004]. Whether or not this ratio is the same (or even of the same origin) for the OIB and MORB mantle sources is still a matter of debate (see Mantle Models section below). 21 Ne/ 22 Ne ratios are generally lower in 26

27 OIBs than MORBs. This has been interpreted as representing a more dominant radiogenic component in the MORB source mantle [E T Dixon, 2003; E T Dixon et al., 2000; Mark D. Kurz et al., 2009; Parai et al., 2009; Mario Trieloff et al., 2000]. Measured 3 He/ 22 Ne ratios, and 3 He/ 22 Ne ratios calculated from measured isotopic ratios, vary widely with some studies suggesting a systematic difference between OIBs and MORBs [E T Dixon, 2003; D W Graham, 2002; Honda and McDougall, 1998]. The varying ratios may represent mantle heterogeneity but the issue is complicated by the possibility of mixing between melts from different mantle reservoirs, which may also have experienced He/Ne fractionation both during melt formation and during degassing [E T Dixon, 2003]. There is still much work to be done on unpicking the various possible reasons for the observed 3 He/ 22 Ne ratios. 40 Ar/ 36 Ar ratios are generally lower in OIBs than MORBs [C. J. Allegre et al., 1983] with the highest MORB ratio of 40,000 [P Burnard et al., 1997] contrasting with the highest measured OIB ratios of ~ 7000 [D. Harrison et al., 1999; Mukhopadhyay, 2012]. This is usually interpreted as an inherent difference between the MORB and OIB mantle reservoirs K/Ar ratio, but the cause of the difference is debated. While 129 Xe/ 130 Xe excesses are well-established in MORB samples, excesses in OIBs are more uncertain; some OIB data from Hawaii and Iceland have shown more radiogenic 129 Xe/ 130 Xe than the atmosphere, although values are still lower than those found in MORBs [D. Harrison et al., 1999; R. K. Mohapatra et al., 2009; Mukhopadhyay, 2012]. Different geochemical reservoirs within the mantle are also indicated in other isotopic systems: Based mainly on Rb/Sr and Sm/Nd isotopic systems, the various mantle sources are divided into different components [Anderson, 2007; Hart et al., 1992]. The most generic geochemical component is termed the depleted mantle (DMM - [Workman and Hart, 2005]); this reservoir is considered to be synonymous with the MORB mantle source (although other components are also detectable in some MORBs) and is assumed to complement the enriched continental crust. The HIMU component represents mantle with a high (high compared to bulk Earth estimates) where = 27

28 238 U/ 204 Pb and was originally associated with geochemical signatures found in certain ocean island basalts such as St. Helena [Zindler and Hart, 1986]. Since a range of radiogenic lead isotopic ratios are found in both MORBs and OIBs, [Stracke, 2005] tightened this definition to the most radiogenic lead isotope ratios coupled with low 87 Sr/ 86 Sr ratios. This component has been explained as recycled oceanic crust [Brandenburg et al., 2008; Hauri and Hart, 1993; Hofmann et al., 1986]. The mantle component termed FOZO [Hart et al., 1992] displays more radiogenic lead isotope ratios but also includes high 3 He/ 4 He, 143 Nd/ 144 Nd and low (though higher than HIMU) 87 Sr/ 86 Sr; it has been suggested that this component represents a deep mantle source reservoir although recycled origins have also been suggested [Stracke, 2005]. The final categories are EM1 and EM2 ( enriched mantle components primarily defined in terms of Sr, Nd and Pb isotopes) which represent different enriched components associated with OIBs; EM2 has been interpreted as originating from recycled continental crust whereas EM1 is harder to place and could originate from recycled ancient sediments or sub-continental lithosphere [Hofmann, 1997]. 2.4 Halogens as geochemical tracers The data set characterising the halogen composition of the convecting mantle is fairly limited compared to the noble gas data set [Deruelle et al., 1992; Albert Jambon et al., 1995; Mark A. Kendrick et al., 2012; Schilling et al., 1980]. Very few studies of OIB halogen compositions have been carried out [Unni and Schilling, 1978]. The small abundances in oceanic basalts have made studies difficult using conventional techniques. However a technique based on the 40 Ar-K dating method has allowed much lower halogen concentrations to be determined [R. Burgess and Turner, 1995; R. Burgess et al., 2002; Ray Burgess et al., 2009; Mark A. Kendrick, 2012; G. Turner, 1965]. The technique involves sample irradiation which converts the halogens to noble gas isotopes: 37 Cl(n,γ) 38 Cl(β) 38 Ar 28

29 79 Br(n,γ) 80 Br(β) 80 Kr 127 I(n,γ) 128 I(β) 128 Xe This allows noble gas spectrometers to determine halogen concentrations alongside noble gas concentrations. Like the noble gases, the halogens are incompatible during melting processes, meaning the halogen content of oceanic basalts is a good representation of the original mantle halogen composition. Halogen degassing is not thought to be significant at pressures above 1MPa, meaning much of the oceanic and subglacial basaltic sample set is considered undegassed in its halogen composition [Edmonds et al., 2009]. However comprehensive halogen degassing models are still a work in progress [Aiuppa et al., 2009; Baker and Balcone-Boissard, 2009]. Unlike the noble gases, the halogens are far from inert and are therefore more often employed to trace processes within the Earth, particularly mantle mixing, convection and recycling, than looking for primordial signatures. In particular, Iodine has a high affinity for organics and can therefore potentially track sediment recycling [Deruelle et al., 1992]. Mantle degassing over geological time can account for the Cl and Br in the Earth s surface reservoirs, the majority of which reside in the oceans (Cl and Br) and the crust (Br and I) [Muramatsu and Wedepohl, 1998; Schilling et al., 1978]. The composition of MORB glasses has been characterised by a number of studies [Deruelle et al., 1992; Dreibus et al., 1979; Albert Jambon et al., 1995; Schilling et al., 1980]. Variations in halogen composition with location in Iceland show that OIBs potentially carry a different halogen signature to MORBs, with higher halogen concentrations associated with Iceland s OIB signature [Unni and Schilling, 1978]. Using the noble gas technique, studies have been carried out on diamonds from a wide range of regions [R. Burgess and Turner, 1995; R. Burgess et al., 2002; Ray Burgess et al., 2009; Johnson et al., 2000]. A component similar to the MORB halogen composition is often identifiable and some samples have also shown the involvement of mantle processes such as crystallisation and mantle melting [Johnson et al., 2000]. Figure 3 summarises the current data on the halogen composition of the Earth s reservoirs. A combined picture of the noble gas and halogen composition of mantle 29

30 samples can provide a complementary picture of both the mantle s origins and the ongoing processes of its evolution. Figure 3: Halogen compositions for various terrestrial reservoirs [Deruelle et al., 1992; Albert Jambon et al., 1995; Kastner et al., 1990; Mahn and Gieskes, 2001; Martin et al., 1993; Muramatsu et al., 2001]. 2.5 Disentangling mantle signatures The majority of the noble gas and halogen mantle data are based on analysis of volcanic samples. The volatiles in basaltic and diamond samples are generally released into a high vacuum by step heating, crushing, laser ablation or a combination of these techniques. Crushing techniques are advantageous for basaltic glasses as the volatiles trapped in vesicles are easily released by this technique but noble gases within the glass matrix, which may be contaminants, are not. Following the appropriate extraction 30

31 technique, high resolution mass spectrometry using gas source noble gas spectrometers then allows the low noble gas abundances in the sample to be analysed. However, as such a sample journeys from the mantle to the surface, various processes occur which can change its volatile composition. These must be corrected for before the sample composition can be used to represent the original mantle composition and any recycled signatures can be identified Crustal contamination Unwanted crustal noble gas signatures can be added to melt in the magma chamber or during eruption. This problem can be avoided by choosing samples from a volcanic region where crustal contamination is not a significant factor; certain isotopic systems such as Os and 18 O are good indicators of the extent of crustal contamination [Prestvik et al., 2001]. This geochemical evidence can be combined with knowledge of the volcanic region s geophysics to minimise the risk of significant contamination. Crustal signatures can also be indicated by neon isotopic trends as an air/crust mixing trend is distinct from air/mantle mixing trends [Chris J. Ballentine et al., 2005a; S. Niedermann and Bach, 1998]. Alternatively, if the nature of the contaminants is well known or deducible from other isotopic systems, it can be resolved from other components [C. J. Ballentine, 1997; Chris J. Ballentine et al., 2005a] Atmospheric contamination Accurately compensating for late atmospheric contamination and distinguishing late contamination from an atmospheric-like source component are crucial when looking for mantle signatures such as recycled noble gases. Both basalt glasses and phenocryst samples show some signs of contamination from an atmospheric-like component [Chris J. Ballentine and Barfod, 2000; Farley and Craig, 1994; Bernard Marty and Ozima, 1986; Patterson et al., 1990; Raquin et al., 2008]. This can be seen most clearly in the 20 Ne/ 22 Ne and 40 Ar/ 36 Ar ratios which can vary from atmospheric to much higher 31

32 values over a single sample data set, or even within different vesicles of the same sample [P Burnard et al., 1997]. [Chris J. Ballentine and Barfod, 2000] discuss the various mechanisms that have been put forward to explain these air-like signatures: Seawater addition to the magma chamber; seawater contamination on eruption; diffusive input of seawater-derived noble gases and sorption of air-derived noble gases on fresh basalt surfaces. They argue that a sea-water derived signature added before eruption should generally be distributed uniformly within a sample, but analysis of a single sample can produce a large range of, for example, 20 Ne/ 22 Ne ratios [D. Harrison et al., 1999]. An elementally air-like signature is seen in both sub-glacial and sub-sea samples. This argues against mechanisms that place the addition during eruption as these should produce a seawater derived signature only in sub-sea samples. Ballentine and Barford s conclusion is that the source of atmospheric contamination is therefore modern air which may well be added to samples during handling and preparation. A mantle origin from recycled air has also been proposed for this atmospheric component [Sarda, 2004; Sarda et al., 1999] but as this component is elementally unfractionated from modern air; a recycling mechanism which puts unfractionated atmospheric noble gases into the mantle is hard to conceive. Assuming this component is an unwanted contaminant, there are two possible methods of correcting for this particular contamination effect: The first assumes that the 20 Ne/ 22 Ne ratio in the mantle is known (it is either assumed to be of direct nebula origin , or Ne-B ). The difference between the assumed initial 20 Ne/ 22 Ne ratio and the measured ratio then gives an indication of the amount of atmospheric contamination present [Farley and Poreda, 1993] (Figure 4). Other unknown neon ratios can then be corrected for using this quantity as follows: Equation 2.1: Correction for atmospheric contamination from [D W Graham, 2002]. f 22 is the proportion of mantle-derived Ne in a sample and S, E, M and A refer to solar, extrapolated, measured and atmospheric respectively Ne/ Ne ( Ne/ Ne Ne/ Ne ) / f E M A Ne/ 22 Ne A 32

33 Ne/ Ne Ne/ f M Ne/ NeS Ne/ 22 Ne Ne A A An alternative method is to assume that all the 36 Ar in the sample is atmospheric in origin and use the quantities measured in the sample to give the amount of atmospheric contamination [D W Graham, 2002]. The assumption that mantle 36 Ar is negligible is a significant disadvantage of this technique, especially when searching for a recycled mantle signature which would be characterised by the addition of 36 Ar to the mantle. It has also been noted that the more atmospheric-like signatures tend to be released in the early stages of crushing or heating; analysing samples over several steps can therefore allow a better estimate of the uncontaminated composition to be obtained. Detailed studies of single pillow basalts have shown that careful selection of the central portion of the glassy pillow rind can potentially minimise this contaminant [Kumagai and Kaneoka, 1998; 2003]. A few studies have noted that many oceanic Icelandic basalts and mid-ocean ridge popping rock show evidence of an additional component which looks isotopically like air but is elementally fractionated. The origin and prevalence of this component requires further study but its presence has the potential to complicate correction strategies which assume two component mixing [P Burnard et al., 2003; D. Harrison, 2003]. 33

34 Figure 4: An example of samples showing neon ratios which contain a mixture of atmospheric and mantle neon. The atmospheric component can be corrected for using a known mantle value (see Equation 2.1) Cosmogenic Surface samples are also open to the possibility of cosmic ray spallation which can produce noble gas isotopes and can particularly affect rare isotopes such as 3 He and 21 Ne. This problem can be largely avoided by careful choice of sample as production is not significant at depths greater than ~1 m. Collecting from recently eroded surfaces, quarries etc., limits a samples exposure to cosmic rays. Due to the location of the target elements, cosmogenic isotopes are also usually situated in the matrix of a sample rather than in fluid inclusions so using crushing techniques over step-heating will also minimise cosmogenic influences. 34

35 2.5.4 Component resolution As described above, there are a large number of potential components contributing to the noble gas signature of oceanic and subglacial basalts. Much of the literature assumes two-component mixing between an air-like component and a mantle component meaning corrections could be compromised by additional un-noticed components. A few studies take into account the possibility of three-component or even four-component mixing and the added complications of resolving these components [P Burnard et al., 2003; D. Harrison, 2003; D. Harrison et al., 1999]. A formal way of looking at multi-component mixtures is described in [Grenville Turner, 1971]. For an n-component mixture, each final noble gas ratio can be written in terms of its components: Equation 2.1: ( ) I R I n /R n l n any noble gas isotope reference isotope is the noble gas ratio of the nth component the fraction of R n present in the nth component and Equation 2.2: Equation 2.2 can be written for each noble gas isotope under study. As long as the reference isotope is kept the same, each noble gas ratio can be written in terms of n-1 other ratios from equation 2.2 and equation 2.3. For example, in a three-component mixture 130 Xe/ 36 Ar could be written as: 35

36 Equation 2.3: where a, b and c are constants whose values depend on the noble gas ratios of the three components. In contrast a four component mixture would need a further isotope to constrain the 130 Xe/ 36 Ar ratio: Equation 2.4: It is therefore necessary to know how many components are present before creating statistical fits to the data and extrapolating to mantle end members. Various formalized procedures for resolving multi-component mixtures exist [Williamson, 1968; York, 1969], but the most comprehensive is the Deming least-squares fit method outlined by [Moniot, 2009], which allows a fit to be calculated for any number of components and isotopes. This method provides a more complete fitting method than those that calculate an individual fit for each ratio of interest: As the entire dataset is treated simultaneously, fits for individual ratios will be consistent across the whole of the data Degassing For volatiles such as CO 2 that have a reasonably high abundance, analysis can be carried out on melt inclusions; these are inclusions formed before eruption [Saal et al., 2002]. This allows the pre-eruption melt composition to be directly determined. However, the low abundance of the noble gases makes it hard to measure their abundance in melt inclusions with any accuracy and there is the added complication of correctly identifying the melt inclusions (as opposed to later stage inclusions), which is 36

37 often not possible. More often, inclusions formed later in the eruption process must be relied upon. The eruption process can dramatically change the noble gas concentrations and elemental ratios found in these inclusions from the original melt composition. Not only are noble gas concentrations decreased but noble gas fractionation can result in elemental ratios which do not represent the source mantle. Studies of individual vesicles have revealed that even within a single sample, different degrees of fractionation due to degassing can be preserved [P Burnard, 1999]. Radiogenic ratios involving isotopes which have linked production paths (such as 4 He/ 21 Ne) have a known value which can provide a reference point to fix the degree of fractionation and, assuming only solubility-controlled degassing, noble gas fractionation can be modelled based on the relative solubility of the different noble gases in basaltic melts [P Burnard, 2001; Guillot and Sarda, 2006; J. Yamamoto and Burnard, 2005]. However, noble gas elemental ratios are often fractionated under disequilibrium conditions, when the relative diffusivity of the noble gases in the melt must also be considered [Aubaud et al., 2004; Gonnermann and Mukhopadhyay, 2007; Paonita and Martelli, 2006; Paonita and Martelli, 2007; Junji Yamamoto et al., 2009]. [Paonita and Martelli, 2007] use a model based on [Proussevitch and Sahagian, 1996; Proussevitch et al., 1993] and use the concept of disequilibrium degassing to explain variations in He-Ar-CO 2 systematics at mid-ocean ridges; Ar and CO 2 have similar diffusivity at low pressures meaning He/Ar and He/ CO 2 are similarly fractionated, but CO 2 diffuses more slowly than Ar at high pressures fractionating Ar/CO 2. This pressure dependence of CO 2 and Ar relative diffusivity, can explain an impressive array of observed He-Ar-CO 2 variations. Gonnermann and Mukhopadhyay (2007) developed a disequilibrium degassing model for the noble gases, starting from Henry s Law for a given noble gas and introducing a disequilibrium factor. They used this model to show that volatile rich, relatively undegassed mantle can produce surface samples with low helium concentrations via disequilibrium degassing; this is a possible explanation for why OIBs often show high 3 He/ 4 He ratios which are attributed to a volatile-rich mantle reservoir, but also show low He concentrations (the helium paradox -see Mantle models section). Both these models show that some degree of 37

38 disequilibrium degassing would be expected in most eruption scenarios, and consideration of relative noble gas diffusivity can produce very different fractionation patterns to models that only consider the relative solubilities of the noble gases. Disequilibrium degassing models depend on a number of parameters (such as degassing time, initial CO 2 and H 2 O content) that are not easy to constrain; further work is required to explore the implications of diffusivity degassing models for the noble gas abundances in oceanic and subglacial basalts Fractionation during melt generation Fractionation effects could also take place during melt generation due to the varying diffusivity of the noble gases in silicate minerals. For example, [P Burnard, 2004; P Burnard et al., 2004] suggest that the large difference in helium and argon diffusivity could generate primary melts with He/Ar ratios higher than the bulk mantle. As this fractionation is controlled by diffusivity, the degree of fractionation possible depends on the diffusion timescales involved, and fast diffusion pathways are needed to generate significant fractionation [P Burnard, 2004]. Although isotopic ratios such as 3 He/ 4 He could also be affected by such a process, plausible changes (a few percent) are not large enough to explain the large MORB/OIB 3 He/ 4 He differences. Fractionation due to differences in the incompatibility of the noble gases and parent isotopes could also have implications for helium isotope systematics: He, U and Th are usually assumed to be equally incompatible elements (preferentially partitioned into melts) but, if He is less incompatible than U and Th, then under certain conditions some mantle residues will retain a higher 3 He/(U+Th) and hence 3 He/ 4 He than primary melts. In this case high 3 He/ 4 He ratios in OIBs could represent a more depleted rather than less degassed mantle source [Albarede, 2008]. However it is hard to reconcile this model with evidence for mantle plumes from the deep mantle at ocean islands such as Iceland and Hawaii. 38

39 2.6 Models of the mantle The mantle stretches from the base of the crust at a depth of around 33 km to the core mantle boundary (CMB) at around 2900 km. It is generally subdivided by obvious discontinuities observed in seismic studies: The boundary at 410 km marks the end of the upper mantle and the start of the transition zone which stretches to 660 km and has been shown to represent the pressure-driven transformation of olivine to wadsleyite, followed by the spinel structure ringwoodite, which finally breaks down into perovskite and magnesiowüstite [Helffrich and Wood, 2001]. The lower mantle then reaches all the way to the outer core although there is an irregular layer observable near the CMB (the D layer) which varies in width. Although considerable progress has been made in understanding the mantle, its inaccessible nature means that much of its structure, chemistry and dynamics are still topics for debate. For example, a conclusive model for the formation and evolution of the mantle, particularly its relation to atmosphere formation has yet to be developed. Indeed, the relationship between mantle dynamics and plate tectonics, including issues such as the extent of recycling at subduction zones, still leaves many unanswered questions. In terms of mantle structure, the nature, origin and composition of the D layer is a key area of research. Building a model of the mantle which pulls together all of the different geochemical observations discussed above, together with the physical locations of samples showing different components and any geophysical evidence for their origins, is a large task which still requires much work. Such a model must also fit with geophysical evidence for the mantle s structure. Evidence from Rb/Sr, Sm/Nd systems, coupled with the high 3 He/ 4 He in OIB signatures, is often taken to represent a more primitive, less processed, mantle source; one that has retained more of its original composition from the Earth s formation [C. J. Allegre et al., 1983]. This led to the proposal of chemical layering within the mantle to explain how such a primitive reservoir had been preserved over the Earth s history. The upper mantle is depleted due to the extraction of continental crust and is well-mixed providing the reasonably homogenous, depleted geochemistry seen in MORBs. OIBs, on the other hand, sample a deep, primitive, 39

40 undegassed source mantle which has been convectively separate from the MORB source mantle since its formation. The seismic discontinuity at 670 km was considered a natural boundary for this isolated lower mantle layer [Allègre and Moreira, 2004]. The existence of this convectively isolated deep mantle layer also helps to solve other mantle puzzles. U and Th decay are the major source of heat and 4 He in the mantle. [O'Nions and Oxburgh, 1983] compared estimates of the amounts of He and heat produced at mid-ocean ridges and noticed a significant mismatch between the two quantities; considerably more heat is produced than can be explained by the amount of degassed helium. A convectively separate mantle layer allows the possibility that some of the missing helium is trapped in this layer while the produced heat escapes more efficiently. Other mismatch puzzles can be similarly solved by allowing for storage in the deep layer: About half the 40 Ar which should have been produced from 40 K over the Earth s history is not found in the atmosphere and could be stored in the deep mantle [C. J. Allegre et al., 1996]. Embellishments to this two-layer mantle model can also explain some of the heterogeneities involved in MORBs; for example the marble-cake model views the upper mantle as depleted but containing streaks of recycled oceanic crust which can provide a variety of different geochemical signatures [Claude J. Allegre and Turcotte, 1986; C. J. Allegre et al., 1980]. This layered mantle picture also fits neatly with the plume theory for OIB genesis: Theoretical models of the mantle show that at boundary layers, thermal instabilities can be generated and rise through the mantle to the surface where their high temperatures generate volcanism [Loper and Stacey, 1983; Mian and Chase, 1991]; seismic evidence is also compatible with a deep origin for some hotspots [Shen et al., 1998]. The difference in the noble gas composition of MORBs and OIBs is then explained as the difference between separate mantle reservoirs. The existence of plumes also allows for the formation of steady state theories of the mantle, where a small amount of mixing occurs between the deep mantle and the MORB sources as plumes journey through the upper mantle [Caffee, 1999; Kellogg and Wasserburg, 1990; O'Nions and Tolstikhin, 1994; Porcelli and Wasserburg, 1995]. The small primordial addition from lower mantle material to the upper mantle has been used to explain some 40

41 observations such as the high 3 He/ 4 He and 129 Xe/ 130 Xe ratios seen in upper mantlederived samples. Such steady state models require that both the MORB and OIB source mantle reservoirs contain primordial noble gases from the same origin. Evidence for distinct 3 He/ 22 Ne ratios between MORBs and OIBs and a recent detailed study of Xe isotopes which shows different Xe isotopic ratios for the MORB and OIB mantle sources have called into question whether this is possible [C. J. Ballentine, 2012; Mukhopadhyay, 2012]. The wealth of explanations generated by this two-layer mantle model has allowed it to become deeply ingrained in noble gas geochemical thinking. However this model is not without inconsistencies of its own. The helium paradox points out that the isolated deep mantle layer is supposed to be much less degassed than the depleted upper mantle, accounting for high 3 He/ 4 He ratios, but this is in contrast to the observed helium concentrations, which are much lower in OIBs than MORBs [Anderson, 1998b]. A few explanations are available for this problem: It has been shown that disequilibrium degassing of a more volatile-rich, undegassed source can produce samples with lower helium concentrations than depleted mantle samples [Gonnermann and Mukhopadhyay, 2007] which could explain the discrepancy whilst leaving the layered model intact. Alternatively, [J. Hopp and Trieloff, 2008] provide an explanation involving an earlier, pre-degassing stage of melt generation. There remains the issue of why other isotopic systems (e.g. lead [Anderson, 1998a]) fail to show the primitive compositions that would be expected if OIBs sample a primitive, undegassed mantle source. The preservation of this mantle layer was made an increasingly untenable model as overwhelming geophysical evidence for whole mantle convection came to light. As discussed above, the 670 km discontinuity represents a phase change of olivine and is not necessarily indicative of a distinct chemical layer. Tomographic evidence has shown subducted slabs penetrating through this discontinuity [R. D. Van Der Hilst et al., 1997]. This suggests that mantle convection penetrates beyond this point, as does evidence for displacement of the 670 km discontinuity at subduction zones and the presence of heterogeneities in the deep mantle which could represent recycled material [Helffrich and Wood, 2001]. Theoretical geophysical models of the mantle have shown 41

42 that the ability of the mantle to preserve a convectively isolated layer is limited and models which try to incorporate such a layer are incompatible with geochemical observations such as 3 He/ 4 He variations [Morgan, 1972; Naliboff and Kellogg, 2007; Peter E. van Keken and Ballentine, 1998; P. E. van Keken and Ballentine, 1999]. Mounting evidence that the enriched mantle components found in some OIBs represent recycled material again suggest the OIB reservoir was not isolated from mantle convection [Matthew G. Jackson et al., 2009; Kogiso et al., 1997; Kokfelt et al., 2006; Lassiter and Hauri, 1998]. In fact, geodynamical models of the mantle suggest recycled subducted crust is stored in the deep mantle and plays a role in the generation of the mantle plumes thought to be responsible for much ocean island volcanism [Christensen and Hofmann, 1994; Steinberger, 2000]. This weight of evidence caused a rethink in noble gas models of the mantle and various models arose which considered both whole mantle convection and the paradoxes solved by the two-layer mantle model (e.g.[c. J. Ballentine et al., 2002; Ellam and Stuart, 2004; Parman, 2007; Parman et al., 2005; Phipps Morgan, 1998; Mario Trieloff and Kunz, 2005; P. E. van Keken et al., 2001]. [Gonnermann and Mukhopadhyay, 2009] proposed a model in which the 660 km provides a partial boundary, allowing whole mantle convection but at low enough rates to preserve a unique lower mantle signature [Elliott, 2009]. However the rates of convection required to make this work are not compatible with the rates required by geophysical mantle models to explain observed heat flux and surface plate motions [Brandenburg et al., 2008]. Although the evidence for whole mantle convection is overwhelming, there is still much evidence to indicate the existence of at least a small primordial reservoir within the mantle: Nd isotopic evidence indicates the existence of an untouched primordial reservoir within the mantle [Boyet and Carlson, 2005; D Graham, 2010] and the expected primordial Nd signature of this reservoir has potentially been observed in samples from Baffin Island [M. G. Jackson et al., 2010]. The high 3 He/ 4 He ratios observed in some OIBs compared to MORBs are hard to explain without a distinct high 3 He/ 4 He reservoir within the mantle, as whole mantle convection should otherwise 42

43 homogenise this ratio throughout the mantle [C. J. Ballentine and Holland, 2008; Porcelli and Ballentine, 2002; Peter E. van Keken and Ballentine, 1998]. Some geophysical models of the early Earth have described possible mechanisms for the generation of a deep primordial reservoir, for example a sunken crystallised melt from early in the Earth s history [C-T A Lee et al., 2010]. In order to keep the positive explanations generated by the layered mantle model, alternative candidates to the lower mantle have been suggested for the isolated reservoir. Improvements in the understanding of seismic anomalies in the deep mantle still allow for potential reservoirs near the core-mantle boundary [Tackley, 2012; Trønnes, 2010; Rob D. van der Hilst and Karason, 1999]. Such a reservoir would provide an additional 3 He flux from somewhere deep within the mantle, producing the observed high 3 He/ 4 He ratios, as well as providing a potential source for other mantle signatures distinct from the depleted mantle. The D anomaly, could represent a distinct chemical layer with a more primitive geochemistry [I Tolstikhin and Hofmann, 2005; I N Tolstikhin et al., 2006]. The core [Porcelli and Halliday, 2001] is another possible source: Os isotopic systems have shown some evidence for a contribution from the outer core in samples studied by [Brandon et al., 2007]. The extent of mixing between these reservoirs and whether or not MORBs and OIBs have a common source for primordial volatiles is still debated [D W Graham, 2005; Darrell Harrison and Ballentine, 2005]. In addition, attempts have been made to explain the various geochemical observations which lead to the two-layer model without the need for an isolated mantle layer. [C. J. Ballentine et al., 2002] show that an upper mantle He concentration that removes the need for an isolated, volatile-rich lower mantle is a realistic figure (the zero paradox ). [Castro et al., 2005] and [Albarede, 2005] proposed that it may be possible to reconcile the heat/he paradox without a mantle boundary layer by considering different transport mechanisms for heat and He. [Bouhifd and Jephcoat, 2006] show that consideration of the effect of aluminium on argon solubility in silicate liquids could allow argon solubility to be higher than previously proposed at deep mantle pressures, providing a potential mechanism for the storage of primordial Ar in the Earth which does not require a separate mantle reservoir. Variations on the marble-cake mantle model have 43

44 attempted to explain helium isotopic variations in terms of shallow mantle heterogeneities without the need for a deep primordial reservoir [Albarede, 2008; Castro et al., 2009; Class and Goldstein, 2005], although the physicality of preserving such heterogeneities, particularly for an element as mobile as helium, is debated [Hart et al., 2008]. [Anderson, 1998a] rejects the plume theory for ocean-island volcanism and explains the various geochemical data in terms of shallow mantle heterogeneity; for example, high 3 He/ 4 He ratios come from low U environments stored in the shallow mantle and represent low 4 He, not high 3 He. Although shallow heterogeneity may explain some features of OIBs, this explanation does not fit with seismic evidence for mantle plumes whose origin must be a boundary layer in the deep mantle, nor with geochemical evidence such as helium isotope systematics [C. J. Ballentine et al., 2002; C. J. Ballentine et al., 2005b; Brandenburg and van Keken, 2007a]. 2.7 Volatile recycling A vital question impacting the constraints of mantle models is the nature of volatile recycling at subduction zones: are a significant amount of volatiles recycled back into the mantle and, if so, how they are distributed? Any significant amount of volatile recycling will need to be incorporated into the mantle models discussed above before their validity can be determined. In particular, recycling influences the mantle s water content which can impact mantle viscosity, conductivity and dynamics, melting processes and seismic studies [Hilton et al., 2002; Kelbert et al., 2009; Richard et al., 2002]. The presence of recycled material within both the convecting and the deep mantle is well established [Hauri and Hart, 1993; Stracke, 2005]. However it is often assumed that any volatiles carried down with this recycled material are returned to the surface through shallow cycling in arc-volcanism. This assumption stems mainly from a study by [Staudacher and Allegre, 1988], which looked at estimating the amount of volatiles erupted through subduction-related volcanism and compared the amounts to estimates of volatiles going down with the subducted slab. They concluded that ~98% of volatiles were recycled only in the vicinity of the subduction zone and never made it 44

45 into the mantle. This meant that any possible effects of volatile recycling were almost completely ignored. However, more recent studies of magmatic well gases have called this assumption into question: results show a heavy noble gas signature of mantle origin that matches the modern atmosphere, modified by the relative noble gas solubility in seawater [C. J. Ballentine and Holland, 2008; Holland and Ballentine, 2006]. Similar signatures in basalts have been interpreted as representing seawater incorporated during eruption but this mechanism is not plausible for magmatic CO 2. The explanation proposed by Holland and Ballentine is that volatile recycling at subduction zones does contribute a significant inventory to the mantle and in fact dominates the heavy noble gas mantle signature. They note that a subduction efficiency of only 5% would be needed to explain the observed seawater signature. [Matsumoto, 2006] questions whether it is reasonable to expect an unfractionated seawater signature to survive through the entire recycling process; however, Holland and Ballentine point out that the fact that the signature has survived unfractionated implies the noble gases have not decoupled from the host seawater. This would allow noble gas recycling to provide direct information on water recycling. [Manuel Moreira and Raquin, 2007] reviewed the evidence for the noble gas subduction barrier by considering the plausibility of a model that allows noble gas recycling: They conclude that a model which provides the atmospheric-like ratios seen in the mantle today would require xenon isotopic ratios during the Archean that are not seen in Archean samples and suggest noble gas fractionation before the Earth s accretion as an alternative mechanism for generating air-like signatures within the mantle. However, it is not clear that Archean samples provide a good test of the recycling hypothesis, nor how such a mechanism would leave the mantle s heavy noble gases with an atmospheric-like signature whilst resulting in very different light ratios. Nitrogen and noble gas data from Botswana diamonds has also been interpreted as showing a recycled volatile component [Ratan K. Mohapatra and Honda, 2006; R. K. Mohapatra et al., 2009] although alternative interpretations exist [C. Gautheron et al., 2005; Cécile Gautheron et al., 2006]. Other evidence for recycled volatiles in the mantle come from studies of coated diamonds: Noble gas analysis shows an air-like 45

46 composition for the heavy noble gases [Ozima and Zashu, 1991]. The mantle halogen composition provides a further test for a recycled volatile signature in the mantle. Some evidence for this can be seen in a study of submarine basalts along the Reykjanes Ridge in Iceland. The data are shown to be consistent with a model which assumes mixing involving an enriched Cl and Br source mantle and MORB source mantle [Unni and Schilling, 1978]. Studies of Cl in OIBs [Byers et al., 1985; John et al., 2010; Stroncik and Haase, 2004] and halogens in diamonds [R. Burgess et al., 2002] both show evidence for a recycled signature of mantle origin. A study of the behaviour of halogens in subducting serpentinite concluded that subduction can strongly affect the halogen budget of the mantle [John et al., 2011]. Lab studies of the properties of Br in diamond anvil cell experiments imply that although shallow Br recycling in arc magmas is extensive, a significant proportion of Br would be expected to be subducted with the down-going slab into the mantle [Bureau et al., 2010]. The mechanism by which these volatiles are carried into the deep mantle, avoiding the various dehydration stages during subduction, is still unclear although a study of the volatile content of a subducting slab have shown a preserved seawater noble gas component preserved at least as deep as 100 km [M. A. Kendrick et al., 2011; Sumino et al., 2010]. To confirm the presence of a seawater component in the mantle, models which describe how such subduction would progress are important. In the picture of whole mantle convection, modelling has shown that recycled material from downgoing slabs will end up concentrated in the deep mantle [Brandenburg and van Keken, 2007a]. As this matches the region where mantle plumes are thought to originate [Zhao, 2004], the OIB mantle source region should show the clearest evidence for a recycled signature. Evidence for lower 40 Ar/ 36 Ar ratios in OIBs is consistent with this theory [C. J. Ballentine and Holland, 2008; C. J. Ballentine et al., 2005b] but further heavy noble gas data from OIBs are needed to enhance this picture. 46

47 2.8 Iceland Iceland is an unusual ocean island in that it is situated on both on a spreading ridge and a hotspot; the two sources of volcanism combine to produce a highly heterogeneous geochemical signature. The mid-atlantic ridge runs as shown in Figure 5. Iceland s major neo-volcanic zones coincide with the onshore continuation of the mid-atlantic spreading ridge and lie along the Reykjanes peninsula and the eastern, western and northern volcanic zones (EVZ, WVZ and NVZ). Off-rift volcanism is also seen at the stratovolcano Oræfajökull, Snæfell, the Snæfellsnes peninsula and Vestmannaeyar at the S end of the EVZ. The WVZ and NVZ are dominated by tholeiitic and picritic basalts whereas the off-rift zones and EVZ tend towards transitional to alkali basalts [Peate et al., 2010]. The best seismic inversion of the hotspot anomaly beneath Iceland suggests a hot (~200 o C hotter than surrounding mantle), narrow (~100 km) plume which reaches a depth of at least 400 km [Wolfe et al., 1997]. This thermal anomaly leads to greater melt production towards central Iceland. Seismic studies have also shown the transition zone to be 20 km thinner under central and southern Iceland which suggests a deep mantle origin for the Iceland plume [Helmberger et al., 1998; Shen et al., 1998], although some seismic studies place the plume s origin no deeper than the transition zone [Foulger et al., 2000]. The current centre of the plume is placed under the Vatnajökull icecap. This location matches some of the highest 3 He/ 4 He ratios [Mark D. Kurz et al., 1985] although a low 3 He/ 4 He mantle source component has also been implicated for parts of central Iceland [Macpherson et al., 2005]. 47

48 Kolbeinsey Ridge plume centre NVZ WVZ EVZ Reykjanes Ridge Figure 5: Map of Iceland showing the major volcanic zones and the proposed centre of Iceland s mantle plume, adapted from [Sigmarsson and Steinthórsson, 2007]. The mid- Atlantic ridge continues as the Reykjanes Ridge in the SW and the Kolbeinsey Ridge in the N. The presence of both mid-ocean ridge and hotspot volcanism allows their interaction to be studied. Iceland s source mantle must be highly heterogeneous given the range of observed geochemical signatures [Hards et al., 1995; Hart et al., 1973; O'Nions et al., 1976; Sigmarsson and Steinthórsson, 2007; Sun et al., 1975; Thirlwall, 1995; Thirlwall et al., 2006]. [Thirlwall et al., 2004] find evidence for at least four mantle components. Much of Iceland s noble gas and halogen geochemistry has been explained by the mixing of a DMM component with more enriched or primordial plume components showing high 3 He/ 4 He and solar neon [P Burnard and Harrison, 2005; E T Dixon, 2003; E T Dixon et al., 2000; Mark D. Kurz et al., 1985; Unni and Schilling, 1978]. 48

49 Trace element, radiogenic element and oxygen studies have also inferred the presence of at least one recycled component in Iceland s source mantle: low 18 O found in basalts from central Iceland and the Reykjanes peninsula could originate from hydrothermally altered, recycled oceanic crust [Macpherson et al., 2005; Burnard and Harrison, 2005]. An increase in water content in samples going from the Reykjanes ridge towards central Iceland could indicate higher water content in the plume s mantle source due to a recycled component [Nichols et al., 2002]; Os and He isotopic systems in Iceland s neovolcanic zone suggest a complex mix of components in the Iceland plume including some primitive material along with ancient recycled crust ([Brandon et al., 2007; Debaille et al., 2009]. Evidence from Sr and Pb isotopic signatures for an enriched component (thought to have originated from recycled sediments) is seen at Oræfajökull [Prestvik et al., 2001]. Evidence for Cl and Br enrichments (Unni and Schilling, 1978) is compatible with a recycled seawater-derived signature in Iceland s volcanism. However the nature of the enriched recycled signature varies significantly across Iceland that has led many authors to conclude that Iceland s plume must itself be heterogeneous, with different recycled components sampled by different volcanic zones [Chauvel and Hemond, 2000; Hilton et al., 1990; Kokfelt et al., 2006; Peate et al., 2010]. For example, [Kokfelt et al., 2006] identify at least three plume components: a HIMU-like enriched component seen in the alkali basalts and associated with recycled material seen in many OIBs; a depleted component in Icelandic picrites that is nevertheless distinct from normal MORB in some trace element patterns and its low δ 18 O (this component could be attributed to recycled ocean-floor gabbros); and a high 87 Sr/ 86 Sr and 207 Pb/ 204 Pb component at Oræfajökull that is attributed to the involvement of recycled sediments. Although noble gas studies often associate Iceland s depleted signature with a MORB signature due to the proximity of the spreading ridge, the geochemical pattern of this component is not identical with N-MORB ( normal MORB) and it remains uncertain as to whether this component is part of the plume or of convecting mantle (DMM) origin. A coherent picture that marries this variety of enriched, anomalous and depleted signatures with the pattern of Iceland s noble gas compositions has yet to emerge. A key question for whole mantle convection models is the extent to which the different recycled components are associated with primordial 49

50 plume signatures such as higher 3 He/ 4 He ratios. This has implications for whole mantle convection models of the mantle that expect to find recycled slabs concentrated in the deep mantle [Hilton et al., 1990]. 2.9 Noble gas and halogen studies of Iceland to date The abundance of subglacially erupted samples in Iceland makes it an ideal place to study volatile compositions. Iceland s 3 He/ 4 He ratios have been extensively studied across much of the island [Condomines et al., 1983; E T Dixon, 2003; E T Dixon et al., 2000; Füri et al., 2010; Hilton et al., 1990; Hilton et al., 2000; Mark D. Kurz et al., 1985; Starkey et al., 2009]. The range of 3 He/ 4 He ratios from lower than MORB-like values along the Rekyjanes ridge to values over 30R A in central Iceland, is generally interpreted as representing varying degrees of influence from a high 3 He/ 4 He source mantle region associated with a mantle plume. A few ratios conflict with this picture; some of the highest 3 He/ 4 He ratios are seen in NW Iceland, well away from the proposed plume centre [Hilton et al., 1998; Hilton et al., 1999]. These ratios may be an indication that the mobility of helium leads to a more complex relationship between parent melts and the helium content of erupted samples [Breddam et al., 2000; Hofmann et al., 2011]. Ratios lower than MORB in the Oræfajökull-Snæfells volcanic system may indicate a radiogenic contribution to 4 He [Hilton et al., 1990]. Neon is also reasonably well studied across Iceland with a studies by [E T Dixon et al., 2000] and [E T Dixon, 2003] covering much of the Reykjanes peninsula, and an extensive dataset collected by [Füri et al., 2010] covering both basalt samples and geothermal fluids. The high 20 Ne/ 22 Ne ratios and low calculated 3 He/ 22 Ne ratios are attributed to a solar-like neon signature associated with Iceland s plume [D. Harrison et al., 1999; M. Moreira et al., 2001]. [Mario Trieloff et al., 2000] prefers an implanted solar origin to a direct nebula origin for this neon signature (an implanted signature could produce a 20 Ne/ 22 Ne of 12.5) but subsequent analyses of the same sample found consistently higher 20 Ne/ 22 Ne ratios ruling out a purely implanted origin [Mukhopadhyay, 2012]. The He/Ne mixing picture across Iceland is complex and 50

51 complicated by the large number of assumptions required to calculate a mantle ratio based on data from surface samples (for example, the nature of degassing-based fractionation). [E T Dixon, 2003] devises a model that allows MORB/OIB mixing to fit the He/Ne picture across the Reykjanes peninsula. This is extended to a wider data set by [Füri et al., 2010] but requires late-stage processes such as helium depletion and different 3 He/ 22 Ne ratios in the MORB and OIB source mantles to work. This has implications for steady-state models of the mantle that require a common source for MORB and OIB noble gases. Studies including the heavy noble gases in Iceland have tended to focus on a single sample area where particularly volatile rich samples are found. These samples are taken from picritic pillow mounds which are part of the Miðfell volcanoes in the Hengill system in SW Iceland. These picrites are thought to be the first products of a period of increased eruptive activity during early deglaciation caused by decreased lithostatic pressure [Tronnes, 1990]. Several noble gas studies [D. Harrison, 2003; Mukhopadhyay, 2012; Parai et al., 2009; Mario Trieloff et al., 2000] have been carried out on basalt samples from this area, often on the same sample set (samples named DICE 10 and DICE 11) originally introduced by [D. Harrison et al., 1999]. These samples show high 20 Ne/ 22 Ne (the original study by Harrison et al found a ratio of ± 0.32 in one run) and low (compared to MORB) 21 Ne/ 22 Ne. High 3 He/ 4 He ratios are found (~ 17R A ). 40 Ar/ 36 Ar ratios are some of the highest found in Iceland (up to ~7000) but low compared to MORB values, as is typical of many OIBs. Excesses in 129 Xe/ 130 Xe and 136 Xe/ 130 Xe are also resolvable in these samples although the highest values found are again much lower than typical excesses observed in MORBs. High quality data from DICE has been used to resolve mantle endmembers for this sample, and the resolved xenon isotopic ratios and He/Xe elemental ratio show a mantle endmember distinct from the MORB source mantle [Mukhopadhyay, 2012]. In particular the xenon isotopic ratios involving daughter isotopes of extinct 129 I and 244 Pu are distinct from the MORB source mantle, indicating that the OIB component of Iceland s source mantle cannot be the source of the convecting mantle s noble gases, contradicting steady state mantle models [C. J. Ballentine, 2012]. However, 51

52 Mukhopadhyay s resolved isotopic composition for this sample is based on the assumption of two component air-mantle mixing, whereas [D. Harrison et al., 1999] found evidence for at least three components in this sample. The effect of three (or more) component models on mantle endmember resolutions warrants further investigation, particularly as the quality of samples from this region means it has become a proxy for Iceland s entire source mantle. A wider sample base for Icelandic heavy noble gas data is also needed, as Iceland s mantle is known to be highly heterogeneous. The halogens across Iceland are less comprehensively studied than the noble gases. [Unni and Schilling, 1978] provide estimates of 17±2ppm and 0.06±0.01 ppm for the Cl and Br concentrations of Iceland s MORB component, and 61±6 ppm and 0.21±0.02 ppm for the Cl and Br concentrations of the more halogen-enriched component seen in the direction of Iceland s plume. Variations in halogen composition from the Rekyjanes ridge into SW Iceland show that OIBs potentially carry a different halogen signature to MORBs, with higher halogen concentrations associated with Iceland s OIB signature [Unni and Schilling, 1978], but the central Iceland halogen signature that shows the most plume influence has not been characterised. A more extensive data set for halogen concentrations across Iceland would provide a further constraint on models of Iceland s mantle, in particular the involvement and interaction of primordial and recycled signatures. 52

53 3 Chapter Three Experimental methods 3.1 Sample collection and field relationships Basalt samples were collected from Iceland s western and eastern volcanic zones (WVZ, EVZ), the Rekyjanes peninsula, the Snæfellsnes peninsula and the Oræfajökull- Snæfells volcanic system in SE Iceland (see Figure 6 and Table 3). Iceland s rift zones exhibit a number of subglacial volcanic features (tuyas see Figure 7 - and tindars) whose lowest sequence is often a thick pillow pile formed during the early stages of an eruption [Jones, 1968; Tuffen et al., 2010]. These features show clear evidence of iceconfinement in their shape and are often generated in a single eruption event. Tuyas are distinguished by a subaerial lava cap indicating the water depth at the point where an eruption broke through the overlying ice sheet (Figure 7 and Figure 8). Tindars are similar subglacial features to tuyas but do not display the distinctive flat top as they are formed in fissure eruptions that never broke through the overlying ice-sheet. The height of these features can therefore give a minimum value for the thickness of the overlying ice-sheet. The glassy rims of pillows from the base of these features are ideal for volatile analysis as the high subglacial eruption pressures and fast quenching nature of these samples minimises volatile loss. The subglacial origin of these samples also avoids the issue of seawater contamination often present in MOR glassy samples. 53

54 Figure 6: Map of Iceland showing locations of samples collected in this study (solid diamonds) and those analysed in this study donated by Alex Nichols (solid squares) [Nichols et al., 2002]. Map adapted from [Sigmarsson and Steinthórsson, 2007]. 54

55 Table 3: Sample locations and descriptions (nr = not recorded) Sample Name Location Coordinates (dd0mm.mmm') N W Elevation (m) UND1 Undirhlíđar Quarry ' ' nr UND2 Undirhlíđar Quarry ' ' nr UND3 Undirhlíđar Quarry ' ' nr Notes Pillows at SE end of quarry UND4 Undirhlíđar Quarry ' ' nr Pillows at NE end of quarry UND5 Undirhlíđar Quarry ' ' nr Sample from lava cave UND6 Undirhlíđar Quarry ' ' nr Pillows at NW end of quarry HLO1 Hlöđufell, N-side ' ' nr HLO2 Hlöđufell, N-side ' ' nr HLO3 Hlöđufell, N-side ' ' nr HLO4 Hlöđufell, N-side ' ' nr Ice-confined pillows near base of tuya BREK1 Brekknafajöll Jarlhettar ' ' nr Pillows at base of tindar - BREK2 Brekknafajöll Jarlhettar ' ' 392 samples show extensive alteration LAUN1 Launöldur ' ' 584 LAUN2 Launöldur ' ' 584 LAUN3 Launöldur ' ' 584 LAUN4 Launöldur ' ' 584 LAUN5 Launöldur ' ' 584 LAUN6 Launöldur ' ' 584 Extensive pillows and pillow breccia on lake shore 55

56 BRIDGE Road nr. Hrauneyjar ' ' nr Pillows. Approx. location only MID1 Dagmálafell Quarry ' ' nr Whole pillow MID2 Dagmálafell Quarry ' ' nr MID3 Dagmálafell Quarry ' ' nr Sample from under overhanging MID4 Dagmálafell Quarry ' ' nr outcrop MID5 Dagmálafell Quarry ' ' nr Pillow + gabbro xenolith HŪS1 Quarry nr. Húsafell ' ' nr HŪS2 Quarry nr. Húsafell ' ' nr HŪS3 Quarry nr. Húsafell ' ' nr LOM1 Lómagnúpur ' ' nr Pillows at base of outcrop - in situ? LOM2 Lómagnúpur ' ' nr Pillows under overhang - very little glass left LOM3 Lómagnúpur ' ' 214 Phenocryst rich sample KEL1 Keldunúpur ' ' 55 Pillows KEL2 Keldunúpur ' ' 55 Pillows KEL3 Keldunúpur ' ' 81 From chilled margin KEL4 Keldunúpur ' ' 81 From base lava KEL5 Keldunúpur ' ' 81 Not in situ KEL6 Keldunúpur ' ' 84 Pillows KEL7 Keldunúpur ' ' 140 Higher pillow sequence KEL8 Keldunúpur ' ' 140 Higher pillow sequence Gođ1 Gođafell, tuff cone ' ' 576 Gođ2 Gođafell, tuff cone ' ' 726 Gođ3 Gođafell, tuff cone ' ' 726 FÖG1 Fögruntungubrýr ' ' 92 Hyaloclastite margin 56

57 FÖG2 Fögruntungubrýr ' ' 115 FÖG3 Fögruntungubrýr ' ' 140 FÖG4 Fögruntungubrýr ' ' 140 Pillows FÖG5 Fögruntungubrýr ' ' 190 Dyke margin KRIS1 nr. Kristinatindar ' ' 734 KRIS2 nr. Kristinatindar ' ' 344 MOS1 Morsadalur ' ' 107 MOS2 Morsadalur ' ' 192 Gođ4 Gođafell ' ' 217 Gođ5 Gođafell ' ' 217 VAL1 Slaga ' ' 274 Glassy fragments VAL2 Slaga ' ' 274 Whole rock VAL3 Slaga ' ' 274 Glassy intrusion VAL4 Slaga ' ' 274 Clast from overlying diamict VAL5 Slaga ' ' 232 Base lavas VAL6 Slaga ' ' 245 Peperitic - some pillows VAL7 Slaga ' ' 254 VAL8 Slaga ' ' 340 Tracydacite flow VAL9 Slaga ' ' 342 Pillow breccia VAL10 Slaga ' ' 342 Hyaloclastite from pillow breccia VAL11 Slaga ' ' 342 Same unit as 9&10 - more coherent lava VAL12 Slaga ' ' 393 Columnar lobe at top of sequence HOF1 nr. Litli Höff ' ' 297 Late stage pillows HOF2 nr. Litli Höff ' ' 443? Clast from hyaloclastite 57

58 HOF3 nr. Litli Höff ' ' 475 nr. Entablature HOF4 nr. Litli Höff ' ' 453 Pillow breccia HOF5 nr. Litli Höff ' ' 225 Lava HOF6 nr. Litli Höff ' ' 225 Very phenocryst rich lavas in HOF7 nr. Litli Höff ' ' 225 gully beneath tuff cone HOF8 nr. Litli Höff ' ' 225 HOF9 nr. Litli Höff ' ' 225 Includes xenolith HOF10 nr. Litli Höff ' ' 225 Not in situ - from scree Gođ6 Gođafell ' ' 417 Basalt below chilled rhyolite margin HLO100 Hlöđufell, gully on S side ' ' 536 Base pillow mound HLO101 Hlöđufell, gully on S side ' ' 533 Base pillow mound HLO102 Hlöđufell, gully on S side ' ' 548? Pillows in hyaloclastite - GPS dodgy on elevation HLO103 Hlöđufell, gully on S side ' ' 600? Higher up base pillow mound HLO104 Hlöđufell, gully on S side ' ' 637 Pillows near transition to hyaloclastite HLO105 Hlöđufell, gully on S side ' ' 637 Pillows near transition to hyaloclastite HLO106 Hlöđufell, gully on S side ' ' 637 Pillows near transition to hyaloclastite HLO107 Hlöđufell, gully on S side ' ' 637 Hyaloclastite HLO108 Hlöđufell, above gully ' ' 769 From lava-fed delta HLO109 Hlöđufell, above gully ' ' 847 HLO110 Hlöđufell, above gully ' ' 829 Hyaloclastite JARL1 Jarlhettur - WVZ tindar ' ' 400 Pillow glass JARL2 Jarlhettur - WVZ tindar ' ' 380 Pillow glass 58

59 JARL3 Jarlhettur - WVZ tindar ' ' 373 Pillow glass JARL4 Jarlhettur - WVZ tindar ' ' 380 Pillow glass JARL5 Jarlhettur - WVZ tindar ' ' 444 Pillow glass JARL6 Jarlhettur - WVZ tindar ' ' 385 Pillow glass JARL7 Jarlhettur - WVZ tindar ' ' 384 Pillow glass JARL8 Jarlhettur - WVZ tindar ' ' 384 Pillow glass JARL9 Jarlhettur - WVZ tindar ' ' 458 Pillow glass KALF1 Kalfstindar - WVZ tindar ' ' 221 Pillow glass KALF2 Kalfstindar - WVZ tindar ' ' 221 Pillow glass KALF3 Kalfstindar - WVZ tindar ' ' 230 Pillow glass KALF4 Kalfstindar - WVZ tindar ' ' 230 Pillow glass KALF5 Kalfstindar - WVZ tindar ' ' 219 Pillow glass KALF6 Kalfstindar - WVZ tindar ' ' 279 Pillow glass KALF7 Kalfstindar - WVZ tindar ' ' 283 Pillow glass KALF8 Kalfstindar - WVZ tindar ' ' 222 Pillow glass KALF9 Kalfstindar - WVZ tindar ' ' 373 Pillow glass EST1 Efstadasfjall - tuya ' ' 273 Glass pieces in hyaloclastite EST2 Efstadasfjall - tuya ' ' 244 Pillow glass KAR1 Karanjukur ' ' 619 Pillow pieces in pillow breccia KAR2 Karanjukur ' ' 639 Pillow pieces in pillow breccia KAR3 Karanjukur ' ' 636 Hyaloclastite KAR4 Karanjukur ' ' 639 Pillow pieces in pillow breccia KAR5 Karanjukur ' ' 639 Pillow pieces in pillow breccia - not in situ ERIK1 Eiriksjokull - ice-capped tuya ' ' 410 Hyaloclastite 59

60 ERIK2 Eiriksjokull - ice-capped tuya ' ' 466 Broken pillow pieces in hyaloclastite ERIK3 Eiriksjokull - ice-capped tuya ' ' 548 Broken pillow pieces in hyaloclastite ERIK4 Eiriksjokull - ice-capped tuya ' ' 549 Broken pillow pieces in hyaloclastite ERIK5 Eiriksjokull - ice-capped tuya ' ' 549 Broken pillow pieces in hyaloclastite ERIK6 Eiriksjokull - ice-capped tuya ' ' 549 Broken pillow pieces in hyaloclastite ERIK7 Eiriksjokull - ice-capped tuya ' ' 549 hyaloclastite ERIK8 Eiriksjokull - ice-capped tuya ' ' 550 pillow glass ERIK9 Eiriksjokull - ice-capped tuya ' ' 605 hyaloclastite ERIK10 Eiriksjokull - ice-capped tuya ' ' 605 not in situ - from scree SN1 Tuffacious Sediments ' ' 258 pyroxene clast SN2 Tuffacious Sediments ' ' 258 hyaloclastite matrix SN3 Tuffacious Sediments ' ' 258 hyaloclastite matrix SN4 Tuffacious Sediments ' ' 258 pyroxene SN5 Tuffacious Sediments ' ' 258 olivine SN6 Tuffacious Sediments ' ' 258 diorite SN7 Tuffacious Sediments ' ' 258 olivine-rich clast SN8 Tuffacious Sediments ' ' 258 whole-rock sample SN9 Tuffacious Sediments ' ' 258 SN10 Tuffacious Sediments ' ' 258 SN11 Tuffacious Sediments ' ' 258 whole-rock sample SN12 Tuffacious Sediments ' ' 258 whole rock plus phenocrysts 60

61 Figure 7: Schematic sketch of the formation of a tuya from [Jones, 1968] 61

62 Figure 8: Hloðufell: a tuya in Iceland s WVZ (686 m from base to top). The confinement imposed by the ice-cap overlying this eruption gives the tuya its distinctive shape. The subaerial lava cap at the top of the tuya indicates the depth of the surrounding meltwater at the point where the eruption broke through the overlying ice. Sample collection in the WVZ and EVZ focused on the pillow mounds at the base of tuyas and tindars. Hyaloclastite samples were also collected although these were found to be too water-rich for noble gas analyses. Eiríksjökull, Hloðufell and Efstadasfjall are classic Icelandic tuyas in the WVZ [Skilling, 2009]. Jarlettur and Kalfstindar are Icelandic tindars. Samples from the Rekyjanes peninsula are also from the glassy rinds from pillow basalts, mainly from quarries showing freshly exposed surfaces, thus avoiding additions from cosmogenic noble gas isotopes (e.g. 3 He and 21 Ne). The MID (Miðfell) sample set are pillows from picritic basalts containing large olivine phenocrysts and gabbroic nodules (see Figure 9) and are from the same sample area as the DICE10 sample that has shown unusually high mantle noble gas concentrations and that have been taken as representative for the noble gas composition 62

63 of the Icelandic mantle [D. Harrison et al., 1999; Mukhopadhyay, 2012; Mario Trieloff et al., 2000]. Figure 9: Sample MID5 showing a gabbroic xenolith. These are ubiquitous in the DICE sample area in SW Iceland. Glassy samples from pillows or chilled margins were collected from a series of sheet sequences found in SE Iceland; a thick-ice origin has been proposed for this sequence making the retention of significant mantle volatiles a possibility [Smellie, 2008]. Glassy and phenocryst-rich samples were also collected from Oræfajökull, the largest stratovolcano in Iceland located at the S edge of the Vatnajökull ice sheet. Previous Nd and Pb isotope studies of basalts from Oræfajökull volcano revealed an enriched geochemical signature attributed to an EM2-like recycled mantle component [Prestvik et al., 2001]. Samples from the Snæfellsnes peninsula in western Iceland are from a single outcrop of tuffaceous sediments which contained an extraordinary variety of xenoliths (Figure 10). A set of pillow glass samples from the NVZ, EVZ and central Iceland were also analysed. These were provided by Alex Nichols and have been 63

64 studied previously for their volatile contents [Nichols et al., 2002]. Figure 10: A large variety of xenoliths were found in this tuffaceous sediment, Snæfellsnes Peninsula (samples SN4, SN6 and SN12 in Table 11). As far as possible, samples showing visible alteration were avoided and fresh glass was collected. However, alteration is ubiquitous across Iceland and it is likely that many samples have been affected to some extent. To minimise the presence of cosmogenically produced isotopes, sample collection focused on obtaining fresh samples that had not obviously been exposed for long periods of time (e.g. quarries). 64

65 3.2 Sample preparation Samples were partially crushed and sieved to separate chips mm in size. Smaller pieces were avoided to minimise the risk of losing larger vesicles containing mantle volatiles. For fragments above 2 mm, a large number (over 3000) crushing strokes in UHV were needed to powder samples (compared to between 5 and 1000 strokes for smaller samples) increasing the potential for air contamination released during crushing from the action of the crusher and sample fragments against the crusher walls. The amount of mantle noble gas released during each crush is variable as the vesicle sampling by this crushing method is random, and so keeping the number of crushes, (and hence the amount of crusher-induced air contamination) low is advantageous in picking up distinct mantle noble gas signatures. Fresh, unaltered glasses were hand-picked using a binocular microscope. For non-glass samples (mostly olivine), fragments showing no signs of alteration or adhering rock matrix were similarly hand-picked. Hand-picked fragments were cleaned in an ultrasonic bath in 2% HNO 3 to remove contaminants adhering to the fragments surfaces, followed by ethanol and de-ionised water for 10 minutes. Samples were then allowed to dry completely under a heat lamp before loading into the crushing system. The loaded samples were baked at 150 (±10) o C for 12 hours to release as much air as possible adhering to the crusher walls, as well as any absorbed air contamination in the samples. Four loaded crushers were left to pump down for a further 36 hours or longer until a vacuum pressure close to 1x10-9 torr was achieved. Analyses of noble gases are carried out under ultra-high vacuum (UHV) conditions with a low extraction line blank, but glassy basalts frequently show a significant atmospheric component that can mask mantle signatures of interest for all the noble gases except He. Some of this will be due to release from the equipment walls of the UHV extraction line which the baking and prolonged pumping steps described above are designed to minimise. However, a component which resides in the sample itself has proved hard to reduce. The exact mechanism for the addition of this component is not 65

66 clear but as the majority of this component is believed to be a late-stage addition, possibly sited in self-healing fractures created during sample preparation [Chris J. Ballentine and Barfod, 2000]. In an attempt to reduce this component, one particular glass sample (MID1) was prepared in a glove box filled with a nitrogen atmosphere at a pressure slightly above atmospheric pressure. The sample was crushed, picked, cleaned and loaded inside the glove-box in an attempt to minimise the contamination by this atmospheric component. A matching sample was prepared in the same way but under normal atmospheric conditions for comparison. In addition, a small sample of the glove-box atmosphere was analysed to check the noble gas content, which showed significantly lower concentrations than normal laboratory air. The results from the analyses of the two samples are shown in Table 4. 66

67 Table 4: In an attempt to minimise the addition of atmospheric contamination added during sample preparation, a sample was prepared in a glove-box filled with a nitrogen atmosphere. Compared with a sample prepared in laboratory air, no significant improvement was achieved: 40 Ar/ 36 Ar ratios measured in the glove-box prepared sample showed similar influence from air contamination and 36 Ar concentrations (a good indicator for the atmospheric contamination) were actually higher than for the sample prepared in laboratory air. 36 Ar x Sample Crushes (cm 3 STPg -1 ) 40 Ar/ 36 Ar MID1_7 Prepared in nitrogen atmosphere (glovebox) MID1_7 Prepared in laboratory air ± ± ± ± ± ± 5 TOTAL 6913 ± ± ± ± ± ± ± ± 5 TOTAL 2323 ± ± Noble gas mass spectrometry Equipment overview Noble gas isotopic analyses were carried out using an UHV extraction line, a VG5400 noble gas spectrometer running the ISOWORKs software (GV software) and an MAP noble gas spectrometer running a custom-written FORTRAN program (see Figure 11). The system is kept at UVH (better than 1 x 10-9 torr) by ionization pumps; two situated at either end of the VG5400 flight tube, one connected to the MAP flight tube and one 67

68 connected to the extraction line. A turbomolecular pump and two-stage rotary pump are also connected to the extraction line to pump the system down from higher pressures after sample loading or system repairs. Figure 11: VG5400 Spectrometer and gas extraction system (foreground) Mass spectrometer details The VG5400 noble gas spectrometer uses two detectors, a Faraday cup and a Burle Channeltron electron multiplier set at 90 o to the Faraday detector. The electron multiplier works in conjunction with the electrode of a Daly detector, used to deflect the ion beam into the electron multiplier. The VG5400 uses a Nier-type electron impact source consisting of a 1 mm diameter, four-turn tungsten filament. The VG5400 is 68

69 designed to allow the electromagnetic field to be optimized for either the light or heavy noble gases using shims which can alter the alignment of the electromagnet. Since this study measured both the light and heavy noble gases, the magnet shims could not be employed. The electromagnetic field was therefore optimised for maximum sensitivity at 40 Ar as a compromise. Tuning of source and detector parameters was carried out using the GV ISOWORKs software except for the Daly electrode voltage, which must be tuned manually. These parameters were optimised for each individual noble gas based on achieving the best sensitivity and peak shape for an aliquot from the calibration bottle (see Extraction line section). The noble gas quantities present in a calibration shot are larger than the majority of sample analyses: [P Burnard and Farley, 2000] have shown that optimum tuning parameters such as the half-plate voltage can be sensitive to spectrometer pressure and tuning at a maximum pressure may minimise inaccuracies introduced by setting optimum tuning parameters at a different pressure to sample gas pressure. The spectrometer s resolution is ~600, allowing common interferences such as HD on 3 He and hydrocarbons in the heavy noble gas mass regions ++ to be resolved. The exception is Ne, which has irresolvable interferences from CO 2 and 40 Ar at mass 20 and mass 22 respectively (see CO 2 and 40 Ar ++ corrections section). An SAES NP10 getter situated at the spectrometer source reduces the levels of active gases to a minimum and the blank spectrum of the spectrometer was monitored routinely for any significant changes in interference peaks. Data collection and preliminary data analysis are carried out by the ISOWORKs software. The MAP spectrometer follows a similar set-up to the VG5400, also using a SAES NP10 getter situated at the spectrometer source, and is connected to the same extraction line. It also uses a Faraday cup and multiplier but does not include a Daly detector. Tuning is carried out manually, but since the MAP was only used to analyse Ar isotopes (due to having a lower sensitivity than the VG5400), re-tuning for different noble gases was not necessary. Data collection and preliminary data analysis is carried out using a custom-designed FORTRAN program SPEC. 69

70 3.3.3 Noble gas extraction line The extraction line is connected to both spectrometers and has been dedicated to noble gas analysis and so is relatively clean of contaminants. The majority of the extraction line consists of a series of Nupro 0.5" bellows-sealed valves connecting short stainless steel pipe sections. Several activated charcoal traps connected to heating coils allow for separation of the heavy noble gases when necessary. In addition, a charcoal trap close to the spectrometer source reduces Ar and CO 2 interferences during analysis of He and Ne isotopes. A cryotrap cooled by compressed helium reaching temperatures of 45K, enables separation of He and Ne at 55K. Two Zr-Al alloy getters (a SAES NP10 getter run at room temperature and a SAES GP50 getter run at ~230 o C) are used to remove active gases such as CO 2 and H 2 O from the noble gases released by crushing. The extraction line also includes a calibration bottle having an elemental and isotopic composition shown in Table 5. This bottle contains laboratory air spiked with 3 He and 4 He to produce a composition which is close to many mantle samples. As there was some doubt as to the exact composition of this bottle, the 3 He/ 4 He ratios and concentrations were calibrated to the atmospheric 3 He/ 4 He ratio of x 10-6 using a 2 cm 3 shot of laboratory air. Ne, Ar, Kr and Xe concentrations and isotopic ratios were calibrated using a second bottle (of known volume 1850 cm 3 ) containing laboratory air filled at ± torr. 70

71 Table 5: Composition of the first 1 cm 3 shot taken from the helium-spiked calibration bottle. Isotope Concentration in 1 cm 3 shot (cm 3 STPg -1 ) 3 He 4 He 3 He/ 4 He (R/R A ) 20 Ne 21 Ne 22 Ne 36 Ar 38 Ar 40 Ar 80 Kr 82 Kr 83 Kr 84 Kr 86 Kr 129 Xe 130 Xe 131 Xe 132 Xe 134 Xe 136 Xe (3.57 ± 0.13) x (1.746 ± 0.006) x ± 0.5 (6.38 ± 0.02) x (1.89 ± 0.01) x (6.50 ± 0.02) x (1.208 ± 0.003) x (2.276 ± 0.007) x (3.606 ± 0.008) x (9.97 ± 0.09) x (5.09 ± 0.05) x (5.07 ± 0.05) x (2.52 ± 0.02) x (7.69 ± 0.07) x (8.92 ± 0.10) x (1.37 ± 0.02) x (7.16 ± 0.08) x (9.07 ± 0.11) x (3.52 ± 0.04) x (2.99 ± 0.04) x

72 3.4 Sample Analysis Sample crushers Samples were loaded into cylindrical steel crushers based on a design especially for this purpose by [Sumino et al., 2001]. Samples were crushed under UHV using an alternating electromagnetic field produced by a solenoid to move a nickel piston inside the vacuum system. Samples were crushed in stages ranging from 5 to 1000's of crushes. The exact number required depends on the sample and small quantities of each sample were run before the main analysis to ascertain the best crushing pattern. The first crushing step was kept to a minimum number of strokes as experience showed that this step often released the largest atmospheric noble gas component Noble gas purification and separation procedure The released gas was expanded into a SAES GP50 getter run at ~230 o C for ten minutes to minimise the active gas content. The heavy noble gases (Ar, Kr and Xe) were allowed to condense on a liquid-nitrogen cooled activated charcoal finger for five minutes. Neon and some helium were then trapped using the cryotrap run at 45K for a further five minutes. Any trapped helium was then released by heating the cryotrap to 55K. The separation temperature was calibrated using an aliquot from the calibration bottle and gradually increasing the cryotrap temperature to determine the optimum separation temperature. After a further five minutes cleaning on an NP10 getter at room temperature, helium was expanded into the GV5400 mass spectrometer for analysis. Neon content was monitored during the helium (and vice versa) run to assess the efficiency of He-Ne gas separation with ~0.01% of the total Ne seen in the He run and vice versa. A further liquid-nitrogen cooled charcoal finger near the source removed most of the remaining heavy noble gases. 3 He was measured using the channeltron, whereas the more abundant 4 He was measured using the Faraday cup collector. 72

73 After the helium run, the remaining, neon was released by heating to 105K. After five minutes cleaning, 20 Ne, 21 Ne and 22 Ne were analysed. 44 CO 2 and 40 Ar were also analysed to allow interferences of 44 CO ++ 2 at 22 Ne and 40 Ar ++ at 20 Ne to be corrected for. The correction procedure is described below. Further potential interferences on the 20 Ne peak such as H 18 2 O + and HF + were monitored by regular checks between runs of the blank peak shape. In addition, H 16 2 O + was recorded during each run, with levels indicating a contribution from H 18 2 O + at mass 20 of less than 0.1% of the total count for a typical sample. All gases were recorded using the electron multiplier detector. The heavy noble gases were released by heating the cold finger to ~500K. Ar, Kr and Xe were then either analysed immediately or the gases were further separated. Separation of Ar from Kr and Xe was achieved by trapping the Kr and Xe on a cold finger typically held at -163K (as for He and Ne, separation temperature was determined using a shot from the calibration bottle) whereas the Ar was trapped onto a separate cold finger held at 77K. Separation efficiency was monitored by recording 84 Kr in the Ar run and 36 Ar in the Kr/Xe run with typically 10% of the total 84 Kr observed in the Ar run. Separation of Ar from Kr/Xe allowed Ar to be analysed on the MAP spectrometer simultaneously with Kr/Xe analysis on the VG5400, which considerably reduced sample run times as well as improving the accuracy of regression lines (due to shorter run times). However, instabilities in the MAP s software resulted in the loss of some Ar data due to software crashes and so, for the majority of samples, Ar, Kr and Xe were analysed together on the VG Ar and 36 Ar were recorded using the Faraday detector; 38 Ar was recorded on both collectors and all other isotopes were recorded on the electron multiplier detector Isotopic analysis On both spectrometers, the software allows method or sequence files for each analysis to be written in advance. These files determine which masses, baselines, detector, settling times and number of readings/counting times are to be used. When the 73

74 appropriate gas fraction was released into the spectrometer a clock was started on the software and the pre-written method files began. The correct tuning settings must be set before the start of each run and the magnet was cycled through the masses to be analysed (a pre-run ) to minimise any hysteresis effects. If enabled, peak-centring software first found the exact peak centre for each mass of interest, meaning small dayto-day movements in mass locations were taken into account. Peak centring was not enabled for the Ne analyses due to the need to measure several different parts of the 20 Ne peak to correct for interferences. The software then recorded background noise levels on the detectors at set masses close to those that were about to be analysed. Levels above zero are corrected for automatically by the software. The software cycles the magnet through the masses to be analysed (from low to high) for 10 cycles. After entering the spectrometer, gas levels will usually drop as gases are absorbed by the detectors (and source) so this allows a best-fit regression to determine the initial gas concentrations at time zero (an example is shown in Figure 12). Generally for each cycle, each peak was measured for 30s, but the high abundance of 40 Ar and 4 He meant 20s per cycle was sufficient. For samples with very low 3 He concentrations, 3 He peaks were analysed for 60s for each cycle. The ISOWORKs software only uses a basic linear regression program to calculate the gas concentration at time zero, which takes no account of error-weighting (each point is the average measurement over the 30 s or 60 s measurement on each peak with readings diverging more than 2σ discarded; each average has an associated standard deviation) or potential non-linear trends. For this reason regressions were carried out using the raw data and the MatLab robust regression tool ( robustfit ). 74

75 Robust regression line: y = [-2.68(±0.06)] x 10 4 *x + [3.19(±0.02)] x 10 8 Figure 12: Example regression for a 10-cycle analysis of 40 Ar. The concentration at t = 0 is given by the intercept of a robust regression through the data Extraction line blanks and calibrations Blanks and calibration analyses followed the same procedure as described above, including expansion into the sample crusher volume before analysis. There was no measurable 3 He blank: Other typical blank levels were 1.8 x cm 3 STP 20 Ne, 2.2 x 10-9 cm 3 STP 40 Ar, 1.7 x cm 3 STP 84 Kr and 1.8 x cm 3 STP 132 Xe. Calibration runs allowed sample noble gas concentrations and any mass discrimination effects to be corrected for and were run on a regular basis. An empty crusher was also used during some blank analyses to check that a significant amount of gas was not released from crusher walls during the crushing procedure. The first few crushes after 75

76 sample loading produced levels of gas up to ten times higher than a static extraction line blank, presumably as vibrations and impact from the piston caused additional gas release from the extraction line walls. This is considered a worst case blank as during a real sample run the impact of the crusher on the crusher walls would be cushioned by the powdered sample. However, after this initial crush blank, gas levels released when using the empty crusher were not significantly higher than static extraction line blanks, although a large number of crushes (over ~ 1000) could produce blank levels up to three times higher than static extraction line blanks. To minimise inaccuracies due to active crusher blanks, the first crush of each sample was kept to a minimum number of strokes and an extraction line blank using the same number of strokes in the empty crusher was run before sample analyses. This was used to correct sample data. Static blank levels remained reasonably consistent and typical levels released during this study were around 1% or less of sample releases, although this figure rose to 10% for a few low release sample crushes. The composition of the extraction line s calibration bottle is given in Table CO 2 and 40 Ar interference corrections 40 Ar ++ and 20 Ne + were partially resolved, allowing both 40 Ar ++ and ( 40 Ar Ne + ) to be measured. In the absence of other interference, a value for 20 Ne + can be recorded on the peak shoulder. However a sharp enough peak shoulder could not be obtained and so the 20 Ne + value was obtained from ( 20 Ne + (cps) + 40 Ar ++ (cps)) - 40 Ar ++ (cps) (cps = counts per second). 44 CO ++ 2 and 22 Ne + were not resolved so a correction factor was needed to get a result for 22 Ne + alone. Correcting for this interference was not straightforward as CO /CO 2 was found to depend on the concentrations of 20 Ne + and 4 He +. Although He and Ne were separated during gas extraction, even a high He/Ne separation efficiency (<0.01% of 4 He observed in the neon run) can leave a potentially significant amount of 4 He present in the Ne run. He and Ne concentrations were variable and unpredictable for each analysis, leading to different CO /CO 2 values, even between different crushes 76

77 of the same sample. Other background concentrations (CO 2 and H 2 O) were monitored and CO /CO 2 showed no significant variation within the observed range. Correction problems are compounded by the fairly large CO 2 background in the VG5400, meaning ++ the CO 2 signal could be as much as 50% of the total signal at mass 22 during the smallest sample runs. Calibration of CO /CO 2 was achieved by running number of shots from both the helium spiked calibration bottle and a calibration bottle containing laboratory air. Different volumes of the extraction line were used for each run, giving a spectrum of Ne and He concentrations covering the range observed in samples. Assuming minimal discrimination between masses 20 and 22, this data allowed the CO ++ 2 /CO + 2 ratio for each run to be calculated assuming 20 Ne/ 22 Ne = 9.8. The CO ++ 2 /CO + 2 depended linearly on 20 Ne + and 4 He + concentrations and a best fit equation using MatLab s robust regression tool was determined after multiple calibration analyses at varying 20 Ne + and 4 He + partial pressures (see Figure 13). A further run of calibration tests then checked the reliability of this correction. This calibration procedure was done after each source retune and also if significant changes in background CO 2 and H 2 O were observed. Using corrections based on this calibration procedure, 77% of 20 Ne/ 22 Ne ratios measured in calibration runs between 29/01/2010 and 01/04/2011 were within 1σ of the atmospheric ratio of 9.80 ±0.08 (Figure 14). 77

78 Robust regression line: y = [-9.05(±1.25)] x 10-8 *x + [0.0060(±0.0002)] Robust regression line: y = [-8.82(±0.51)] x 10-9 *x + [ (±0.0001)] Figure 13: The correction factor for CO2 ++ at mass 22 showed a linear dependence with both 20 Ne + and 4 He +. Calibrations run over a variety of 4 He + pressures at constant 20 Ne + and a variety of 4 He + pressures at constant 20 Ne + allowed this pressure dependence to be quantified and corrections at mass 22 to be made for the range of 4 He + and 20 Ne + observed in this study s samples. 78

79 AIR Figure 14: Corrected 20 Ne/ 22 Ne ratios from 38 calibration shots run between 29/01/2010 and 01/04/2011. Error bars are 1σ. The 22 Ne correction was re-done several times during this period as sensitivity changes or changes of source filament required retuning of the source. The circled cluster of high 20 Ne/ 22 Ne ratios coincide with a significant decrease in the spectrometer s background H 2 O + and CO + 2, resulting in an incorrect pressure correction. Once background H 2 O + and CO + 2 concentrations stabilized the 22 Ne correction calibration was re-done. 79

80 3.5 Halogen analysis Sample preparation and irradiation Samples were analysed for halogen concentrations using an extension of the Ar-Ar dating method, making use of the following neutron-induced reactions to allow halogen and K concentrations to be determined in a noble gas spectrometer [Chris J. Ballentine et al., 2002; Böhlke and Irwin, 1992; Mark A. Kendrick, 2012; G. Turner, 1965]: 39 K(n,p) 39 Ar 37 Cl(n,γ) 38 Cl(β) 38 Ar 79 Br(n,γ) 80 Br(β) 80 Kr 127 I(n,γ) 128 I(β) 128 Xe. Prior to neutron irradiation in a reactor, hand-picked basalt glass chips were cleaned in acetone and deionised water in an ultrasound bath to remove any surface contaminants. Between and mg of each sample (0.5-2 mm size fragments) was wrapped in aluminium foil and placed, along with the Hb3gr hornblende neutron flux monitor (t = ± 5.3 Ma; [Jourdan et al., 2006]) and pyroxene from the Shallowater enstatite achondrite meteorite I-Xe standard [Gilmour et al., 2009; Hohenberg, 1968], in an evacuated silica glass tube. Sample and reactor details are given in Table 6. The fast and thermal neutron fluxes were determined from 39 Ar K and 38 Ar Cl production respectively in Hb3gr which contains ± wt% K and 2379 ± 16 ppm Cl [Roddick, 1982]. The neutron fluxes and irradiation parameters are given in Table 6. Conversion of noble gas isotopes to parent abundances of Cl, K, Br, I followed procedures outlined previously [Johnson et al., 2000; Kelley et al., 2005; Mark A. Kendrick, 2012]. For 80 Kr Br and 128 Xe I there is significant additional production via epithermal (resonant) neutrons in the reactor. Epithermal neutron production of 128 Xe I is measured directly from the Shallowater standard giving a factor of 1.7 higher production than that obtained from the thermal fluence alone. The epithermal neutron 80

81 flux calculated from the 128 Xe I production in the Shallowater is then combined with the resonance cross-section integral for 79 Br(n, ) 80 Kr to calculate the 80 Kr/Br production ratio. The value obtained indicates a factor of 1.3 higher production than calculated from only the thermal neutron fluence. Table 6: Irradiation details for halogen analyses MN10-a irradiation 2010 Analysed MN11-a irradiation 2011 Analysed A1 24-Mar-11 HL02 15-Aug-12 T1 25-Mar-11 JARL8 30-Apr-12 MID1 29-Mar-11 SN12 ol 17-Aug-12 HF1 29-Mar-11 HLO Aug-12 HOF10 30-Mar-11 ERIK3 01-May-12 VAL12 O HB2 HOF10ol FOG 4 KS1 HUS-2 LAVN-5 SA02 31-Mar Mar Mar Apr May May May May May-11 GOD2 not analysed Irad date 15/06/2010 Irrad date 10/05/2011 alpha 0.58 ± ±0.010 beta ± ±0.050 J ± ± position B2W Safari-1 reaactor, NECSA, Pelindaba, South Africa rodeo facility, high flux reactor, NRG Petten, The Netherlands (24 hr irradiation) fast neutron flux 2.56E+18 n/cm^2 1.37E+18 n/cm^2 thermal neutron flux 4.46E+18 n/cm^3 6.93E+18 n/cm^3 K Cl 38 Ar 39 Ar Hb3gr K/Cl ; Hb3gr 5.242; 39 Ar K K J 40 K e (mole/mole); / e 9.54; 40-4 K/K

82 3.5.2 Spectrometry Irradiated samples were loaded into separate arms of a furnace connected to a UHV extraction line and the MS1 argon-argon dating spectrometer. The MS1 and extraction line follow a similar set-up to that described above for the VG5400. A Faraday Cup and DeTech channeltron electron multiplier are available on the MS1 spectrometer for noble gas analysis. Noble gases were released from the samples using a step-heating technique with 3-6 steps between 600 o C and 1600 o C. The first sample was also run after hand-crushing under vacuum but released negligible gas during this step so future samples were not crushed. Released gases were cleaned for 10 minutes on a Zr-Al getter and trapped on an activated charcoal finger cooled by liquid nitrogen, before being released into the spectrometer for analysis. Ar, Kr and Xe were analysed in the same run. Furnace blanks were monitored regularly throughout and were typically 2.0 x 10-9 cm 3 STP 40 Ar, 1.4 x cm 3 STP 84 Kr and 3.6 x cm 3 STP 132 Xe at 1250 o C. Aliquots of air were used to calibrate noble gas concentrations and to correct for any mass discrimination. 40 Ar/ 36 Ar ratios were also determined during the halogen analysis. 82

83 4 Chapter Four Disequilibrium degassing model constrains the 3 He concentration of the MORB and OIB mantle sources 4.1 Abstract Models of the dynamics of eruptive degassing provide a unique insight into the preeruptive concentrations of the major volatiles (CO 2, H 2 O), noble gases, and the mantle reservoirs supplying them: a fundamental component in describing the accretion and evolution of the Earth. We investigate and develop a disequilibrium degassing model, exploring the parameters required to reproduce noble gas compositions observed in two ocean island basalt (OIB) and one mid-ocean ridge basalt (MORB) sample suites from the East Pacific Rise (EPR), Loihi Seamount, and Iceland. The original model assumed an identical loss of major volatile components for each degassing step. We recalculate the major volatile vapor phase composition for each degassing step, taking account of the degassing history over previous steps. Final noble gas elemental ratios, using the same eruption parameters, can differ by orders of magnitude from the original model s calculations. We further adapt our model variant to take into account both decompression-driven degassing during magma ascent and degassing at constant pressure during sample quenching. Our results show that elemental ratios and noble gas concentrations can be effectively decoupled by the different degassing stages of an eruption. Ascent rate determines whether degassing during magma ascent is modeled as predominantly closed or open system degassing. The closed system conditions that dominate for the slowly ascending MORB result in a less dramatic decrease in degassed elemental ratios than for the OIB models, consistent with the observed data. The MORB model also constrains the initial MORB source melt 3 He/ 22 Ne ratio, allowing that the MORB mantle could have a 3 He/ 22 Ne as low as the OIB 3 He/ 22 Ne, a 83

84 feature required by steady state mantle models. The two OIB sample suites show very similar noble gas ratios and concentrations despite a large difference in final eruption pressures. We propose this can fit our model if the effect of gas loss by quenching is negligible. Our model then shows that only a small amount of degassing takes place during magma ascent, so initial OIB He melt concentrations must be similar to or lower than MORB source melt concentrations, although greater initial CO 2 concentrations are required. This result is entirely consistent with mantle models which see a significant recycled component in the OIB-source mantle: such material would be expected to have higher CO 2 but lower 3 He concentrations than the depleted mantle. 4.2 Introduction The noble gases are invaluable geochemical tracers as their rare, inert nature means mantle signatures are little affected by chemical processes and, in the case of He and Ne, are not significantly recycled from the surface back into the mantle system. The noble gases include a number of radiogenic isotopes that are used as dating tools, tracers of parent isotopes and constrain the timing of events during the Earth s history. The non-radiogenic noble gas isotopes are not produced in significant quantities within the mantle and, because of their distinct isotope signatures, have been used as a key tracer of mantle reservoir preservation and interaction since the Earth s formation [Claude J. Allegre and Turcotte, 1986; C. J. Allegre et al., 1983; Chris J. Ballentine and Barfod, 2000; C. J. Ballentine et al., 2005b; E T Dixon et al., 2000; D. Harrison et al., 1999; M. D. Kurz et al., 1982; M. Moreira et al., 2001; Samuel Niedermann et al., 1997; Porcelli and Ballentine, 2002; Sarda et al., 2000]. Noble gas isotope ratios are well studied within the mantle, although disentangling the different noble gas origins is still ongoing [C. J. Ballentine and Holland, 2008; Holland and Ballentine, 2006; Greg Holland et al., 2009b; Mukhopadhyay, 2012; Parai et al., 2009; Stroncik et al., 2007]. Elemental noble gas ratios from different mantle reservoirs also have a key role to play in our understanding of the mantle system. For example, a difference in 3 He/ 22 Ne ratios, derived from 3 He/ 4 He and Ne 84

85 isotopes, between ocean island basalts (OIBs) and mid-ocean ridge basalts (MORBs) lead to the suggestion that the mantle is heterogeneous in 3 He/ 22 Ne [Darrell Harrison and Ballentine, 2005; Honda et al., 1993a; M. Moreira et al., 1996]. Models of the mantle differ in whether such heterogeneity is acceptable and how it should be distributed, with steady state mantle models - which involve a common origin for MORB and OIB He and Ne - requiring the same source 3 He/ 22 Ne ratio for the MORB and OIB sources ([Porcelli and Ballentine, 2002] give a more detailed discussion). There is as yet no consensus that reconciles the geophysical observations and numerical models suggesting whole mantle convection, with the geochemical observations used to argue for the preservation of at least two distinct mantle reservoirs over the lifetime of the Earth [Gonnermann and Mukhopadhyay, 2009]. Another critical use has been the identification of non-radiogenic elemental ratios of Ar, Kr and Xe in magmatic gas [Holland and Ballentine, 2006] that match those in fluids subducted to at least 100 km depth [Sumino et al., 2010] and together are indistinguishable from marine pore fluids. These observations have been used to argue for recycling of heavy noble gases and associated volatiles into the mantle system. Recycled noble gases could also account for the systematically lower 40 Ar/ 36 Ar and near-atmosphere Kr and Xe isotopic composition of OIB s, compared to MORB [C. J. Ballentine et al., 2002]. This distinction is important as it identifies the presence of airlike noble gases in the deep mantle to be the result of dynamic mantle processing and not the preservation of an air-like accretionary signature in an isolated deep mantle reservoir. Noble gas concentrations in the mantle are also critical in model generation and validation [C. J. Ballentine et al., 2002; Porcelli and Ballentine, 2002]. Four classic examples are 1) the 40 Ar mass balance between the MORB-source mantle and atmosphere that has been used to argue for ~50% mantle degassing efficiency over its lifetime [C. J. Allegre et al., 1996]; 2) the mantle heat/ 4 He ratio used to argue for a boundary layer separating heat and helium in the mantle [O'Nions and Oxburgh, 1983; P. E. van Keken et al., 2001]; 3) the helium paradox [Anderson, 1998b], which arises as models that explain high 3 He/ 4 He in OIBs by a high 3 He reservoir in the mantle 85

86 sampled by plumes are confronted with helium concentrations in OIBs that are often lower than those in MORB (this paradox is reinforced when noble gas elemental ratios that are the reverse of those predicted for equilibrium degassing are observed in OIB) and 4) the low 3 He concentration in the MORB-source mantle relative to U and a high 3 He/ 4 He that has been fundamental in developing models that require a 3 He flux from a volatile rich geochemical reservoir deep in the mantle system [Claude J. Allegre and Turcotte, 1986; O'Nions and Oxburgh, 1983; Porcelli and Wasserburg, 1995]. A key issue in the use of elemental noble gas compositions as a diagnostic tool is being sure which signatures are representative of the mantle source being investigated. Noble gas abundances are depleted by degassing during eruption and elemental ratios are fractionated. During magma ascent, volatiles such as the noble gases will degas from the melt as their solubility in melt decreases due to decompression. A quantitative description of the degassing process is necessary to extrapolate back from the observed noble gas composition in an erupted sample to the composition of the parent melt. Given enough time, the exsolved gas bubbles will be in equilibrium with the volatiles dissolved in the melt and the noble gas concentrations in the respective melt and gas phases will be controlled solely by their relative solubility at the eruption pressure and the concentration of the major volatiles. Degassing under such equilibrium conditions has been modeled with a single step, or with a mix of open and closed system steps [M. Moreira and Sarda, 2000; Sarda and Moreira, 2002; J. Yamamoto and Burnard, 2005]. However, volatiles diffuse through the melt into the gaseous bubble phase at different rates and, if the available degassing time is shorter than this diffusion rate, the relative diffusivity of the different volatiles will also affect the final composition of a sample. Recent noble gas degassing models that have also included disequilibrium effects have had some success in explaining features of the observed noble gas data. For example, [Paonita and Martelli, 2006; Paonita and Martelli, 2007] used a full bubble growth approach to model He-Ar-CO 2 systematics at mid-ocean ridges to conclude that the effects of disequilibrium degassing are seen in MORBs. Kinetic fractionation during degassing is also discussed as a possible mechanism for generating the observed array 86

87 of 4 He/ 40 Ar ratios [Aubaud et al., 2004; P Burnard et al., 2003]. [Gonnermann and Mukhopadhyay, 2007] used a disequilibrium degassing model to show that diffusive fractionation during degassing can provide a possible explanation for the Anderson [1998b] helium paradox. In their model disequilibrium degassing of a more volatile rich source with the high helium concentration can produce basalts that have a relatively low helium concentration with an elemental ratio inverse to that of equilibrium degassing. There are nevertheless many variables in this model s approach that are poorly constrained. We investigate here, and vary, the formulation of the model by Gonnermann and Mukhopadhyay [2007] and further investigate the model s sensitivity to the full range of input parameters. 4.3 Model formulation The starting point for our degassing model is that of Gonnermann and Mukhopadhyay s [2007] disequilibrium model (G & M s model). A full derivation for this model is given in Appendix A but average noble gas concentrations in the melt phase ( c i ) are given as c i 2 1 i psi c0 i vc vw psi 1 i (1) (see Table 7 for notation key and Appendix A for full derivation). To model open system magma degassing, with bubbles being continually formed and lost from the melt [Paonita and Martelli, 2007], equation (1) is applied repeatedly over several degassing steps, with initial noble gas melt concentrations for each step after the first given by final noble gas melt concentrations from the previous step. To model multi-step closed system degassing (for example during decompression degassing during magma ascent, when each step takes place at a different pressure), the original melt concentrations of the relevant volatiles must be taken into account at each step. 87

88 Equation (1) for each step becomes c i c w 0 1 vc vw psi 1 i c v v c ps p i i i. (2) c p c 0 process final noble gas melt concentrations from the previous step initial noble gas concentrations at the start of the closed system degassing In G & M s original model, the major volatile vapor phase composition is first calculated for equilibrium conditions using Dixon s [1997] model [J E Dixon, 1997]. The total major volatile concentration in the vapor phase is then calculated as if the entire multi-step process took place in a single step, with the disequilibrium factors, c and w, calculated using the total available degassing time. This vapor phase is divided equally between individual steps to allow the noble gas evolution to be modeled using equation (1) over the specified number of steps. By using a constant value for v c and v w at each step, this method does not take into account the effects of the changing major volatile concentrations in the melt. For example, in a near equilibrium process, a large proportion of CO 2 will move into the vapor stage in the first few steps. Then, as the melt moves from being over-saturated in CO 2 towards saturation, progressively smaller amounts of CO 2 will enter the vapor phase during subsequent steps. Additionally, G & M s original model does not allow that, during decompression degassing, each step may take place at a different pressure. 88

89 Table 7: Model parameters and default values Parameter Definition Default value EPR Loihi Iceland i Fractional gas retention for volatile species i during disequilibrium degassing. Determined from (A4 Appendix A). n/a n/a t degas Total available degassing time expressed as a fraction of the characteristic diffusion time of Ar log 10(t degas) is varied between -2 and +2 in 0.25 intervals as default D i Diffusivity of species i in melt (m 2 s -1 ) D He = ; D Ne = ; D Ar = ; D H2O = ; D Ar /D CO2 = 1 D He = ; 0.5D Ne = D Ar ; D Ar from equations (3) and (6); D H2O from equation (4); D CO2 from equation (5) p Eruption pressure (bar) S i Solubility constant for i (cm 3 STPg -1 bar -1 ) S Ar = 6 x 10-5 ; S Ne = 18 x 10-5 ; S He = 60 x 10-5 as default c 0 Initial concentrations and ratios of the relevant noble gas (concentration units: cm -3 STP g -1 ) 3 He: 1x10-8 ; 3 He/ 22 Ne = He: 1x10-9 ; 3 He/ 4 He = 8.8R A; 3 He/ 22 Ne = 11; 4 He/ 40 Ar = 5 3 He/ 4 He = 24.6R A; 3 He/ 4 He = 16.0R A; 3 He: 1x10-8 ; 3 He/ 22 Ne = 3.6; 4 He/ 40 Ar =5 CO 2 (H 2O) concentrations (per gram of rock) in the vapor v c(v w) phase at equilibrium: Calculated from Dixon s solubility model (Dixon, 1997) at the relevant temperature and n/a n/a pressure). vc v w CO 2 (H 2O) concentrations in the vapor phase accounting for any disequilibrium effects n/a n/a T Eruption temperature (K) as default 89

90 Parameter Definition Default value EPR Loihi Iceland [CO 2] 0 Initial melt CO 2 concentration (wt %) Varied between 0.07 and 0.82 in 0.05 intervals Varied between and 0.18 in intervals Varied between and 0.63 in 0.05 intervals C w Initial melt H 2O concentration (wt %) C w = [CO 2] Steps Degassing may take place over a number of open or closed system steps 10 Stage 2: 0-17 Stage 2: 0-5 Stage 2: 0-7 Stage 3: 10 Open vs. closed In open system degassing the vapor phase is lost at the end of each step whereas closed system degassing carries the existing vapor phase into the next step. Open Closed (stage 2) Open (stage 3) Open 90

91 In order to take these factors into account, we use a different sequence to model a multi-step degassing process. First, the available degassing time for each step is calculated as the total degassing time divided by the number of steps. i for each noble gas, c and w are then calculated using this value. v c and v w are calculated for the first step (using Dixon s [1997] model) and corrected for disequilibrium effects using c and w to give v c and v w. Assuming an open system, equation (1) is applied for the first step, after which c and w are recalculated using the new initial major volatile melt concentrations. This sequence is then repeated for the remaining number of steps, recalculating c and w at each step. If required, degassing at each step can be calculated at different pressures. The difference between the two models is only relevant for a multi-step process; for a single, closed system step, the two models are identical. Figure 15 shows schematically the difference between the two models for a three step scenario. 91

92 Figure 15: Schematic representation of the different model formulations. Original Model after Gonnermann and Mukhopadhyay (G & M) [2007]: First the CO 2 and H 2 O vapor phase composition is calculated for the total available degassing time. This vapor phase is then split equally between the number of steps and the noble gases diffuse at each step according to equation (1). Model variation: We calculate the CO 2 and H 2 O vapor phase composition at each step. A degassing time equal to the total degassing time divided by the number of steps is allowed for each step. This takes into account the evolution of the major volatiles over several steps rather than assuming an identical major volatile vapor phase at each step. 92

93 4.4 Model comparison Figure 16 shows calculated results for melt concentrations of 3 He and 3 He/ 22 Ne ratios matching the default parameters listed in Table 7. For the purpose of comparing the two models, volatile diffusivity and solubility, initial volatile concentrations, eruption temperature, number of steps and open vs. closed system conditions were chosen to match the conditions used in G & M s original calculations. A reference eruption pressure of 350 bar is used. We model a ten step process as the difference between the two models is only apparent for a multi-step process. The central parameter which determines the extent of disequilibrium conditions present during degassing is t degas. This is the total available degassing time (or for a multi-step model, the total available time for a single step) expressed as a fraction of the characteristic diffusion time of Ar (τ Ar ). As well as D Ar, t degas depends on the available degassing time, bubble radius and vesicularity; parameters that can vary widely under different eruption conditions. Hence we chose a range of values for t degas that covers the full range of conditions from full equilibrium to extreme disequilibrium. Model iterations over a range of values for t degas show that equilibrium conditions are achieved when log10 deg as t = 2 for the new model and0.25 for G & M s model; increasing the degree of disequilibrium beyond log 10 deg as t = -2 for both models does not significantly alter the results. A more detailed discussion of the choice of all parameters for particular sample suites follows this section. 93

94 Figure 16: 3 He concentrations and 3 He/ 22 Ne ratios in a melt after multi-step degassing following G & M s original model (left) and our model variation (right). Parameters are set to their default values, chosen to match G & M s original model formulation (see Table 7). The open square shows the initial 3 He concentration and 3 He/ 22 Ne ratio. The results grid represents calculated values for a range of initial CO 2 concentrations (from 0.07 to 0.82 wt. % in 0.05 intervals) and varying degrees of disequilibrium (from full equilibrium to extreme disequilibrium in 0.25 intervals). The G & M model results are shown in the faded grid on the right and differ by orders of magnitude from our model variation. For a single step process the two models are identical, so this result emphasizes the importance of taking into account the effect of major volatile evolution over a multi-step process. Results from both models are plotted for comparison in Figure 16. For G & M s model, the largest increase in the 3 He/ 22 Ne ratio (up to ~56000 from an initial ratio of 3.6) is seen at the highest initial CO 2 concentration (0.87 wt %) and under equilibrium degassing. These conditions also give the greatest extent of degassing and hence the lowest final 3 He concentration in the basalt glass of ~6x10-14 cm -3 STP g -1. The lowest 3 He/ 22 Ne ratio of 0.25 is seen at 10 deg as log t 1.75 (disequilibrium conditions), again at the highest initial CO 2 concentration. Similar to G & M s model, our model variation produces the highest 3 He/ 22 Ne ratios at time scales close to equilibrium and at the highest major volatile concentration, although the highest value produced is only 25 and is found slightly away from full equilibrium conditions: In general, G & M s model produces much higher 3 He/ 22 Ne ratios for a given value of t degas than our model 94

95 variation. The lowest 3 He concentration of ~2.6x10-14 cm -3 STP g -1 for our model is of the same order as that produced by G & M s model, and also occurs at the highest initial major volatile concentration, although it occurs under disequilibrium conditions ( 10 deg as log t 0.75 ) rather than at equilibrium. For our model, the calculated range of 3 He/ 22 Ne ratios falls considerably lower than in G & M s model, with the lowest value of 8x10-5 found at 10 deg as log t Some of the produced trends are common to both models. Higher initial CO 2 concentrations in the melt will give rise to a larger vapor phase on degassing than would be produced by a smaller initial concentration. This higher initial concentration leads to more extensive noble gas loss and fractionation effects are more apparent than for lower initial major volatile concentrations. Both models also show a general trend of lower 3 He/ 22 Ne ratios as the extent of disequilibrium increases. However, over a multi-step process, the difference in 3 He concentrations and 3 He/ 22 Ne ratios calculated by the two models from the same initial conditions can be orders of magnitude apart. Our model variation allows for much more dramatic disequilibrium fractionation of the noble gases than G & M s original model, whereas increases in 3 He/ 22 Ne ratios near equilibrium are less marked. This difference is due to the fact that the original model defines equilibrium over the entire degassing time whereas for equilibrium conditions to be present in our model variation, equilibrium must be reached at each individual step. 4.5 Modeling MORB and OIB sample suites Noble gas degassing models are essential in unraveling the link between the (often unknown) starting composition of the mantle and the observed array of results seen in MORBs and OIBs. Three basaltic sample suites were chosen for comparison with our degassing model to represent a range of differing eruption conditions. The first is a suite of MORB data from the East Pacific Rise taken from [Samuel Niedermann et al., 95

96 1997]. Two OIB sample suits are also considered, one from the Loihi seamount [Valbracht et al., 1997] and another from an Icelandic subglacially erupted sample suite [M. Moreira et al., 2001]. These suites were chosen as all the samples within a single suite were collected at similar depths (or were from the same subglacial unit for the Icelandic samples), indicating similar eruption pressures. This reduces the number of variables in the model. In applying our model variation to the three sample suites, each eruption is divided into three separate stages of degassing during which different conditions are present (see Figure 17). Stage 1 (magma chamber) represents the melt and vapor phases reaching equilibrium before the start of the magma ascent, at the depth at which the magma is stored. This is modeled as a single, closed system step according to equation (1). Stage 1 (magma chamber) will only be significant if the melt is oversaturated in volatiles at the depth of magma storage. For the storage depths and initial major volatile concentrations assumed for the three sample suites used in this study (see sections 4.6 to 4.8), melts are calculated to be undersaturated during stage 1 (magma chamber) and so stage 1 (magma chamber) will not be discussed further. In stage 2 (magma ascent), degassing is driven by decompression as the magma ascends to the surface. This is modeled as a multi-step process and can be open or closed system; each step is modeled according to equation (1) (open system) or equation (2) (closed system) with final volatile concentrations from one step representing initial volatile concentrations for the next step. Pressure is gradually decreased at each step at a rate that depends on the ascent rate of the magma (the rate of change of pressure with depth is assumed to be constant). The final stage, stage 3 (quenching), takes place on the surface and represents degassing that takes place as the melt cools and finally quenches at the eruption depth. Stage 3 is modeled as a second multi-step process, this time at constant pressure. As any volatiles degassed during stage 3 (quenching) are lost to the surrounding water, this stage is always modeled as an open system process. 96

97 Figure 17. Schematic sketch of the three stage model used to calculate final melt concentrations of He, Ne and Ar for the three sample suites. Stage 1 (magma chamber) will only be relevant if the melt is oversaturated in CO 2 under the pressure and temperature conditions found in the magma chamber. For the three sample suites investigated in this study, stage 1 is not relevant. Stage 2 (magma ascent) can be open or closed system, depending on the magma ascent rate, and each degassing step takes place at a different pressure as the melt ascends. Stage 3 (quenching) takes place at constant pressure and under open system conditions as bubbles are continually lost into the surrounding water. 97

98 4.5.1 Number of degassing steps There are two parameters to set here, the number of steps during stage 2 (magma ascent) and the number of steps during stage 3 (quenching). An accurate setting for the number of steps would require a full bubble growth model. An approximation can be derived by modeling a bubble growing in the melt as a sphere moving in a viscous fluid at low Reynolds number (see Appendix B for details). The melt in a single eruption can be divided up into melt cells, each of which surrounds a single bubble. Knowing the speed of a bubble relative to the ascent rate of the melt then constrains how long a bubble will take to move through one melt cell ( timescale of bubble loss ). The timescale of bubble loss must be individually calculated for each step until the total time elapsed equals the total time available for the eruption, which is determined from ascent rate and depth of the magma chamber. This will then give the number of steps for stage 2 (magma ascent). The number of steps for stage 2 is calculated separately for each suite as ascent rates and pressure differ across the three sets of samples. Stage 2 (magma ascent) degassing takes place in between 0 and 17 steps (the exact number of steps depends on initial CO 2 concentration) for the EPR sample suite, 0 to 5 steps for the Loihi sample suite and 0 to 7 steps for the Icelandic sample suite. The number of steps is calculated following the method set out in Appendix B. For stage 3 (quenching), the number of degassing steps can be determined using the same method, the only difference being that the total available degassing time is now determined by the quenching time of the sample in question. Paonita and Martelli [2007] suggest quench times ranging from minutes to hours, which equates to between 1 and 100 steps. During quenching, degassing occurs at the lowest pressure of the entire eruption process and the first few degassing steps have a large effect on noble gas concentrations and ratios. For example, for the Icelandic sample suite, moving from a quenching time of 21 seconds to 3 minutes equates to the difference between a one step and a ten step stage 3 model, and the resulting difference in final noble gas compositions is large. However, moving from a quenching time of 17 minutes to 34 minutes (50 to 100 steps) produces little difference in final compositions. As the time 98

99 of quenching increases, changes to the final noble gas composition of the sample become less dramatic. As the samples taken for noble gas studies are usually the fast quenching, glassy rims of pillow basalts, a representative value at the lower end of the spectrum - 10 steps - is used as a default and reference value Open or closed system conditions A further parameteris whether each degassing step should take place under open or closed system conditions. During stage 2, a single melt cell will see something inbetween a completely open and a completely closed system. Newly nucleated bubbles will rise through the melt due to their greater buoyancy. To a single melt cell, this looks like an open system. However, old growing bubbles will also rise through the melt. In this case to a single melt cell, an old bubble passing through will look like a closed system, since it already contains degassed noble gases from the melt cells below. The key parameter in determining the interplay between these two factors is ascent rate. [Bottinga and Javoy, 1990] suggested that melt must be oversaturated in CO 2 by a factor of between 1.5 and 7 before bubble nucleation can occur. Setting this factor to a default value of 4, the CO 2 degassing path of an EPR melt cell with an ascent rate of 0.1 ms -1 shows only three nucleation events (see Figure 18) [Paonita and Martelli, 2006]. With so few nucleation events, closed system conditions will dominate and models for this sample suite used a closed system for stage 2 (magma ascent). However, for OIB samples with a faster ascent rate (1 ms -1 - [Paonita and Martelli, 2006]), the oversaturation levels required for bubble nucleation are reached many times during the ascent. For this reason we have used an open system to model stage two for OIB samples. This has its limitations, as some closed system degassing will occur as well, but including both factors is beyond the scope of this model. Stage 3 (quenching) is always modeled as an open-system process, as the melt is now no longer ascending and bubbles are continuously lost into the surrounding water. 99

100 Figure 18. Melt oversaturation and CO 2 concentrations during stage 2 degassing are shown for a slow (0.1 ms -1 ) and a fast (1 ms -1 ) eruption starting at bar. Bubble nucleation is set to occur when CO 2 oversaturation exceeds four (see text): For the slow eruption (EPR sample suite) there are only three nucleation events (dashed line), whereas the fast eruption (OIB sample suites) shows multiple nucleation events (dotted line). The slow eruption shows only three generations of bubbles. As each melt cell sees old bubbles which already contain degassed noble gases, closed system conditions will be dominate. In the fast eruption, new bubble generations are continually formed and ascend through the melt due to their buoyancy allowing open system conditions to dominate. 100

101 4.5.3 Diffusivity The relative diffusivity of the noble gases and major volatiles controls the extent of noble gas fractionation caused by disequilibrium degassing. Absolute noble gas and major volatile diffusivity will affect the time scale at which disequilibrium degassing becomes important. A number of studies have investigated Ar, H 2 O and CO 2 diffusivity in silicic melts (see review by [Zhang et al., 2007]) and have shown that D H2O and D CO2 depend on temperature (T), pressure (p) and water content of the silicic melt (C w ) as follows [Zhang, 2007]: p pcw ln DAr: silicic ; (3) T T DHO ln (4) Cw T DCO 2 is determined from D Ar [Zhang, 2007] as follows: D CO D. (5) 2 : Ar silicic Equations (3) to (5) give values for a silicic melt, but this study is concerned with basaltic melts. [Nowak et al., 2004] investigated Ar and CO 2 diffusivity across a range of melt compositions and found that Ar diffusivity increases from rhyolitic to basaltic melts by a factor of approximately four whereas anhydrous melt composition. Using the silicic values for D CO 2 = m 2 s -1, D CO 2 is relatively independent of D HO and 2 D CO 2 this gives D HO = m 2 s -1 and D 2 Ar = m 2 s -1 [Zhang, 2007] at 300 bar, K, C w = 0.37 wt%. This value for D Ar is considerably lower than the value 101

102 determined from an empirical study by [Lux, 1987] with a value of m 2 s -1, (even taking into account the higher temperature of Lux s experiments) but is higher than the value used by Gonnermann and Mukhopadhyay [2007] (from Nowak [2004]) of m 2 s -1. However it compares well with the diffusivity of K, ( m 2 s -1 ) used by [Albert Jambon et al., 1986] as a proxy for Ar. Equations (3) to (5) will be used to determine D HO and 2 DCO 2 for all three sample suites, with D Ar given by D 4D. (6) Ar: basaltic Ar: silicic The difference in calculated results made by using equations (3) to (6) to calculate volatile diffusivity, rather than the constant (default) values shown in Table 7, can be seen in Figure 19. A key difference is the effect of differing initial H 2 O concentrations. Since water does not degas significantly until lower pressures are reached (<50 bar), the vapor phase is generally dominated by CO 2 and it might be expected that initial H 2 O content has little effect on degassing. However, Figure 19 shows that the dependence of diffusivity on C w means melts with different initial H 2 O can follow different degassing paths, with higher C w allowing much lower 3 He/ 22 Ne ratios to be achieved under disequilibrium conditions: water content becomes a key parameter. 102

103 a b C W doubled Figure 19. New model results for 3 He and 22 Ne with D CO 2, D HO and D 2 Ar calculated using equations (3)-(6) (solid lines). All other parameters are set to default values except C w which is doubled in the right-hand figure (b) to test the model s sensitivity to increased water content. The dashed grid represents results calculated using the defaults parameters (Table 7) for comparison. Under extreme disequilibrium conditions, calculated results are similar to those calculated using the default parameters. Nearer to equilibrium, and at the highest CO 2 concentration of 0.82 wt. %, He concentrations and elemental noble gas ratios can be lowered by an order of magnitude or more. The high water content used in Figure 19b is unlikely to be realistic but the results show that even though H 2 O does not significantly degas at 350 bar, initial water content can impact calculated results due to the dependency of melt water content. D CO 2 and D Ar on 103

104 The relative diffusivity of He and Ne is not well known. Lux s empirical study [1987] gives values for diffusivity across the entire range of noble gases. Gonnermann and Mukhopadhyay [2007] use D Ar ~ 0.5D Ne ~ 0.001D He (see Table 7). By comparison, values from Lux (1987) give D Ar ~ 0.25D Ne ~ 0.125D He. Starting from the same value for D He, the difference in calculated results can be seen in Figure 20. Helium diffusivity is fast enough that it can be assumed to degas at equilibrium conditions under all realistic eruption conditions. In contrast neon and argon degassing can take place under disequilibrium conditions. Neon diffusivity is considerably less well studied than argon and this adds an extra uncertainty into the calculations for 3 He and 22 Ne, which is not an issue for the 4 He and 40 Ar calculations. 104

105 Figure 20. New model results calculated using noble gas diffusivity relationships from Lux [1987]: D Ar ~ 0.25D Ne ~ 0.125D He (solid lines) rather than the default relationship of D Ar ~ 0.5D Ne ~ 0.001D He (dashed lines) used by Gonnermann and Mukhopadhyay [2007] The higher value used here for D He /D Ne results in calculated values that generally show higher elemental ratios and He concentrations than those calculated using the default values for noble gas diffusivity, as Ne diffuses more quickly into the vapor phase and disequilibrium effects are not so prevalent. We continue to use the default relationship for noble gas diffusivity shown in Table 7, but uncertainties in D Ne are an important factor in the accuracy of this model. 105

106 4.5.4 Solubility Noble gas solubility in melt is determined via a range of experimental techniques carried out at high temperature and pressure (see review by [Paonita, 2005]). Solubility in basaltic melts decreases as the noble gases get heavier and shows some dependence on melt composition, a linear relationship with pressure, and a weak dependence on temperature [Paonita, 2005; Shibata et al., 1998]. An empirical study of noble gas solubility by Jambon [1986] gives values for tholeiite as 56 x 10-5 cm -3 STP g -1 bar -1, 25 x 10-5 cm -3 STP g -1 bar -1 and 5.9 x 10-5 cm -3 STP g -1 bar -1 for helium, neon and argon respectively. Other studies found similar results for helium, neon and argon [Carroll and Stolper, 1993; Lux, 1987]. Although these studies are in reasonably good agreement, [Paonita et al., 2000] found helium solubility in basaltic melts increased with the addition of H 2 O. They found an increase in solubility by a factor of around 3 on the addition of 3 wt. % H 2 O. This is a much higher value than the maximum water content of 0.63 wt % investigated in this study, but it should be considered that this may introduce a further sensitivity of the model to initial H 2 O content. Higher solubility acts against the decrease in light to heavy ratios caused by diffusive fractionation as this would result in higher 3 He/ 22 Ne and 4 He/ 40 Ar ratios under disequilibrium conditions than for the default solubility parameters Temperature Although magma temperatures may have varied across the three sample suites, modeling results for both the original model and our model variation proved relatively insensitive to this parameter. For example, a temperature increase of 200 o changes ratios and concentrations by less than 6%. For this reason the temperature is set at a constant K. 106

107 4.5.6 East Pacific Rise (EPR) sample suite The initial CO 2 content of the MORB source mantle is hard to determine as even deep sea samples have degassed to some extent. This is less of a problem for H 2 O, as water will not significantly degas until low pressures (< 50 bar) are reached [Dixon, 1997]. Taking into account the effect of both degassing and crystallization, [Nichols et al., 2002] estimates the water content of the Icelandic mantle to range from 165 ppm away from the plume to 920 ppm near the plume, suggesting higher water content in the plume source mantle than in the MORB source mantle. The lower value of 165 ppm, within the range suggested by [Gose et al., 2009] of ppm, is used in our model as an estimate of MORB source mantle H 2 O content. For this model, pre-eruptive CO 2 concentrations in the melt are varied between the minimum EPR estimate of 44 ppm [Saal et al., 2002] and a maximum of 1800 ppm in 150 ppm intervals. Initial noble gas ratios and concentrations in the melt are determined as follows: 3 He concentrations follow Gonnermann and Mukhopadhyay s [2007] estimate of 1x10-9 cm -3 STP g -1 for the MORB source mantle. This gives a CO 2 / 3 He ratio of between 3x10 9 and 1x10 10 which compares well to the estimated MORB CO 2 / 3 He value of 2-4 x 10 9 [B. Marty and Jambon, 1987]. 3 He/ 4 He is then set to the sample suite average of 8.8 R A (where R A is the atmospheric ratio of 1.399x10-6 ): 4 He/ 21 Ne and 4 He/ 40 Ar are set to the known mantle production ratios of 2.22 x 10 7 and 1-5 respectively, and 3 He/ 22 Ne is set to the MORB source mantle estimate of 11 [D W Graham, 2002]. The degassing stages take place as follows: Stage 2 (magma ascent): The EPR is considered to have shallow magma chambers so the magma chamber depth is set to 1 km below the ocean floor [Hussenoeder et al., 1996]. Stage 2 therefore starts at 1000 bar (using a pressure gradient of 1bar m -1 [Head et al., 1996]). The melt is oversaturated at this depth for initial CO 2 concentrations over 493 ppm and, as discussed in section 4.2, bubble nucleation only begins when the melt is oversaturated by a factor of 4 which is not a scenario that arises for this sample suite. The ascent rate for the EPR suite is set to 0.1 ms -1 [Paonita and Martelli, 2007]. Using the method described in section 4.1, the number of steps is between 0 and 17, depending on [CO 2 ] 0 (see Appendix B). log t as is varied between full disequilibrium ( 10 tdeg as 10 deg log 2 ) to equilibrium 107

108 conditions ( 10 deg as log t 2). Stage 3 (quenching): Degassing during quenching at the sea floor takes place at 273 bar based on the average sample depth of m [Samuel Niedermann et al., 1997]. log10 deg as t is again varied between -2 and Loihi sample suite The CO 2 content in OIB samples is often higher than in MORB: We vary initial CO 2 melt content from to 0.63 wt% [J E Dixon and Clague, 2001]. Water content for OIB source mantle is set to 920 ppm [Nichols et al., 2002]. Other initial noble gas melt concentrations are set as follows: Initial 3 He is set to 1x10-8 cm -3 STP g -1, giving a CO 2 / 3 He ratio of between 5x10 8 and 6x10 9 [Gonnermann and Mukhopadhyay, 2007]; 4 He concentrations are then determined using the average sample suite 3 He/ 4 He ratio of 24.56; 21 Ne and 40 Ar are then determined from the mantle production ratios as for the EPR data suite; Initial 22 Ne concentration is based on the OIB source mantle and 3 He/ 22 Ne estimate of 3.6 [D W Graham, 2002]. Stage 2 (magma ascent): [Clague, 1988] proposed a reservoir at 16 km below sea level for the Loihi seamount. A shallower chamber proposed at 8-10 km is likely to be a more recent addition [Garcia et al., 1998] and hence not involved in the production of the alkalic lavas investigated by Valbracht et al. [1997] and used here. Therefore we have modeled stage 2 as beginning at a depth of 16 km below the sea floor. Magma ascent rates at the Hawaiian hot spot are generally faster than the 0.1 ms -1 used to model a MOR eruption. An ascent rate of 1 ms -1 is used in modeling degassing at the Loihi seamount [Paonita and Martelli, 2006]. This gives an estimate for the number of steps during decompression of between 0 and 5 (depending on [CO 2 ] 0 ). As for the EPR data suite, we vary log10 tdeg as disequilibrium ( log10 tdeg as 2 ) to equilibrium conditions ( 10 deg as between full log t 2). Stage 3 (quenching): Degassing during quenching at the sea floor takes place at bar based on the average sample depth of 4649 m [Valbracht et al., 1997]. log 10 deg as t is again varied between -2 and

109 4.5.8 Iceland sample suite The parameters for the Iceland sample suite [M. Moreira et al., 2001] are identical to those used for the Loihi suite with the following exceptions: this sample suite is subglacially erupted so the final eruption pressure is set to 100 bar (this is an estimate as ice thickness during the eruption is unknown). This difference in final eruption pressure changes the maximum estimate for the number of steps in stage 2 (magma ascent) to 7. The average 3 He/ 4 He ratio for this sample suite is 16 rather than the used for the Loihi sample suite. 4.6 Model results Calculated results are compared to sample data for all three suites in Figure A small decrease in initial helium concentrations (from 1 x 10-9 cm 3 STP g -1 to 0.9 x 10-9 cm 3 STP g -1 ) allows our EPR model results to match measured noble gas concentrations and 4 He/ 40 Ar seen in the sample suite [Samuel Niedermann et al., 1997]. However, the model s 3 He/ 22 Ne ratios do not cover the range of observed 3 He/ 22 Ne ratios, with five data points showing ratios lower than model results seen in the sample suite data (Figure 21). The use of full closed system conditions could explain this discrepancy: As discussed above, the degassing melt-vapor system is not a fully closed system, although closed system conditions will dominate. The influence of some open system degassing in stage two would be to produce a broader range of noble gas ratios whilst still allowing an initial 4 He/ 40 Ar ratio consistent with the mantle production ratio. However, the initial 3 He/ 22 Ne ratio would need to be lower than 4.4 for the degassing model to match the observed data range. The noble gas ratios and concentrations for the Icelandic sample suite fall within the model results (Figure 22). Model results also cover the range of noble gas ratios seen in the Loihi sample suite (Figure 23) although initial 3 He concentrations must be reduced to 2x10-9 from 1x10-8 cm 3 STPg -1 to explain the Loihi samples helium 109

110 concentrations. The higher initial major volatile content of the two OIB sample suites allow lower elemental ratios and concentrations to be generated under disequilibrium conditions than for the EPR sample suite. The lower eruption pressure for Iceland gives calculated final basalt glass noble gas concentrations which can be lowered by up to five orders of magnitude compared to two orders of magnitude for the Loihi calculations (relative to the original melt concentration). Unless initial helium concentrations are further lowered, the data from Loihi s sample suite only match the model towards the highest values of the possible initial CO 2 melt concentrations (between 0.42 and 0.62 wt. %): this is in contrast to the Icelandic data where, due to the lower eruption pressures, an initial CO 2 concentration at the lower end of the allowed range (between 0.12 and 0.22 wt. %) fits the sample suite data. 110

111 Stage 2 (magma ascent) Stage 3 (quenching) Figure 21. Calculated results and measured compositions [Niedermann et al., 1997] for He, Ne and Ar for the EPR sample suite. The open square represents the initial noble gas composition used for modeling. Crosses represent measured noble gas compositions (corrected for atmospheric contamination) for the EPR sample suite. The grid of results represents calculated noble gas compositions for a range of initial CO 2 concentrations and conditions from full equilibrium to extreme disequilibrium (see Table 7 for parameter details). Figures 21a and 21c show results after stage 2 (magma ascent) for [ 3 He] vs. 3 He/ 22 Ne and [ 4 He] vs. 4 He/ 40 Ar respectively Final results after stage 3 (quenching) are shown in 21b for [ 3 He] vs. 3 He/ 22 Ne and 21d for [ 4 He] vs. 4 He/ 40 Ar. The final grid of results in figures 21b and 21d show a smaller range of final He/Ne and He/Ar ratios than for the two OIB models and these ratios are also invariably higher than their starting point which is not the case for the OIB results (see Figures 22 and 23). This is due to a combination of the closed system conditions in stage 2 (magma ascent) and low initial CO 2 concentrations. Some of the lower 3 He/ 22 Ne ratios are not covered by the final grid of results (21b). A combination of partial open system conditions for stage 2 and a lower initial 3 He/ 22 Ne (lower than 4.4) will allow the model to cover the full array of measured 3 He/ 22 Ne compositions. This demonstrates the model s ability to constrain initial noble gas compositions. 111

112 Stage 2 (magma ascent) Stage 3 (quenching) Figure 22. Calculated results and measured compositions [Valbracht et al., 1997] for He, Ne and Ar for the Loihi sample suite. Symbols are as for Figure 21. Figures 22a and 22c show results after stage 2 (magma ascent) for [ 3 He] vs. 3 He/ 22 Ne and [ 4 He] vs. 4 He/ 40 Ar respectively. Final results after stage 3 (quenching) are shown in 22b for [ 3 He] vs. 3 He/ 22 Ne and 22d for [ 4 He] vs. 4 He/ 40 Ar. The higher initial CO 2 concentrations for this sample suite allow more dramatic reductions in He concentrations than for the EPR sample suite. A reduction of less than an order of magnitude in the initial 3 He and 4 He concentrations allows calculated results to match the samples with the lowest measured He concentrations. Alternatively much lower initial 3 He and 4 He concentrations coupled with fewer steps in stage 3 degassing would also cover the range of measured compositions (see Figure 24). Noble gas ratios in stage 2 (magma ascent), shown in 22a and 22c, can be altered by more than an order of magnitude while helium concentrations are reduced by less than an order of magnitude due to the high pressures involved for most of this stage. In contrast, figures 22b and 22d shows that during stage 3 (quenching), both ratios and concentrations can be altered by orders of magnitude. This means it is possible for noble gas ratios to be significantly fractionated during stage 2 (magma ascent) without much decrease in concentrations (for example, a degassing path which ends at the star in 22a). Conversely, the star in figure 22b shows the end point of a degassing path where a sample can be significantly degassed during stage 3 (quenching) without much change in noble gas ratios. Care must be taken when implying links between noble gas elemental ratios and concentrations. 112

113 Stage 2 (magma ascent) Stage 3 (quenching) Figure 23. Calculated results and measured compositions [Moriera et al., 2001] for He and Ne (Ar data is not available) for the Iceland sample suite. Symbols are as for Figures 21 and 22. Figure 23a shows results after stage 2 (magma ascent) for [ 3 He] vs. 3 He/ 22 Ne. Final results after stage 3 (quenching) are shown in 23b. The low eruption pressure during stage 3 degassing (100 bar) results in the potentially dramatic reduction in He concentrations during stage 3 degassing (a reduction of up to five orders of magnitude). The calculated results match the range of measured compositions, but at much lower initial CO 2 concentrations (0.12 to 0.22 wt. % CO 2 ) than for the Loihi sample suite (0.42 to 0.62 wt. % CO 2 ): It is surprising that two OIB sample suites should have originated from melts with such different initial 3 He and CO 2 compositions. An alternative explanation is that initial He concentrations are much lower (~ cm 3 STP g -1 3 He and ~ cm 3 STP g -1 4 He) and degassing during stage 3 is minimal (i.e. quenching happened very quickly). This scenario is shown in Figure

114 4.7 Discussion Three principle observations can be picked out of the model results that have important implications for the interpretation of noble gas measurements. Both noble gas elemental ratios and concentrations are often used in interpreting a sample s history within the mantle [Claude J. Allegre and Turcotte, 1986; Anderson, 1998b; C. J. Ballentine et al., 2002; O'Nions and Oxburgh, 1983; Porcelli and Ballentine, 2002; P. E. van Keken et al., 2001]. This study shows that care is needed here as both G and M s original model and our model variation allow that, under certain conditions, extensive degassing can take place without much fractionation of elemental ratios (see Figure 21) [Gonnermann and Mukhopadhyay, 2007]. This means, for example, that 4 He/ 40 Ar* and 4 He/ 21 Ne* ratios close to mantle production values do not necessarily indicate an undegassed sample the sample may be significantly degassed and other heavier noble gas elemental ratios may have been more significantly altered. Our model results also provide a further caveat. Figure 21(a&c) shows an example of how it is possible for elemental ratios to be significantly fractionated during stage 2 (magma ascent) without extensive loss of the noble gases to the vapor phase. Elemental ratios and noble gas concentrations can be effectively decoupled by the different degassing stages of an eruption. When interpreting links between concentrations and ratios in erupted samples, the degassing history of a sample is of prime importance. Secondly, it is observed that OIBs generally have lower light to heavy elemental ratios than MORBs [Fisher, 1985]. OIBs higher initial CO 2 compared to MORBs can mean they are more extensively degassed [Gonnermann and Mukhopadhyay, 2007]. Under disequilibrium conditions, this can lead to lower elemental ratios. Our model adds a further consideration. Ascent rate determines whether open or closed system degassing dominates during magma ascent. Lower ascent rates often indicate that closed system degassing will dominate (see Figure 18). Closed system degassing does not lower elemental noble gas ratios to the same extent as open system degassing (see Figure 21), so samples that have seen slower ascent rates may well show higher He/Ne and He/Ar ratios. This has the further implication that MORB initial 3 He/ 22 Ne ratios can be similar to OIB initial ratios and still generate 114

115 the higher final 3 He/ 22 Ne ratios observed in MORBs: a similar 3 He/ 22 Ne ratio in the MORB and OIB source mantles is a requirement of mantle models that involve a common source for this ratio, such as steady state models of the mantle. Finally, Gonnermann and Mukhopahyay s [2007] explanation for the apparent helium paradox is that the higher initial CO 2 content of the OIB source mantle means final He concentrations in OIBs are lower than MORBs after degassing, even starting with a higher OIB mantle source helium concentration. If initial OIB source mantle helium concentrations are higher than MORB source mantle values, our model results for Iceland and Loihi still require a higher initial OIB source mantle CO 2 content than for the EPR sample suite. It is also possible for our model to reproduce the OIB results from an initial melt 3 He concentration, which is higher than for the EPR sample suite. However, given that stage 3 (quenching) degassing takes place at such different pressures for the two OIB sample suites, very different final He and Ne concentrations would be expected. In fact, the two sample suites are very similar and so a counterbalancing parameter to eruption pressure needs to differ significantly between the two suites. For example, different initial CO 2 concentrations (between 0.42 and 0.62 wt. % for Loihi compared to between 0.12 and 0.22 wt. % for Iceland) combined with different initial helium concentrations (1x10-8 cm 3 STPg -1 for Iceland but 2x10-9 cm 3 STPg -1 for Loihi) would be needed for our model to match the observed OIB data. Such different initial concentrations would mean Iceland s and Loihi s source mantle s CO 2 / 3 He and initial 3 He concentrations differed by an order of magnitude. Although very different quenching conditions cannot be ruled out, a simpler explanation is that both samples quenched very quickly (a few minutes or less) and hence experienced most of their degassing during stage 2 (magma ascent). Modeling the degassing of the Loihi and Iceland samples assumes similar conditions are present during stage 2 (magma ascent), so similar end results, as observed in the two sample suites, would then be predicted. Figure 24 shows how this model scenario can provide a good fit to the sample suite data if an Initial 3 He concentration of cm 3 STPg -1 and initial 3 He/ 22 Ne of 2.4 are used. All other parameters unchanged, except that the number of steps for stage 3 (quenching) is set to 0 so that all degassing takes place during stage 2 (magma ascent). Higher than 115

116 MORB initial CO 2 concentrations are still required in this scenario for model results to match He and Ne data. A potential difficulty with this interpretation is that the quenching conditions for Loihi and EPR are very similar, but at least one quenching step is needed for stage 3 degassing if the EPR model is to match the sample suite. However the difference in quench times is only of the order of 40 seconds or so, a factor that could easily be accounted for by small differences in quenching conditions or even sampling at a different depth into the pillow rind. Allowing no degassing during the stage 3 (quenching) process for both OIB sample suites has consequences for the possible range of initial He concentrations. Contrary to the conclusion reached by G & M, initial OIB-source melt helium concentrations must then be similar to or lower than those used in the EPR model (~1 x cm -3 STP g -1 ). Figure 24: Calculated results (for stage 2 - magma ascent - degassing only) and measured compositions for He and Ne for the Loihi (24a) and Iceland (24b) sample suites. Model results have been calculated using the same parameters as for figures 22a and 23a except that initial 3 He concentration is set to cm 3 STPg -1 and 3 He/ 22 Ne is set to 2.4. The results show how stage 2 degassing produces similar results for both sample suites. If both sample suites quenched very quickly, allowing stage 2 to dominate the degassing process, this would simply explain the very similar sample suite data for Iceland and Loihi even though they have very different final eruption pressures. However, such an explanation requires lower initial helium concentrations than for the MORB sample suite. This would make sense within mantle models which involve a significant recycled component within the OIB source mantle, but such low initial He concentrations for the two OIB source melts would still require a further source of helium to explain the high 3 He/ 4 He ratios observed in these samples. 116

117 Given the greater extent of melting of MORBs compared to OIBs this would lead the OIB source mantle to have a 3 He concentration up to four times lower than the MORB source mantle [Klein and Langmuir, 1987; Watson and McKenzie, 1991]. Lower OIB-source mantle initial 3 He concentrations are problematic for models of the mantle that require the OIB-source mantle to be a less depleted reservoir rich in helium. However, if the OIB source mantle contains significant recycled oceanic crust [Brandenburg and van Keken, 2007b; Hofmann, 1997; Holland and Ballentine, 2006; White, 1995], then low He concentrations would be unsurprising. A small volatile rich ( 3 He) component, perhaps sampled from the deep mantle, would then account for high OIB 3 He/ 4 He [C. J. Ballentine and Holland, 2008]. The recycled nature of the OIB-mantle source could also explain the higher major volatile concentrations that lead to disequilibrium degassing paths giving lower than MORB 4 He/ 40 Ar ratios [Fisher, 1985]. Most mantle models require a He source to explain both OIB and MORB 3 He/ 4 He ratios and He concentrations [Claude J. Allegre and Turcotte, 1986; O'Nions and Oxburgh, 1983], but our results suggest that its primary source is not a helium-rich reservoir in the mantle sourcing the OIB melts. 117

118 4.8 Summary and conclusions We vary Gonnermann and Mukhopadhyay s (G & M s) [2007] disequilibrium degassing model to take into account the evolution of the major volatiles over a multi-stage process. We further extend the model to take into account the different conditions present during the magma ascent and quenching stages of an eruption and apply our model to two OIB and one MORB sample suites. The key results of this study are as follows: o o o o o Our model variation can produce calculated results that are orders of magnitude different from those produced by G & M s original model: in particular our model allows much lower elemental noble gas ratios to be reached under disequilibrium conditions. Our degassing model is relatively insensitive to eruption temperature. In contrast, initial CO 2 concentration, pressure, diffusivity and degree of disequilibrium can affect noble gas concentrations and elemental ratios by an order of magnitude or more. Initial water content is a more important parameter than would be expected from vapor phase composition considerations alone. In our three stage model, stage 2 (magma ascent) can have a large effect on elemental ratios whereas stage 3 (quenching) has the most effect on noble gas concentrations. This means that different eruption conditions can affect noble gas ratios and concentrations, and it may be difficult to infer properties of the source mantle from links between these two values. Closed system conditions during stage 2 (magma ascent) degassing do not lower elemental noble gas ratios as much as open system degassing: the ascent rate is the key factor in determining whether open or closed system conditions dominate. This means it should be expected that samples from eruptions with a slow ascent rate (such as the EPR sample suite) will have higher elemental noble gas ratios than those (such as the Iceland and Loihi samples) that saw faster ascent rates. Initial calculations for EPR provide a good fit to the noble gas concentration data but do not match measured 3 He/ 22 Ne ratios. The model calculations produce higher ratios than seen in some of the sample suite data. This is likely due to the assumption that fully closed system conditions are present during 118

119 o stage 2 (magma ascent) degassing; allowing for partial open system conditions allows model calculations to match results for 4 He/ 40 Ar, but a lower value for the 3 He/ 22 Ne ratio of the MORB source mantle is still required. A value lower than 4.4 (down from the estimate of 11 given by Graham [2002]) would provide a better fit to the data, which compares well with the source estimate for OIBs of 3.6 [D W Graham, 2002]. The calculations for Iceland and Loihi can provide a good fit to all the data. However the simplest way of doing so requires initial He concentrations similar to or lower than the MORB initial helium concentrations. If the OIB source mantle contains a significant recycled component, lower He and higher major volatile concentrations are not surprising. However this scenario will either leave the helium paradox standing, or require a model in which a high 3 He/ 4 He component (along with neon) is added to a helium-poor protolith [C. J. Ballentine and Holland, 2008]. This study shows that understanding the elemental fractionation and decrease in concentration of the noble gases that takes place during the degassing processes must be taken into account when using noble gas data to speak to the nature of the mantle. Our three stage model shows how different degassing stages impact noble gas ratios and concentrations and that full account of the degassing history of a sample must be taken when inferring links between these values. Understanding a sample s degassing history allows limits to be placed on the initial noble gas composition of the source melt and mantle, and has here shown that such limits can provide valuable constraints to mantle models. 119

120 4.9 Appendix A: Detailed model derivation. The starting point for our degassing model is that of Gonnermann and Mukhopadhyay s [2007] disequilibrium model which is derived as follows. At the interface between melt and vapor phases, noble gas concentrations in the melt are described by Henry s law: c S px. (A1) i i i c i g -1 ) equilibrium concentration of noble gas species i in the melt phase (cm 3 STP S i solubility of noble gas i in melt (Henry s constant) (cm 3 STP g -1 bar -1 ) P X i total pressure (bar) mole fraction of noble gas i in the vapor phase under equilibrium conditions Given the low noble gas concentrations, it is reasonable to assume the bulk of the vapor phase will consist of the major volatiles, H 2 O and CO 2. Thus the mole fraction (X i ) of noble gas i can be approximated as X i vi v v c w. (A2) v i vapor phase concentration of noble gas i v c (v w ) equilibrium vapor phase concentration of CO 2 (H 2 O). These are determined according to the solubility model detailed in Dixon [1997]. The required inputs are melt concentrations of CO 2 and H 2 O before degassing, temperature and pressure. After degassing under equilibrium conditions and for a closed system, the average melt concentration for noble gas i would be c i, and the vapor phase composition (v i ) would be equal to c 0- c i (where c 0 is the initial concentration of i in the melt). However, under disequilibrium conditions, the average melt concentration of i will be higher than that at the interface as the noble gases take time to diffuse towards the vapor phase. This can be expressed as the fractional amount of the equilibrium 120

121 vapor phase that is retained in the melt due to disequilibrium conditions: fraction retained in melt c c. (A3) i 0 i An expression for i is determined for each volatile species assuming diffusion in the layer of melt surrounding a bubble can be approximated as diffusion from a plane sheet [Crank]: 8 1 D i exp 2 j 1 t Ar 2 deg as D j j i. (A4) D i diffusivity of volatile species i t degas is the parameter which determines the extent to which disequilibrium conditions are present and is given by t degas t. (A5) Ar t total available degassing time Ar characteristic diffusion time of Ar Ar is given by r Ar bubble radius vesicularity 2 a a r DAr 1 3 with 1. (A6) The average noble gas concentration in the melt ( c i ) is then given by i i i 0 i c c c c. (A7) Major volatiles are also affected by disequilibrium conditions and so final vapor phase concentrations for H 2 O and CO 2 must be modified by a disequilibrium factor 121

122 as follows: 1 and w 1 vc c vc v w vw. (A8) v c ( v w ) vapor phase concentration of CO 2 (H 2 O) c and w are determined from (A4) using the appropriate values for the diffusivity of CO 2 and H 2 O. The noble gas vapor phase composition (v i ) is c 0 - c i and so Henry s law becomes c i Si p c c vc vw 0 i. (A9) Combining (A7) and (A8) gives the result used in this model to determine final average noble gas concentrations in the melt phase as c i 2 1 i psi c0 i vc vw psi 1 i. (A10) 4.10 Appendix B: Method for calculating the number of steps for stage 2 and 3 of the model. The number of steps for stage 2 and stage 3 can be set using the following considerations: In our model variation, the number of degassing steps controls the timescale of bubble loss. For stage 2, this will depend on melt ascent rate, melt viscosity, bubble size and vesicularity. 122

123 In a steady state, balancing the viscous drag on the bubble, the bubble s buoyancy and the bubble s weight, the speed of the bubble relative to the ascending melt is given by v 2 2 gr. (B1) 9 b v m v b m speed of ascending bubble relative to ascending melt density of melt phase. For a basaltic melt this is set to 2600 kgm -3 [Paonita and Martelli, 2007] g acceleration due to gravity (9.81 ms -2 ) r bubble radius; a reference value for this parameter is found for each sample suite from (B1) using 10 deg as log t 0 and a pressure equal to the eruption pressure. melt viscosity set to 50 Pa s v density of vapor phase. Assuming the vapor phase is dominated by CO 2 this is given by Mp v. (B2) 1000RT M molecular mass of CO 2 p pressure (bar) R gas constant (8.314 x 10-5 m 3 bark -1 mol -1 ) T temperature (K) The melt in a single eruption can then be divided up into melt cells, each of which surrounds a single bubble. Knowing the speed of a bubble relative to the ascent rate of the melt then constrains how long a bubble will take to move through one melt cell. This gives the timescale of bubble loss for each step as 123

124 timescale of bubble loss a, (B3) v b Where v b is determined from (B1) and a (the size of a single melt cell ) is given by 1 a 2r 2r 3 1. (B4) vesicularity; set to a reference value of 0.1 [J E Dixon and Clague, 2001] During stage 2 (magma ascent), the total available time will depend on the melt s ascent rate and the distance it travels: distance of travel total available time. (B5) ascent rate distance of travel depth of first nucleation event (determined using Dixon s [1997] solubility model) minus eruption depth, taking into account a supersaturation factor of 4. (B3) depends on pressure (through v b ) which is decreasing throughout the course of the eruption. The timescale of bubble loss must be individually calculated for each step until the total time elapsed equals the total available time calculated in (B5): this will then give the number of steps for stage 2 (magma ascent). This parameter is calculated separately for each sample suite as ascent rates and pressure differ across the three sets of samples. For stage 3, the number of degassing steps can be determined using the same method, the only difference being that the total available degassing time is now determined by the quenching time of the sample in question. Paonita and Martelli [2007] suggest quench times ranging from minutes to hours, which equates to between 1 and 100 steps. A representative value of 10 steps is used as a default and reference value. 124

125 5 Chapter Five Noble gases and volatile recycling in Iceland's source mantle 5.1 Abstract We present noble gas data from an unusually volatile rich sample from SW Iceland. Iceland combines hotspot volcanism, a spreading ridge and abundant subglacially erupted basaltic samples. This combination allows for samples which erupted under high enough pressures to retain a measurable noble gas content, and also display signatures representing mixing between ocean island and mid-ocean ridge basalt mantle sources. In terms of the isotopic composition of the light noble gases, this interaction has been the subject of a number of studies. However, the elemental heavy noble gas composition of Icelandic basalts has been less well investigated. Noble gas studies are hampered by the large, isotopically atmospheric component typically found in Icelandic subglacial samples, which can swamp other signatures. In addition, the degassing process results in both elemental fractionation and loss of the noble gases. Taking full account of both these processes is crucial to resolving the elemental noble gas composition of Iceland s source mantle: Evidence for volatile recycling, volatile sources during the Earth s history and the nature of different mantle source zones are just a few topics that require elemental data as well as isotopic. Detailed analysis of a volatile rich sample from SW Iceland shows evidence for more than one contaminant air-like component. Component resolution based on two-component, mantle-air mixing models can produce misleadingly precise source mantle noble gas ratios. Multi-component best fits to noble gas elemental ratios find that four components are present in samples from this region. These components are unfractionated air, fractionated air and a mantle component which shows some variation due to degassing. Combining a disequilibrium degassing model with component resolution allows limits to be placed on the source mantle composition for this sample. The light noble gas source isotopic composition is compatible with mixing between a direct nebula component 125

126 and a MORB-like component. This direct nebula signature is at odds with an implanted signature seen in both Ne and Kr for the convecting mantle, and shows that both accretionary volatile origins must have contributed during the earth s formation. The heavy noble gases show an elemental abundance pattern that is distinct from air and solar patterns, and trends towards seawater. This confirms the presence of a recycled volatile signature in Iceland s mantle but it is not possible to further constrain the origin of this signature. 5.2 Introduction Basalts from ocean-island settings such as Iceland have shown noble gas elemental and isotopic ratios that differ from the well-studied mid-ocean ridge composition [C. J. Ballentine et al., 2001; E T Dixon et al., 2000; D. Harrison et al., 1999; M. Trieloff et al., 2002; Mario Trieloff et al., 2000; Valbracht et al., 1997]. Compared to mid-ocean ridge basalt (MORB) values, 3 He/ 4 He ratios are often high and this has contributed to the suggestion that the ocean-island basalt (OIB) source mantle represents a primitive, undegassed component from the deep mantle [C. J. Allegre et al., 1983; Macpherson et al., 1998]. The nature of such a reservoir and the possible extent of its interaction with both the depleted mantle and other reservoirs such as the atmosphere and the core is still a matter of debate [Albarede, 2008; Anderson, 1998b; C. J. Ballentine et al., 2002; P. E. van Keken et al., 2001]. Tomographic and seismic evidence for whole mantle convection has shown that the barrier for such a reservoir cannot be the 660 km phase boundary [Helffrich and Wood, 2001; R. D. Van Der Hilst et al., 1997]. Models exist which allow this division of the mantle with limited interaction although the mechanism for retaining such a division is unclear [Gonnermann and Mukhopadhyay, 2009]. Models of recycling at subduction zones have shown that recycled components from downgoing slabs are likely to be concentrated in the deep mantle and play a significant role in plume genesis [Brandenburg and van Keken, 2007a]. Coupled with the discovery of recycled volatiles in the convecting mantle, this has given rise to the suggestion that a recycled noble gas signature may play a role in determining the deep mantle s composition (the OIB source mantle). This argument is strengthened by the low 40 Ar/ 36 Ar values and low 3 He concentrations seen in OIBs which could 126

127 be explained by a recycled component. [C. J. Ballentine et al., 2005b; Holland and Ballentine, 2006]. Iceland s geological setting offers a unique testing ground for noble gas models of the mantle and mantle processes. Its setting on a spreading ridge and a probable plume mean both MORB and OIB mantle signatures can be expected to contribute to its noble gas collection. This is evident in the range of 3 He/ 4 He ratios and 3 He/ 22 Ne ratios already investigated [E T Dixon, 2003; E T Dixon et al., 2000; Hilton et al., 1990; Mark D. Kurz et al., 1985]. The abundance of subglacially erupted material allows for eruption pressures high enough for samples to retain a measurable volatile content of mantle origin. Elemental noble gas ratios from Iceland s source mantle may allow recycled signatures to be resolved and can also shed light on the different reservoirs present within Iceland s source mantle. Air contamination is ubiquitous in basaltic samples and Iceland is no exception [Chris J. Ballentine and Barfod, 2000]. This causes particular difficulties in looking for a potential deep mantle recycled signature as this will appear isotopically similar to late-stage air contamination. Two-component air-mantle mixing models are often used to deduce the noble gas composition of Iceland s mantle: such an analysis of recent high quality data from a single sample in western Iceland has been used to suggest that Iceland s mantle shows a signature that must represent a mantle reservoir held in isolation from the MORB-source mantle [Mukhopadhyay, 2012]. However both this study and previous Icelandic studies have shown that a third fractionated air component is common in Icelandic samples [D. Harrison et al., 1999]. This compromises endmembers calculated from two-component mixing models and requires different fitting strategies [D. Harrison, 2003]. Studies of disequilibrium degassing models have also shown that the role of degassing in fractionating elemental ratios cannot be ignored even when the radiogenic noble gas ratios ( 4 He/ 21 Ne and 4 He/ 40 Ar) are close to mantle production ratios [Gonnermann and Mukhopadhyay, 2007] (see chapter four). In this study data are presented from a glassy, volatile-rich basaltic sample taken from a region of Iceland that is often considered representative of Iceland s source mantle [D. Harrison et al., 1999; Mukhopadhyay, 2012; Mario Trieloff et al., 2000]. A multi-component fitting technique [Moniot, 2009] is used to constrain the 127

128 number and nature of noble gas components observed in both this data and existing literature data from the same region. Taking account of these components and the possible effects of degassing allows constraints to be placed on the noble gas isotopic and elemental ratios for this region of Iceland s source mantle. 5.3 Geological background Iceland is an ocean island that it is situated on both on a spreading ridge and a hotspot; the two sources of volcanism combine to produce a heterogeneous geochemical signature which varies with location across Iceland. The mid-atlantic ridge runs as shown in Figure 25. Iceland s major volcanic zones lie along the Reykjanes peninsula, the eastern, western and northern volcanic zones (EVZ, WVZ and NVZ) and in central Iceland. Off-rift volcanism is also seen at the stratovolcano Oræfajökull and at the Snæfellsnes peninsula. Seismic inversions of the hotspot anomaly beneath Iceland suggests a hot (~200 o C hotter than surrounding mantle), narrow (~100 km) plume which reaches a depth of at least 400 km [Wolfe et al., 1997]. Seismic studies have also shown the transition zone to be 20 km thinner than average under central and southern Iceland, which suggests a deep mantle origin for the Iceland plume [Helmberger et al., 1998; Shen et al., 1998]. The current centre of the plume is placed under the Vatnajökull icecap. This location shows some of the highest 3 He/ 4 He ratios [Mark D. Kurz et al., 1985] although low 3 He/ 4 He ratios are also observed for parts of central Iceland due to crustal contamination [Macpherson et al., 2005]. The presence of both mid-ocean ridge and hotspot volcanism allows their interaction to be studied. Iceland s source mantle must be highly heterogeneous given the range of observed geochemical signatures [Hards et al., 1995; O'Nions et al., 1976; Sigmarsson and Steinthórsson, 2007; Thirlwall, 1995; Thirlwall et al., 2006]. [Thirlwall et al., 2004] find evidence for at least four mantle components. Much of Iceland s geochemistry has been explained by the mixing of a depleted (DMM) component with more enriched or primordial plume components showing high 3 He/ 4 He and solar neon [P Burnard and Harrison, 2005; E T Dixon, 2003; Hart et al., 1973; Mark D. Kurz et al., 1985; Sun et al., 1975; Taylor et al., 1997; Unni and Schilling, 1978]. Various studies have also inferred the presence of a recycled component in Iceland s source mantle 128

129 [Chauvel and Hemond, 2000; Prestvik et al., 2001]. Low δ 18 O found in basalts from central Iceland and the Reykjanes peninsula could originate from hydrothermally altered, recycled oceanic crust [P Burnard and Harrison, 2005; Macpherson et al., 2005]. An increase in water content in samples going from the Reykjanes ridge towards central Iceland could indicate higher water content in the plume s mantle source due to a recycled component [Nichols et al., 2002]. Os and He isotopic systems in Iceland s neovolcanic zone suggest a complex mix of components in the Iceland plume including some primitive material along with ancient recycled crust [Brandon et al., 2007; Debaille et al., 2009]. Whether or not this recycled component is associated with primordial plume signatures such as higher 3 He/ 4 He ratios has implications for whole mantle convection models of the mantle, which expect to find recycled slabs concentrated in the deep mantle. Mixing scenarios may well be complicated by a plume that is itself likely to be heterogeneous [Hilton et al., 1990]. 129

130 Snæfellsnes Peninsula DICE Sample Area NRZ Central Iceland Rekyjanes Peninsula 3 He/ 4 He (R/R A ): WRZ ERZ Oræfajökull Figure 25: The shaded zones represent Iceland s major volcanic zones and spreading ridge. The variation in 3 He/ 4 He ratios, in multiples of R A, across Iceland ratios is also shown [E T Dixon, 2003; E T Dixon et al., 2000; Füri et al., 2010; Hilton et al., 1999; Mark D. Kurz et al., 1985] (and unpublished results from this work). Most of the highest ratios are seen towards central Iceland, near the centre of Iceland s proposed plume. Away from central Iceland, moving offshore along the rift zones, ratios trend towards lower MORB-like values. This has been interpreted as mixing between the MORB-source mantle with 3 He/ 4 He = 8R A and the plume-source mantle with a much higher 3 He/ 4 He. Lower than MORB 3 He/ 4 He ratios seen in SE Iceland (Oræfajökull) and the highest ratios seen in NW Iceland do not fit with this straight forward picture. 130

131 5.4 DICE sample area In this study, we present data from a volatile-rich glassy basaltic sample from Iceland s WVZ (sample MID1-64 o 10.5 N, 21 o W). We also discuss literature data from the same exposure from [Mario Trieloff et al., 2000] and [Mukhopadhyay, 2012]; both these data sets refer to the same sample (DICE 10). Trieloff s (2000) study also includes DICE 11, a second sample from the same area. Data from the original study of these samples are not used because of potentially compromised Ne results [D. Harrison, 2003; D. Harrison et al., 1999]. In terms of the noble gases, the best studied location in Iceland is that of the DICE10, DICE11 and MID1 (this study) samples. These samples are taken from picritic pillow mounds that are part of the Miðfell volcanoes in the Hengill system (called the DICE area hereafter). These picrites are thought to be the first products of a period of increased eruptive activity during early deglaciation caused by decreased lithostatic pressure [Tronnes, 1990]. These samples have been so well studied because they show some of the highest concentrations of mantle noble gases seen in Iceland to date. They therefore either represent a relatively undegassed melt, or a melt that was anomalously volatile-rich for Iceland. As the DICE samples are often used as a proxy for Iceland s entire mantle a full understanding of the components involved in the formation of these samples is essential. In this study, we look in detail at the components observed in these samples and their likely relationship to the Icelandic mantle. 5.5 Sample analysis Glass fragments between 0.5 and 2 mm were hand-picked using a binocular microscope to avoid visible alteration. Olivine and pyroxene samples were similarly hand-picked to choose the freshest samples available. Hand-picked samples were ultrasonically cleaned for 5 minutes in 2% HNO 3, de-ionised water and finally acetone to remove surface contamination. Samples were loaded into magnetic crushers and left to bake under vacuum conditions for 48 hours at 150 o C to reduce 131

132 atmospheric contamination from laboratory air. All analyses were carried out on a VG5400 mass spectrometer using a Daly multiplier detector for 3 He, 20,21,22 Ne, 36,38 Ar and all Kr and Xe isotopes. A Faraday cup was used to detect 4 He, 36 Ar and 40 Ar. After each set of samples was loaded, an initial blank analysis followed by a calibration was carried out to ensure blank noble gas levels were low and the spectrometer s sensitivity and tune settings were consistent. There was no measurable 3 He blank: Other typical blank levels were 1.8 x cm 3 STP 20 Ne, 2.2 x 10-9 cm 3 STP 40 Ar, 1.7 x cm 3 STP 84 Kr and 1.8 x cm 3 STP 132 Xe. An empty crusher was also used during some blank analyses to check that a significant amount of gas was not released from crusher walls during the crushing procedure. The first few crushes after sample loading produced levels of gas up to ten times higher than a static extraction line blank, presumably as vibrations and impact from the piston caused additional gas release from the extraction line walls. This is considered a worst case blank as during a real sample run the impact of the crusher on the crusher walls would be cushioned by the powdered sample. However, after this initial crush blank, gas levels released when using the empty crusher were usually not significantly higher than static extraction line blanks, although a large number of crushes (over ~ 1000) could produce blank levels up to three times higher than static extraction line blanks. To minimise inaccuracies due to active crusher blanks, the first crush of each sample was kept to a minimum number of strokes and an extraction line blank using the same number of strokes in the empty crusher was run before sample analyses and used to correct sample data. Typical blank levels were generally around 1% or less of sample releases, although this rose nearer to 10% for a few low release crushes. A calibration bottle containing air plus a helium spike with a 3 He/ 4 He of 14.6 ± 0.5 R A (where R A is the atmospheric ratio of x 10-6 ) was used to determine the spectrometer s sensitivity for each of the noble gases. Calibration and blank analyses exactly mimicked sample analyses (excepting the use of the crusher) and were run daily. For each sample, an electromagnet was used to operate a steel crusher for the required number of crushes. The released gases were then cleaned for 10 minutes on a SAES Zr-Al getter. The heavy noble gases were trapped on a charcoal trap held at 77 K by liquid nitrogen. Neon was separated from helium by means of a cryotrap held at 55 K. The light noble gases were analysed in turn. For most analyses, argon was separated from krypton and xenon by means of a cold charcoal trap held at the 163 K 132

133 (separation temperature was determined by calibration using a calibration shot). Early analyses measured all the heavy noble gases by separating argon from krypton and xenon but the separation process introduced larger errors for the Kr and Xe data and so later analyses ran Ar, Kr and Xe without separation. The analysis of each noble gas fraction was preceded by a further 5 minutes cleaning on a getter. During the helium and neon analysis, a charcoal trap held at 77 K by liquid nitrogen near the spectrometer s source was used to minimise interference from 40 Ar ++ and 44 CO The spectrometer s resolution was sufficient to obtain both a 40 Ar ++ and a 20 Ne + plus 40 Ar ++ measurement to allow 20 Ne + ++ to be calculated. The peaks for CO 2 and 22 Ne + were not resolvable and the amount of CO ++ 2 present for each run was found to depend both on both 20 Ne + and 4 He + counts. This interference was corrected for by calibrating shots of air noble gases combined with varying spikes of 4 He across a range of pressures and using a statistical best fit for the dependence of the 20 Ne/ 22 Ne ratio on both 20 Ne + and 4 He +. Levels of CO ++ 2 ranged from 1% of the total peak at mass 22 Ne to 50% for the lowest concentrations in some sample runs. These high levels of CO ++ 2 compared to 22 Ne + made this correction the largest source of error for 20 Ne/ 22 Ne ratios. 18 H 2 O + and CO + 2 were also monitored as a potential source of pressure dependence but background levels remained reasonably constant. 5.6 Results Results for individual crushes of the sample MID1 are displayed in Table 8. 3 He/ 4 He ratios across Iceland vary from MORB-like (~8 R A ) to much higher values found in central Iceland (see Figure 25). The MID1 analyses have an average 3 He/ 4 He of ± 0.08 R A, which is consistent with previous studies of the DICE area and towards the maximum value observed for the WVZ. A neon isotopic plot also shows good agreement with previous data from this sample area, following a trend that is close to an air-solar mixing line with much less radiogenic 21 Ne/ 22 Ne ratios than the air-morb mixing line (Figure 26). 40 Ar/ 36 Ar ratios also show mixing between air and a component with high 40 Ar/ 36 Ar. The highest measured 40 Ar/ 36 Ar value of 6451 is not far below that highest value of 7047 found to date in the DICE area [Mukhopadhyay, 2012]. Kr and Xe isotopic ratios are within error of 133

134 atmospheric values. Elemental abundances show a clear atmospheric component but do not show clear mantle/air mixing trends. This may indicate the presence of more than two components. SOLAR AIR MID1 this study DICE10 Mukhopadhyay, 2012 DICE10 Trieloff et al., 2000 DICE11 Trieloff et al., 2000 Figure 26: Three-isotope plot of neon data for samples from the DICE sample area. Each data point represents an individual crush of the specified sample. A linear trend, distinct from that observed for MORB samples, between air and a mantle component with higher 20 Ne/ 22 Ne and 21 Ne/ 22 Ne is clearly defined. 134

135 Table 8: Noble gas data for individual crushes of the MID1 sample from the DICE area. Concentrations are cm 3 STP g -1 and errors are 1σ. MID1 crushes 3 He X10-12 ± 22 Ne X10-12 ± 36 Ar X10-12 ± 84 Kr X10-12 ± 130 Xe X10-14 ± 3 He/ 4 He (R/R A) ± 20 Ne/ 22 Ne ± 21 Ne/ 22 Ne ± 40 Ar/ 36 Ar ± 129 Xe/ 130 Xe ± 136 Xe/ 130 Xe ± Sample TOTAL Sample TOTAL

136 MID1 He crushes X10-12 ± Ne X10-12 ± Ar X10-12 ± Kr X10-12 ± Xe X10-14 ± 3 He/ 4 He (R/R A) ± 20 Ne/ 22 Ne ± 21 Ne/ 22 Ne ± 40 Ar/ 36 Ar ± 129 Xe/ 130 Xe ± 136 Xe/ 130 Xe ± Sample TOTAL

137 5.7 Component analysis Noble gas data from Icelandic samples frequently show evidence for signatures derived from more than one noble gas reservoir [E T Dixon, 2003; E T Dixon et al., 2000; D. Harrison et al., 1999]. Identifying and quantifying these components is crucial if the mantle component of interest is to be resolved from the data. The most obvious component in the DICE area samples is unfractionated air added at a late stage, possibly during sample collection and/or processing [Chris J. Ballentine and Barfod, 2000]. This component is clearly seen in isotope mixing diagrams and often dominates the early crushing releases of a sample (see Table 8). Given the high 3 He/ 4 He ratios of the DICE samples it is also reasonable to expect at least one mantle component to be present. A simple two-component mixing scenario (mantle plus unfractionated air) would mean results should always lie on a straight line for three-isotope plots referenced to the same isotope. In this scenario, picking out the mantle component is straightforward as long as you have a single known mantle ratio (such as, for example, 4 He/ 22 Ne or 20 Ne/ 22 Ne) to use as a reference. A threeisotope neon plot fits this scenario with two clear components (see Figure 26). Although some of the noble gas ratios besides neon (such as 40 Ar/ 36 Ar) fit reasonably well with a two-component mixing scenario, considering all the available data makes it clear that two components are not adequate to explain the data (see Figure 27 and [D. Harrison et al., 1999]). 137

138 MID1 this study DICE10 Mukhopadhyay, 2012 DICE10 Trieloff et al., 2000 DICE11 Trieloff et al., 2000 AIR WATER Figure 27: 130 Xe/ 36 Ar vs. 40 Ar/ 36 Ar for the DICE sample area. A cluster of points with ratios close to unfractionated air indicate this as a clear component. A trend towards higher 40 Ar/ 36 Ar ratios (typical of a mantle component) correlates with higher 130 Xe/ 36 Ar ratios indicating a second component. However, a model containing only these two components does not provide a good fit to the data. A seawater-like noble gas component with high 130 Xe/ 36 Ar ratios but air-like 40 Ar/ 36 Ar ratios is a possible third component which would provide a better fit to these data [D. Harrison et al., 1999]. 138

139 Harrison et al., (1999; 2003) previously observed that many noble gas analyses, including those carried out on the DICE samples, show evidence of a third component that has some characteristics of a fractionated air component. All the data considered in this study show evidence for this component (for example Figure 27). To be sure of selecting a genuine mantle component, it is therefore necessary to combine multi-component fitting techniques with an understanding of the potential components in each sample. For an n-component mixture, each final noble gas ratio can be written in terms of its components: ( ) (5.1) I R I n /R n l n any noble gas isotope reference isotope is the noble gas ratio of the nth component the fraction of R n present in the nth component and (5.2) Equation 5.1 can be written for each noble gas isotope under study. As long as the reference isotope is kept the same, each noble gas ratio can be written in terms of (n-1) other ratios from equation 5.1 and equation 5.2. For example, in a threecomponent mixture 130 Xe/ 36 Ar could be written as: (5.3) where a, b and c are constants whose values depend on the noble gas ratios of the three components. In contrast a four-component mixture would need a further isotope to constrain the 139

140 130 Xe/ 36 Ar ratio: (5.4) It is therefore necessary to know how many components are present before creating statistical fits to the data and extrapolating to mantle endmembers. Both the light (Figure 28) and heavy (Figure 27) noble gas isotopes show signs of at least three components. One of these components is clearly unfractionated air (C1) and a second must be a mantle component (M). A third fractionated air component (as suggested by Harrison et al., 2003) would not show up as an extra component in the light noble gases as the 3 He/ 22 Ne ratios of air and water are very similar due to the low 3 He concentrations in the atmosphere. However the presence of a third component with high 3 He/ 22 Ne compared to air is obvious in Figure 28, particularly in the data set from Trieloff et al. (2000), and suggests that there are two separate mantle components (M1 and M2). To determine whether there is also a fractionated air component (C2) we have fitted two, three and four-component models respectively to the noble gas data for 36 Ar, 40 Ar, 22 Ne, 129 Xe, 130 Xe following the method of Moniot (2009). This method calculates the goodness of fit as χ 2 /ν, where ν is the number of degrees of freedom for the fit in question given by: m n x n y is the number of data points is one plus the number of independent coordinates is the number of dependent coordinates Values close to one represent a good fit. Values much greater than one indicate a poor fit or underestimated errors in the data. Values lower than one indicate that too many components have been introduced for the fit resulting in a spurious fit improvement or that errors are overestimated. Assuming errors are reasonable, the DICE data are not well fit by a two- or three-component model, but a reasonable fit 140

141 is given by a four-component model (see Table 9). This leaves two minor unknown components beside C1 (unfractionated air) and a mantle component. 20 Ne/ 22 Ne AIR MID1 this study DICE10 Mukhopadhyay, 2012 DICE10 Trieloff et al., 2000 DICE11 Trieloff et al., He/ 22 Ne Figure 28: The range of 3 He/ 22 Ne seen at high 20 Ne/ 22 Ne ratios indicates that a third component is required. The noble gas 3 He/ 22 Ne ratio of seawater is very similar to air and so seawater alone cannot provide the third component seen above, although it cannot be ruled out as an additional component. In the absence of cosmogenic production, the additional component with high 3 He/ 22 Ne is likely to be of mantle origin. 141

142 Table 9: Goodness of fit was calculated for using 130 Xe/ 36 Ar, 22 Ne/ 36 Ar, 40 Ar/ 36 Ar and 129 Xe/ 36 Ar as variables following the method set out in Moniot (2009). A good fit is indicated by (χ 2 /ν) 1. (χ 2 /ν)>1 indicates a poor fit or underestimated errors. (χ 2 /ν)<1 indicates overestimated errors or a spurious improvement to the fit due to adding an unnecessary component. A reasonable fit is provided by a fourcomponent model, whereas a two- or three-component model does not provide a good fit to the DICE area data. Number of components for fit Goodness of fit (χ 2 /ν)

143 5.8 Nature of additional components for the DICE area As fractionated air has previously been reported in many Icelandic samples, a potential additional component is elementally fractionated air with a signature similar to seawater or freshwater [D. Harrison, 2003]. This component would be isotopically identical to air and have similarly low He concentrations, fitting with the trends seen in the heavy elemental noble gas ratios (Figure 27). Unlike the late addition, unfractionated air component, this fractionated air component is seen in conjunction with mantle-like 40 Ar/ 36 Ar and 20 Ne/ 22 Ne ratios. This requires a different mechanism for the addition of this signature compared to the large unfractionated air signature which dominates early crushing steps. A possible mechanism would be the addition of this component through small amounts of mixing with the melt water surrounding the quenching basalt. This would explain the heterogeneous but ubiquitous appearance of this component but further work is needed to investigate the addition and siting of this component as other mechanisms such as air adsorption are possible. The final fourth component has high 3 He/ 22 Ne, 3 He/ 36 Ar, 3 He/ 84 Kr and 3 He/ 130 Xe ratios and so is likely to be of mantle origin. A possible origin for this component is variable degassing during sample eruption. As it is perfectly possible that different vesicles within a single sample could have experienced different degrees of degassing, this would produce a spread of elemental ratios between the most and least degassed vesicle for each sample. This would give the signal of an extra mantle component. Although the DICE sample is generally considered to be an undegassed sample due to 4 He/ 21 Ne* and 4 He/ 40 Ar* ratios close to mantle production values, it is possible for degassing to occur that will produce only small variations in these ratios. Correcting neon isotopes for atmospheric contamination and primordial contributions to 21 Ne (following the method in [D W Graham, 2002],) leaves a linear correlation between log( 3 He/ 22 Ne MANTLE ) and log( 4 He/ 21 Ne*) with a gradient of ~1 scattered around the mantle 4 He/ 21 Ne production ratio (see Figure 29). This is consistent with disequilibrium degassing models as isotopic fractionation during degassing is negligible and light to heavy elemental ratios can 143

144 be raised or lowered depending on the degree of disequilibrium [Gonnermann and Mukhopadhyay, 2007]. Using the model of Gonnermann & Mukhopadhyay (2007) as modified by Weston (this work), we calculate the possible spread of elemental noble gas ratios that could be produced by degassing given the observed range of 3 He/ 22 Ne MANTLE and 4 He/ 21 Ne*. The calculated spread can explain the apparent fourth component observed in the DICE data. A further constraint is provided as extensive degassing would significantly lower the 4 He/ 136 Xe* ratio and this is not observed (see Figure 30). This rules out large amounts of degassing. A small amount of degassing can explain the additional mantle component observed in elemental ratios for the DICE area. The four components in the DICE area samples are therefore unfractionated air (C1), fractionated air with an abundance pattern matching atmospheric noble gases in water (C2) and a mantle component split into lower (M1) and upper (M2) degassing bounds. An example of mixing between these components is displayed in Figure 31. Similar plots for other elemental ratios show that such a four-component mixing scenario can explain the observed data for the DICE sample area. 144

145 Mantle 4 He/ 21 Ne production ratio MID1 this study DICE10 Mukhopadhyay, 2012 DICE10 Trieloff et al., 2000 DICE11 Trieloff et al., 2000 Figure 29: Data for 3 He/ 22 Ne S ( 22 Ne corrected for atmospheric contamination) and 4 He/ 21 Ne* ( 21 Ne * indicates radiogenic 21 Ne) ratios from the DICE sample area. The range of ratios generated by degassing under equilibrium through to disequilibrium conditions for an initial composition matching the mantle 4 He/ 21 Ne* production ratio is log 10 ( 4 He/ 21 Ne*) = 4 to 10. The DICE data fall on a straight line with a gradient ~1, matching the trend expected for degassing. This shows that even though 4 He/ 21 Ne* ratios are close to the mantle production ratio, at least a small amount of degassing needs to be taken into account when looking at elemental ratios for these samples. From an error-weighted least-squares fit to the data, the mantle 3 He/ 22 Ne ratio is calculated to be 2.49±

146 MID1 this study DICE10 Mukhopadhyay, 2012 DICE10 Trieloff et al., 2000 DICE11 Trieloff et al., 2000 Increasing initial CO 2 Increasing initial CO 2 Increasing disequilibrium Figure 30: The grid represents the range of results generated by degassing from an initial composition matching the star, following the model of Gonnermann and Mukhopadhyay (2007) as modified by Weston (see chapter four). Conditions range from equilibrium to full disequilibrium and initial melt CO 2 concentrations range from wt.%. 4 He/ 136 Xe* ratios from the DICE sample area changed by less than an order of magnitude confirming that only a small amount of degassing has taken place. 146

147 MID1 this study DICE10 Mukhopadhyay, 2012 DICE10 Trieloff et al., 2000 DICE11 Trieloff et al., 2000 Highes observed 40 Ar/ 36 Ar ratio C1 C2 Figure 31: Mixing between three components unfractionated air (C1), freshwater (C2) and a degassed mantle component with upper and lower 130 Xe/ 36 Ar values represented by M1 and M2. This four component mixing model can provide a good fit to the data observed from crushes of samples from the DICE sample area. The exact values for M1 and M2 are unknown but show a trend towards higher than air 130 Xe/ 36 Ar ratios with increasing 40 Ar/ 36 Ar ratios. 147

148 5.9 Source mantle composition Assuming a four-component mixture between air (C1), fractionated air (C2) and a small spread of degassed mantle ratios (M1 and M2), it is possible to place some constraints on the pre-degassing mantle component from mixing plots. However there are not enough constrained ratios to calculate exact endmember elemental ratios for all isotopes using a four-component fit. Using the components discussed above as limiting factors, it is possible to place upper and/or lower bounds on initial elemental and isotopic ratios as follows Ne/ 22 Ne C1 and C2 both have 20 Ne/ 22 Ne = 9.80 and M1 and M2 can be assumed to have the same mantle 20 Ne/ 22 Ne ratios as degassing does not significantly affect isotopic ratios. A neon isotopic plot should therefore show a linear two-component mixing trend and this is observed in the DICE data (see Figure 26). However this mixing trend can only give a minimum value (12.88±0.06) for the mantle component s 20 Ne/ 22 Ne ratio as any addition from C1 or C2 will lower the mantle ratio. There has been some debate as to whether the 20 Ne/ 22 Ne ratio of Iceland s mantle should match the Ne-B ratio of 12.5 ± 0.04 found in some MORB studies [Chris J. Ballentine et al., 2005a; Holland and Ballentine, 2006] or a solar wind ratio of ± 0.03 [D. Harrison et al., 1999; Heber et al., 2009; Mario Trieloff et al., 2000]. Ne-B is ruled out by ratios >12.5 found both by Mukhopadhyay (2012) and in this study. However the DICE data show a trend towards low 36 Ar/ 22 Ne ratios that correlates with increasing 20 Ne/ 22 Ne ratios. This trend also rules out the higher solar wind 20 Ne/ 22 Ne ratio of 13.8 as the DICE source mantle ratio (see Figure 32) Heavy isotopic ratios Assuming a four-component mixing model, the mantle s 40 Ar/ 36 Ar, 129 Xe/ 130 Xe and 136 Xe/ 130 Xe ratios can only be given limits as the contributions from air (C1) and fractionated air (C2) cannot be quantified. A two-component hyperbolic fit to the 148

149 DICE area 40 Ar/ 36 Ar vs. 20 Ne/ 22 Ne data, with unfractionated air as the first component and a mantle component with 20 Ne/ 22 Ne=13.78, gives a best fit for a mantle 40 Ar/ 36 Ar ratio of (Figure 33a). This is much lower than estimates for the MORB-source mantle (e.g. >40,000 from popping rock [P Burnard et al., 1997]), in agreement with the 40 Ar/ 36 Ar ratio of 10,745 calculated for DICE 10 by Mukhopadhyay (2012), but this model does not provide a good fit to the data and ignores the evidence for additional components discussed above. Contrasting a two -component model with a three-component fit, using the noble gas composition of seawater as a third component, shows that a three-component (or greater) model with a low 40 Ar/ 36 Ar mantle ratio can provide a good fit to the data. However, the three-component model can also provide a good fit to the data assuming a mantle 40 Ar/ 36 Ar ratio as high as current estimates for the MORB source mantle (Figure 33b). This ratio can therefore only be constrained as lying above the maximum 40 Ar/ 36 Ar value observed to date of 7047 ± 70 [Mukhopadhyay, 2012]. The source value is likely to be somewhat higher than 7047 given that the fractionated air component (C2) is still seen in high 20 Ne/ 22 Ne crushes. This is also the case for 129 Xe/ 130 Xe and 136 Xe/ 130 Xe ratios: both MORB-like values (up to 7.6 and 2.64 respectively - [Mukhopadhyay, 2012]) and lower values are possible mantle endmembers for the DICE area s source mantle (Figure 34). These ratios can only be given a minimum value at the maximum observed values of 7.01±0.08 and 2.33±0.14 respectively He/ 22 Ne Assuming a mantle 20 Ne/ 22 Ne of 13.5±0.1 ( Discussion section) and 4 He/ 21 Ne* ratio of 2.22x10 7, the mantle 3 He/ 22 Ne ratio can be determined from consideration of degassing trends. First neon is corrected for atmospheric contamination using a two-component mixing model (two-component mixing is valid here as the isotopic ratios for C1/C2 and M1/M2 are identical) [D W Graham, 2002]. As discussed above, log 10 ( 4 He/ 21 Ne*) vs. log 10 ( 3 He/ 22 Ne MANTLE ) fits a straight line with a gradient ~1, as would be expected from a small amount of degassing (Figure 29). Assuming an initial 4 He/ 21 Ne* equal to the mantle production ratio, this fit then allows the initial mantle 3 He/ 22 Ne to be calculated as 2.49±

150 MID1 this study DICE10 Mukhopadhyay, 2012 DICE10 Trieloff et al., 2000 DICE11 Trieloff et al., 2000 AIR (C1) WATER (C2) Figure 32: Data from the DICE sample area. The solid line represents a mixing trend between air and a component with 20 Ne/ 22 Ne=13.8 (direct nebula) and 36 Ar/ 22 Ne=0. This mixing trend is in the rough direction of the air-mantle mixing trend but cannot explain the data that fall to the left of this line. A lower 36 Ar/ 22 Ne for the mantle component is not possible so a lower 20 Ne/ 22 Ne mantle ratio is required to fit this low 36 Ar/ 22 Ne data. This implies that the DICE source mantle s 20 Ne/ 22 Ne ratio cannot be as high as the solar wind ratio of However a measured value of 20 Ne/ 22 Ne=12.88 (±0.06) implies that the source value must also be higher than the value of 12.5 (Ne-B) proposed for the MORB-source value. An intermediate value could be the product of MORB-OIB mixing at Iceland. 150

151 a b MID1 this study DICE10 Mukhopadhyay, 2012 DICE10 Trieloff et al., 2000 DICE11 Trieloff et al., 2000 Figure 33: Potential mixing models compared to 40 Ar/ 36 Ar and 20 Ne/ 22 Ne data from the DICE sample area. The single central blue line in figure 33a represents two component mixing between air and a mantle component with 20 Ne/ 22 Ne =13.8 and 40 Ar/ 36 Ar = The area between the two outer red curves on the same figure represents the possible range of ratios for a three component mixing model between air, elementally fractionated air as found in water and a mantle component with MORB-like 40 Ar/ 36 Ar (40,000). This model provides a better fit to the data than a two component mixing model and shows that Iceland s 40 Ar/ 36 Ar ratio is not necessarily lower than the MORB-source mantle. Figure 33b shows a three component mixing model using the same components as for 33a except that the 40 Ar/ 36 Ar ratio of the mantle component is lower than MORB values at 12,000. This shows that with three (or more) component mixing both MORB-like and lower 40 Ar/ 36 Ar ratios are possible endmembers. 151

152 Figure 34: Three component mixing models for 129 Xe/ 130 Xe vs. 20 Ne/ 22 Ne for Mukhopadhyay s (2012) data. The area between the blue curves represents possible ratios for three component mixing between an unfractionated air component, a fractionated air component akin to water and a mantle component with a low 129 Xe/ 130 Xe compared to the MORB source mantle. The area between the red curves represents possible ratios for three component mixing between an unfractionated air component, a fractionated air component akin to water and a mantle component with a 129 Xe/ 130 Xe equivalent to that proposed for the MORB source mantle. Both models provide an equally good fit to the data, showing that Iceland s source mantle is not necessarily distinct from the MORB-source mantle in terms of 129 Xe/ 130 Xe ratios. 152

153 Figure 35: A range for the degassed 36 Ar/ 22 Ne ratio for the mantle (M1/M2) component is represented by the solid blue shape on the degassing grid. The limits on possible values for initial CO 2 concentration and degree of disequilibrium are placed by consideration of the range of observed 4 He/ 21 Ne values compared to the known mantle production ratio. 153

154 5.9.4 Elemental ratios 36 Ar/ 22 Ne ratios show a negative correlation with 20 Ne/ 22 Ne ratios (Figure 32). This allows the lowest measured 36 Ar/ 22 Ne ratio to provide a maximum value for the lower degassed mantle component (M1), since atmospheric contaminants and M2 can only increase this ratio. The actual source mantle value must lie between this ratio and the maximum possible degassed ratio. Given the constraints on degassing determined from 3 He/ 22 Ne and 4 He/ 21 Ne ratios, the potential changes in elemental ratios due to degassing can be calculated (see Figure 35). A minimum value for the source mantle 36 Ar/ 22 Ne can then be calculated from the minimum value for M1. 84 Kr/ 22 Ne and 130 Xe/ 22 Ne also show an inverse correlation with 20 Ne/ 22 Ne ratios, allowing limits to be calculated for these elemental ratios using the same technique as for 36 Ar/ 22 Ne. The resulting mantle ratio ranges are shown in Table Kr/ 36 Ar and 130 Xe/ 36 Ar show a positive correlation with 40 Ar/ 36 Ar ratios, indicating the mantle endmember has higher than air 84 Kr/ 36 Ar and 130 Xe/ 36 Ar ratios. Several data points have higher 84 Kr/ 36 Ar and 130 Xe/ 36 Ar ratios than even the fractionated air component (C2) (see Figure 27). These points can be used to place a lower limit on the M2 component, since all other components can only lower this ratio. A line between the C2 component and the point with highest 84 Kr/ 36 Ar or 130 Xe/ 36 Ar defines a line which tracks the lower limit of the M2 component for a given 40 Ar/ 36 Ar ratio. This lower limit has a minimum value at the minimum 40 Ar/ 36 Ar value of 7050 ± 70. As for the 36 Ar/ 22 Ne ratio, the change in an initial mantle ratio due to degassing can be calculated using the disequilibrium degassing model (Chapter 4) from the range of observed 4 He/ 21 Ne values, allowing the source mantle ratio to be calculated from the lower limit for M2. This technique gives a minimum 130 Xe/ 36 Ar ratio of (2.44 ± 0.03) x 10-4 and a minimum 84 Kr/ 36 Ar ratio of ± for the DICE area source mantle. A similar technique can place a maximum limit on the source mantle 130 Xe/ 36 Ar ratio (Figure 31). A best fit line through the MID1 data (used as it shows the least influence from the C2 component) defines a line that tracks the maximum 130 Xe/ 36 Ar ratio of the M1 component for a given 40 Ar/ 36 Ar. Assuming that the 40 Ar/ 36 Ar ratio of the DICE area source mantle is not higher than the estimate of 40,000 for the MORB source mantle, a maximum 130 Xe/ 36 Ar ratio for M1 is given at 40 Ar/ 36 Ar = 40,000 (Table 10). Possible 154

155 degassing ranges then give a maximum 130 Xe/ 36 Ar of (8.5 ± 1.1) x 10-4 for the DICE area source mantle. The errors on the Kr data are too large to allow a meaningful maximum limit to be calculated for the 84 Kr/ 36 Ar source mantle ratio. Calculated mantle ranges for all the ratios discussed are summarised in Table 10. Elemental ratios are compared to other key elemental noble gas trends in Figure 36. In general the light noble gases show elemental ratios that tend towards solar values. They are also lower than the ratios proposed for MORB from popping rock and Bravo Dome data [Holland and Ballentine, 2006; Raquin et al., 2008]. The heavy noble gas elemental ratios (Xe/Ar and Kr/Ar) are higher than air and similar to values suggested for the MORB source mantle. These values are similar to the fractionated air pattern seen in seawater (Figure 36). Table 10: Calculated ranges for the source mantle of the DICE area. Isotopic Ratio DICE area range for source mantle Elemental Ratio DICE area range for source mantle 3 He/ 4 He ± 0.08 R A 3 He/ 22 Ne 2.49 ± Ne/ 22 Ne 13.5 ± Ar/ 22 Ne < 6.1 ± Ne/ 22 Ne ± Kr/ 22 Ne < ± Ar/ 36 Ar >7050 ± Xe/ 22 Ne < (10.5 ± 0.4) x Xe/ 130 Xe >7.01 ± Kr/ 36 Ar >0.061 ± Xe/ 130 Xe >2.33 ± Xe/ 36 Ar > (2.44 ± 0.03) to (8.5 ± 1.1) x

156 5.10 Discussion He and Ne isotopes across Iceland have been interpreted as a mixture of the MORB-like component seen off Iceland along the Reykjanes ridge and a high 3 He/ 4 He component associated with Iceland s plume [E T Dixon, 2003; Mark D. Kurz et al., 1985]. More generally the geochemistry of Iceland s basalts has been explained by a heterogeneous mantle consisting of three major components: a depleted mantle component; a plume component associated with high 3 He/ 4 He ratios (up to 30 R A ) and solar wind-like neon ( 20 Ne/ 22 Ne = 13.8); and a recycled crustal component that could come from either the deep or shallow mantle. The DICE sample area has a 3 He/ 4 He ratio of ~18R A, a ratio higher than the average MORB value of 8R A, but lower than the highest values associated with plumes of up to 50 R A [Starkey et al., 2009]. Such an intermediate value is indicative of a melt that is a mixture of MORB and plume components. The neon isotopic values for the DICE area match this mixing picture. The MORB source mantle has a 20 Ne/ 22 Ne ratio of 12.5 ± 0.04, equivalent to the Ne-B implanted component [Chris J. Ballentine et al., 2005a; Holland and Ballentine, 2006]. The 20 Ne/ 22 Ne ratio of the DICE area s source mantle of 13.5 ± 0.1 lies above this Ne-B value but below a direct nebula ratio of This intermediate 20 Ne/ 22 Ne ratio can be explained by mixing between a MORB-source reservoir with 20 Ne/ 22 Ne = 12.5 ± 0.04 and 21 Ne/ 22 Ne = ± [Chris J. Ballentine et al., 2005a; Holland and Ballentine, 2006] and OIB mantle reservoir with a solar composition of 20 Ne/ 22 Ne =13.78 ± 0.03 and 21 Ne/ 22 Ne = ± [Heber et al., 2009]. This assumption allows the source 20 Ne/ 22 Ne ratio to be calculated from the intercept of the observed DICE air-mantle mixing line for a neon isotopic plot with a MORBsolar mixing trend (Figure 37). The resulting intermediate ratio is 13.5 ± 0.1, equivalent to 82 (± 8) % of the DICE area s source mantle s 22 Ne coming from a source with 20 Ne/ 22 Ne = 13.8 and the remainder from MORB neon. Similarly, a composition of 60 (± 0.8) % 3 He from a plume component with 3 He/ 4 He = 120R A, and 40% 3 He from a MORB component with 3 He/ 4 He =7.75 ± 0.86 R A would give the 3 He/ 4 He =17.38 ± 0.08 R A observed in DICE samples and is consistent with the plume component having a solar-like 3 He/ 22 Ne ratio. If the plume component of Iceland s mantle represents a solar component, at least as far as the light noble gases are concerned, this could also explain a notable feature of the DICE area: the 156

157 high ratios of neon to the heavy noble gases when compared to both air and typical MORB values (Figure 36). Seawater Bravo Dome Popping Rock Solar Limiting values for DICE source mantle Figure 36: Key non-radiogenic ( 3 He, 22 Ne, 36 Ar, 84 Kr and 130 Xe) elemental noble gas trends are compared with the range of ratios calculated for the DICE area s source melt. Stars for He, Ne and Ar represent minimum source mantle values. Maximum values could not be determined. Stars for Xe represent the maximum and minimum source mantle values. The 130 Xe/ 36 Ar for the DICE area samples show a range close to a seawater composition as observed in the MORB source mantle [Holland and Ballentine, 2006]. The minimum value for 84 Kr/ 36 Ar is also consistent with a seawater composition. The light noble gases show ratios trending away from air and seawater towards solar values [Holland and Ballentine, 2006]. 157

158 Figure 37: Mixing between a solar neon component, a MORB-like neon component and atmospheric neon can produce the intermediate 20 Ne/ 22 Ne and 21 Ne/ 22 Ne ratios observed in the DICE sample area. In contrast to the dominantly solar signature seen in He and Ne, the heavy noble gases show an abundance pattern that tends towards a seawater signature (Figure 36) [Holland and Ballentine, 2006]. The presence of a recycled heavy noble gas signature within the MORB source mantle has shown that significant proportions of noble gases are recycled into the mantle at subduction zones [Holland and Ballentine, 2006]. The distinct trend towards seawater values seen in the heavy noble gases in the DICE area s source mantle confirms the presence of a recycled heavy noble gas signature within the mantle. Physical models of mantle convection have shown that the mantle convects as a whole and the down-going slabs may end up concentrated in the deep mantle, near the core-mantle boundary (CMB) [Brandenburg et al., 2008], which is also the proposed origin of Iceland s mantle 158

159 plumes [Helmberger et al., 1998]. Such a slab component would be depleted in He and Ne, be progressively enriched in the heavier noble gases relative to the light and should look isotopically similar to air. If this recycled component then combined with a deep mantle solar component, the resulting plume signature would be dominantly solar for He and Ne, but dominantly recycled for the heavy noble gases. This matches the elemental abundance patterns seen in the DICE area. However, as discussed above, the DICE area source mantle also requires a small component from the MORB source mantle to explain the observed 3 He/ 4 He and 20 Ne/ 22 Ne ratios. This MORB component is known to carry a recycled abundance pattern in the heavy noble gases and it is not possible to distinguish whether the recycled signature seen in the DICE source mantle is due solely to this MORB component or whether it also resides in the OIB component. A further consideration is the origin of this MORB component. The combination of a spreading ridge and a hotspot at Iceland means it is possible for the MORB component to be added by shallow melt mixing alone. However, similar MORB components such as lower than solar wind 3 He/ 4 He and 20 Ne/ 22 Ne, and higher 36 Ar/ 22 Ne ratios are seen at other hotspots such as Loihi [Mario Trieloff et al., 2000], implying that MORB/plume mixing is not unique to Iceland, although a small degree of shallow melt mixing is still necessary to explain the large range of He and Ne isotopic ratios seen across Iceland. This in turn suggests that the MORB source mantle is a common component of mantle plumes, which raises the question of the location and scale of MORB and OIB source mantle mixing. Xenon isotopic ratios from samples in the DICE area have been used to argue that the heavy (as well as the light) noble gases in OIBs are distinct from the MORB source mantle and preclude deep mantle MORB/OIB mixing [Mukhopadhyay, 2012]. In particular distinct 129 Xe/ 130 Xe ratios for the two mantle sources mean that these two reservoirs must have remained separate since the extinction of 129 Xe s parent isotope, 129 I. This assumption is based on xenon isotope ratios derived from two component mixing models for the DICE data. However, models involving more than two components (as applied here) mean that the noble gas elemental ratios of the combined contaminant components (C1 and C2) are not constant from sample to sample, or even between separate crushes. It is therefore not possible to fix the 159

160 40 Ar/ 36 Ar, 129 Xe/ 130 Xe or similar mantle isotopic ratios from hyperbolic fits with the 20 Ne/ 22 Ne data [Mukhopadhyay, 2012], and ratios equivalent to those proposed for the MORB source mantle are still a possible endmember for the OIB source mantle (see Source mantle composition section). The implication is that MORB and OIB source mantle Xe and Ar isotopic ratios are not necessarily distinct, although Mukhopadhyay s (2012) conclusion that 3 He/ 130 Xe ratios differ between these two reservoirs is still valid (Figure 38). This allows at least two possibilities for deep mantle OIB/MORB reservoir mixing which are summarised in Figure 38: 1) The DICE area source mantle could have heavy isotopic ratios indistinguishable from the MORB composition, but a light noble gas composition that consists of a solar component mixed with the MORB component. The OIB mantle reservoir would then be identical to the convecting mantle combined with a solar addition. A problem with this picture is that it is at odds with source He concentrations for Iceland, which are lower than MORB concentrations (see chapter four). 2) Including a recycled slab component in the OIB source would explain lower than MORB He concentrations, as this recycled component would be depleted in He. This recycled component would move heavy noble gas isotopic ratios towards airlike values and the heavy noble gas abundance pattern towards seawater. Mixing, at some unknown point between plume generation and eruption, with a MORB source would then generate the ratios observed at Iceland. The exact location in the mantle of the high 3 He/ 4 He component needed to generate the observations at ocean islands is debated, although the wide acceptance of whole mantle convection places it in the deep mantle [Peter E. van Keken and Ballentine, 1998]. The data from the DICE area suggests that this deep mantle component has a direct nebula origin. This is in contrast to the signature observed in the convecting mantle, which shows an implanted accretionary signature for both Kr and Ne [Chris J. Ballentine et al., 2005a; G. Holland et al., 2009b]. This indicates that both implanted meteoritic volatiles and volatiles captured directly from the solar nebula may have contributed to the Earth s volatile composition, since both signatures are seen in different mantle reservoirs. It is a requirement of some mantle models, such as steady-state models, that He comes from a common source for the plume and MORB mantle reservoirs [O'Nions and Oxburgh, 1983; Porcelli and Wasserburg, 1995]. A direct nebula noble gas signature in the OIB 160

161 source mantle, contrasting with the implanted noble gas signature seen in the MORB source mantle calls into question whether this is possible. We have shown that the mixing of melt with at least two sources of atmospheric contamination makes it difficult to resolve mantle 129 Xe/ 130 Xe, 136 Xe/ 130 Xe and 40 Ar/ 36 Ar ratios sufficiently precisely to distinguish the OIB source from the MORB source for these ratios, allowing the possibility of a common origin for the 3 He content of the MORB and OIB source. It is also possible that the heavy noble gases in Iceland s plume show the same composition as the MORB source mantle, including an implanted Kr and Xe signature that has yet to be resolved from the contaminants ubiquitously affecting Icelandic data. However the neon picture of the two mantle reservoirs poses more of a problem than the heavy noble gases for steady state mantle theories. If the MORB source s He comes from the direct nebula OIB source, this signature should be observed in MORB neon as well. One possibility is that the MORB 20 Ne/ 22 Ne ratio of 12.5 ± 0.04 for the convecting mantle is the combined mixing product of direct nebula solar neon, implanted solar Ne and a small atmospheric component introduced through recycling. Although recycled slabs do not contain much Ne due to its low solubility in seawater, they will still contain a small Ne component isotopically identical to air which will, over time, reduce accretionary 20 Ne/ 22 Ne ratios. This could explain the distinct 20 Ne/ 22 Ne ratios of the MORB source mantle but still allow the deep plume reservoir to be the source of the convecting mantle s 3 He, but further modelling work and high resolution OIB heavy noble gas data are required. 161

162 Figure 38: 3 He/ 130 Xe vs. 129 Xe/ 130 Xe data from the DICE area. A two component mixing model produces an endmember for Iceland s source mantle that has a distinct 129 Xe/ 130 Xe ratio from the MORB-source mantle. This requires these two reservoirs to have been separated since 129 Xe s parent isotope became extinct. However, a three (or greater) component mixing model allows other possibilities. For example, an OIB component with an air-like 129 Xe/ 130 Xe ratio but a high (solarorigin) 3 He/ 130 Xe ratio could mix with the MORB-source mantle to produce an endmember compatible with the DICE data. This would fit well with mantle models that place a concentration of recycled slabs in the plume-source mantle. Alternatively an additional 3 He source combined with a mantle reservoir with MORB-like 129 Xe/ 130 Xe ratios could produce a mantle endmember compatible with the DICE area data. 162

163 5.11 Conclusions o It is possible to constrain the number of noble gas components present in a sample by best fits to elemental ratios. Two-component fits used incorrectly can produce misleadingly precise mantle noble gas ratios. For example, the mantle 129 Xe/ 130 Xe and 40 Ar/ 36 Ar endmember ratios for the DICE sample area in Iceland could be as high as the ratios proposed for the MORB-source mantle given a three-component (or greater) mixing scenario. o The DICE sample area shows evidence for four noble gas components. We interpret these to be unfractionated air (C1), fractionated air (water C2) and a mantle component that shows some variation due to a small amount of degassing (M1 and M2). The mantle component trend can be identified by higher than air 20 Ne/ 22 Ne and 40 Ar/ 36 Ar ratios. Constraints can then be placed on elemental and isotopic noble gas ratios of this component by using 4 He/ 21 Ne* and 4 He/ 136 Xe* ratios to constrain degassing and considering the possible ranges of 20 Ne/ 22 Ne and 40 Ar/ 36 Ar. o The glass samples from the DICE area show light isotopic and elemental noble gas ratios consistent with mixing between a direct nebula solar component (with 20 Ne/ 22 Ne=13.78 and 3 He/ 4 He=120 R A ) and a MORB-like component (with 20 Ne/ 22 Ne=12.5 and 3 He/ 4 He=7.7 R A ). The origin of the plume component s solar noble gases cannot be fixed, but mantle convection models do not leave much room for a mantle reservoir containing solar ratios: such a signature must have come from deep mantle layers such as the D or even the core. A negative correlation between 36 Ar/ 22 Ne and 20 Ne/ 22 Ne for samples from the DICE area implies a 20 Ne/ 22 Ne source ratio that is lower than the solar wind value of If this intermediate 20 Ne/ 22 Ne ratio is assumed to derive from mixing between the MORB-source mantle ( 20 Ne/ 22 Ne = 12.5) and OIBsource mantle that carries a solar signature for the light noble gases ( 20 Ne/ 22 Ne =13.78), this allows a value of 20 Ne/ 22 Ne =13.5 ± 0.1 to be calculated for the DICE area s source mantle. 163

164 o The heavy noble gas elemental abundance pattern for the DICE area s source mantle is distinct from air and solar patterns, and trends towards a seawater pattern. We interpret this result as further confirmation that a recycled noble gas signature dominates the mantle s heavy noble gases. For Iceland, this recycled component could be derived from a MORB component added either at a shallow level during melt mixing with melts produced by Iceland s spreading ridge, or at a deeper level within the mantle as the plume journeys towards the Earth s surface. However, the DICE data are consistent with some of this recycled component coexisting with the solar component representing Iceland s mantle plume, as would be expected given mantle models that predict the influence of recycled slabs in the deep mantle. This interpretation would also explain the low 3 He concentrations of the DICE source mantle compared to the MORB source mantle, as the recycled slab material would introduce an extra He-poor component. o The direct nebula signature seen for Iceland s plume component is at odds with evidence for an implanted solar origin seen in both Ne and Kr ratios for the MORB-source mantle. It is possible that noble gases with both a direct solar origin and an implanted origin are present in the Earth s mantle. Fourcomponent mixing shows that heavy noble gas isotope ratios cannot be sufficiently constrained to rule out mantle models where the deep mantle, high 3 He/ 4 He reservoir provides the convecting mantle s 3 He. However the difference in 20 Ne/ 22 Ne ratios between the MORB and OIB source mantles is more of a concern for these mantle models. It is possible that the MORB mantle s 20 Ne/ 22 Ne ratio does not in fact represent a purely implanted signature, but is a mixture of solar wind, implanted and recycled signatures. 164

165 5.12 Appendix to noble gas paper Introduction In this appendix noble gas results from samples across Iceland are presented and discussed. Away from the DICE area discussed above, samples generally showed low mantle 3 He concentrations and the dominant air contaminant component present in basaltic glasses made it difficult to resolve any mantle components from the dominant air signatures. However, many samples did show evidence for mantle 3 He and some noble gas results were distinct from an atmospheric composition in Ne and Ar. These results are presented and discussed here Noble Gas Results The data for these samples are displayed in Table 11. Sampling strategy and sample locations are described in chapter three, as are the experimental techniques used to determine the noble gas composition of these samples. 3 He/ 4 He ratios match previous studies across Iceland and are added to literature data to produce the helium map of Iceland shown in Figure He/ 4 He values vary from MORB-like values (~8 R A where R A is the atmospheric ratio of 1.399x10-6 ) away from the postulated plume centre, to much higher values towards central Iceland [Mark D. Kurz et al., 1985]. Values lower than 8R A are also seen in the Oraæfajökull- Snæfells volcanic system in SE Iceland [Peate et al., 2010]. The range of Icelandic 3 He/ 4 He ratios is discussed further in chapter six. 3 He concentrations are typically <10 x cm 3 STP g -1 and hence much lower than found in typical MORB samples, and in samples from the DICE sample area. The majority of samples showed Ne, Ar, Kr and Xe isotopic ratios indistinguishable from air. Only samples KALF9, JARL7, SN6, HLO100, KALF4, ARM1 and showed 20 Ne/ 22 Ne and 40 Ar/ 36 Ar ratios in excess of the atmospheric values. Samples HB2, HLO102, HLO103, HLO104, SN6, JARL8 and T1 showed atmospheric neon isotopic ratios but excesses in 40 Ar/ 36 Ar. Most of the samples analysed which showed higherthan-air 20 Ne/ 22 Ne ratios have a neon isotopic pattern compatible with the air- 165

166 mantle mixing line determined for the DICE area (see Figure 39). A few samples (from the WVZ and Snæfellsnes peninsula) show trends that lie between the MORB-air mixing line and the DICE mixing line, indicating a more pronounced MORB signature than seen in the DICE area. The Snæfellsnes peninsula sample also shows higher 40 Ar/ 36 Ar ratios for a given 20 Ne/ 22 Ne than would fit with the DICE trend, again closer to a MORB-air mixing trend. Sample ARM1 (from a location a few kms north of the DICE area) lies on the DICE trend line from unfractionated air to low 36 Ar/ 22 Ne. Other samples are compatible with the 36 Ar/ 22 Ne range determined for the DICE area, but since few show lower than air 36 Ar/ 22 Ne, mantle endmembers with higher (up to seawater-like) 36 Ar/ 22 Ne are also possible. 84 Kr/ 36 Ar and 130 Xe/ 36 Ar are invariably greater than or equal to air ratios, but possible mantle endmembers can only be constrained as being air-like or higher due to the dominance of atmospheric contamination. The presence of more than two mixing components in these samples is apparent in the 130 Xe/ 36 Ar vs. 40 Ar/ 36 Ar trend (Figure 40). 166

167 Figure 39: Three-isotope Ne plot for samples across Iceland. Many samples show ratios indistinguishable from atmospheric air. Those that are distinct lie on trends between the DICE sample area trend-line and the trend-line for MORB. This is compatible with variable MORB-plume source mantle mixing across Iceland. 167

168 Sample Table 11: Noble gas results for Icelandic samples. The heavy noble gases were not measured for every sample. For sample locations see chapter three. Concentrations are cc STP g -1 and errors are 1σ. 3 He Crushes (x10-12 ) ± 22 Ne (x10-12 ) ± 36 Ar (x10-12 ) ± 84 Kr (x10-12 ) ± 130 Xe (x10-14 ) 3 He/ 4 He ± ± 20 Ne/ 22 Ne ± 21 Ne/ 22 Ne ± 40 Ar/ 36 Ar ± 129 Xe/ 130 Xe ± 136 Xe/ 130 Xe ± (R/R A ) HB g TOTAL HUS g TOTAL LAUN g TOTAL HLO g TOTAL

169 3 He Sample Crushes ± (x10-12 ) 22 Ne (x10-12 ) ± 36 Ar (x10-12 ) ± 84 Kr (x10-12 ) ± 130 Xe (x10-14 ) 3 He/ 4 He ± ± 20 Ne/ 22 Ne ± 21 Ne/ 22 Ne ± 40 Ar/ 36 Ar ± (R/R A ) HLO g TOTAL Xe/ 130 Xe ± 136 Xe/ 130 Xe ± HLO g TOTAL KALF g TOTAL SN g TOTAL SN g (pyroxene)

170 3 He Sample Crushes ± (x10-12 ) 22 Ne (x10-12 ) ± 36 Ar (x10-12 ) ± 84 Kr (x10-12 ) ± 130 Xe (x10-14 ) 3 He/ 4 He ± ± 20 Ne/ 22 Ne ± 21 Ne/ 22 Ne ± 40 Ar/ 36 Ar ± 129 Xe/ 130 Xe ± 136 Xe/ 130 Xe ± (R/R A ) JARL g SN6 (pyroxene) g TOTAL SN12 (olivine) g TOTAL HB g TOTAL

171 3 He Sample Crushes ± (x10-12 ) 22 Ne (x10-12 ) ± 36 Ar (x10-12 ) ± 84 Kr (x10-12 ) ± 130 Xe (x10-14 ) 3 He/ 4 He ± ± 20 Ne/ 22 Ne ± 21 Ne/ 22 Ne ± 40 Ar/ 36 Ar ± 129 Xe/ 130 Xe ± 136 Xe/ 130 Xe ± (R/R A ) HLO g TOTAL JARL g TOTAL T g TOTAL

172 3 He Sample Crushes ± (x10-12 ) 22 Ne (x10-12 ) ± 36 Ar (x10-12 ) ± 84 Kr (x10-12 ) ± 130 Xe (x10-14 ) 3 He/ 4 He ± ± 20 Ne/ 22 Ne ± 21 Ne/ 22 Ne ± 40 Ar/ 36 Ar ± 129 Xe/ 130 Xe ± 136 Xe/ 130 Xe ± (R/R A ) EIRIK g TOTAL KALF g TOTAL HLO g TOTAL g TOTAL KS g TOTAL

173 3 He Sample Crushes ± (x10-12 ) 22 Ne (x10-12 ) ± 36 Ar (x10-12 ) ± 84 Kr (x10-12 ) ± 130 Xe (x10-14 ) 3 He/ 4 He ± ± 20 Ne/ 22 Ne ± 21 Ne/ 22 Ne ± 40 Ar/ 36 Ar ± 129 Xe/ 130 Xe ± 136 Xe/ 130 Xe ± (R/R A ) A g TOTAL HLO g TOTAL MID g TOTAL HLO g TOTAL KALF g TOTAL

174 3 He Sample Crushes ± (x10-12 ) 22 Ne (x10-12 ) ± 36 Ar (x10-12 ) ± 84 Kr (x10-12 ) ± 130 Xe (x10-14 ) 3 He/ 4 He ± ± 20 Ne/ 22 Ne ± 21 Ne/ 22 Ne ± 40 Ar/ 36 Ar ± 129 Xe/ 130 Xe ± 136 Xe/ 130 Xe ± (R/R A ) EIRIK g TOTAL ARM g TOTAL MID2 (olivine) g TOTAL MID2 (xenolith) g TOTAL MID g TOTAL

175 AIR WATER Figure 40: 130 Xe/ 36 Ar vs. 40 Ar/ 36 Ar for Icelandic samples that showed higher than air 3 He/ 4 He ratios. These data confirms that two-component mixing alone is not sufficient to explain the range of Icelandic noble gas abundances, as for the DICE samples area. The majority of samples show mixing between air and water (components C1 and C2 for the DICE samples), but a trend towards high 40 Ar/ 36 Ar and 130 Xe/ 36 Ar is also apparent. This trend is compatible with the trend towards the same mantle component seen in the DICE area Xenoliths in DICE area One particular analysis from the DICE area is worth discussing at this point. The DICE area contains a large amount of both olivine phenocrysts and gabbroic xenoliths. In this study a separate analysis of one of the gabbroic xenoliths found in the DICE area is reported. There is no reason to assume the xenolith component will be exempt from late stage atmospheric contamination, but the noble gases will be sited in fluid inclusions or the matrix itself rather than vesicles as in glass 175

176 samples. It may therefore be unaffected by the fractionated air component seen in glass samples and associated with vesicles. The 3 He/ 4 He ratio is 15.9±1 R A, which is a little lower than the ratio seen in the glass samples from this area, as observed by [P G Burnard et al., 1994]. This does not necessarily indicate a common mantle origin for both as the 3 He/ 4 He remains fairly consistent throughout this region of Iceland. The 4 He/ 21 Ne* range (2.57x10 5 to 3.13x10 5 ) and estimates for 4 He/ 40 Ar* and 4 He/ 136 Xe* ratios are fractionated from mantle production ratios in favour of the heavier noble gases. Degassing, crystallisation and diffusive noble gas loss are possible mechanisms for this fractionation, but these mechanisms are unlikely to produce a wide variety of fractionated ratios within a single sample asdifferent generations of vesicles are not present as they may be in a glass sample. The mantle endmember can therefore be treated as a single endmember composition, which may or may not have been fractionated from its original mantle composition by degassing, melting or crystallisation. This allows us to assume two-component, airmelt mixing for this sample and correct for the unfractionated air component to give the melt component. Air-corrected isotopic and elemental ratios are shown in Table 12 and Figure 41. Ranges are still given as mantle endmember 20 Ne/ 22 Ne can only be assumed to lie somewhere between 12.5 and 13.8 (Chapter 5). The isotopic ratios show a melt which lies on a similar neon trend to the DICE glass source melt while the 40 Ar/ 36 Ar ratio is much lower than the minimum estimate of 7050 for the DICE glass (Table 10). Melt 36 Ar/ 22 Ne, 84 Kr/ 22 Ne and 130 Xe/ 22 Ne ratios are much higher than the corrected ratio ranges for the DICE glasses (Table 10). The 84 Kr/ 36 Ar ratio is also somewhat lower than that of the DICE glasses although still above an air composition. 130 Xe/ 36 Ar falls within the range seen in the DICE glasses. 176

177 Seawater Bravo Dome Popping Rock Solar Limiting values for DICE source mantle Xenolith from DICE area Figure 41: Elemental abundances for various reservoirs are compared to the possible range of mantle-origin endmembers for a gabbroic xenolith from the DICE area. The depleted mantle is represented by the Bravo Dome and Pooping Rock samples Iceland s mantle beyond the DICE area As many of the highest 3 He/ 4 He values seen in central Iceland coincide with the seismically determined location of Iceland s proposed mantle plume [Wolfe et al., 1997], this pattern has been attributed to mixing between the MORB source mantle and a plume mantle component with much higher 3 He/ 4 He [E T Dixon, 2003; E T Dixon et al., 2000; Hilton et al., 1990; Mark D. Kurz et al., 1985]. Similarly, previous studies of 20 Ne/ 22 Ne and 3 He/ 22 Ne ratios have shown the light noble gases to be consistent with MORB/OIB mixing. Few Icelandic samples contain enough mantle neon to obtain 20 Ne/ 22 Ne ratios that are distinct from air. Those that do are consistent with MORB/OIB mixing across Iceland (Figure 39). It is likely that the 177

178 contaminant components identified in the DICE area are also be present in other Icelandic samples. The fractionated air and unfractionated air components will not change in composition, but the degassing range will vary from sample to sample depending on eruption conditions. Figure 42 shows that the DICE sample is unusual for Iceland in showing relatively little degassing. The degree of fractionation seen even within a single (non-dice) sample can be great. This is problematic as it means volatile loss and fractionation due to degassing will generally be high and mantle signatures will tend to be swamped by atmospheric contamination components. Quantifying the degassing process is difficult for the majority of Icelandic samples due to the near atmospheric 20 Ne/ 22 Ne ratios, which makes corrections to determine 4 He/ 21 Ne* prone to large errors. The large degree of degassing in Icelandic samples could be due to thinner overlying ice, slower ascent rate and quenching, or a combination of these factors. Atmospheric contamination in basaltic glasses is often correlated with vesicularity [Chris J. Ballentine and Barfod, 2000]. The potentially low 3 He concentrations of Iceland s source mantle and the high vesicularity of Icelandic samples then combine to make resolving mantle noble gases in Icelandic samples problematic. Nevertheless, some of the samples analysed in this appendix show trends towards mantle-like 20 Ne/ 22 Ne and 40 Ar/ 36 Ar ratios, and the new generation of multi-collector noble gas spectrometers may tease out further information on new areas of Iceland s source mantle from these samples. 178

179 Table 12: The noble gas composition of gabbroic xenoliths from the DICE area. Air contamination has been corrected for assuming two-component mixing between air and a xenolith component. Results are given as ranges as 20 Ne/ 22 Ne ratio of the xenolith component is not known and can only be constrained as being between 12.5 and Isotopic Ratio Xenolith from DICE area Elemental Ratio Xenolith from DICE area 3 He/ 4 He 15.9±1 20 Ne/ 22 Ne Ne/ 22 Ne Ar/ 36 Ar He/ 22 Ne Ar/ 22 Ne Kr/ 22 Ne Xe/ 22 Ne Xe/ 130 Xe 136 Xe/ 130 Xe - 84 Kr/ 36 Ar Xe/ 36 Ar ( ) x

180 DICE Samples Area Range of 4 He/ 136 Xe* ratios across the rest of Iceland. Figure 42: Across the rest of Iceland, degassing is much more extensive than for the DICE area. This results in low noble gas concentrations which, combined with atmospheric contamination, causes difficulties in resolving mantle components. 180

181 6 Chapter Six The halogen composition of Icelandic basalts 6.1 Abstract Volatile studies of Icelandic basalts have shown evidence for heterogeneity in Iceland s source mantle. The combined presence of a spreading ridge and a potential mantle plume make Iceland an ideal location to study the relationship between the deep and convecting mantle source reservoirs. In this study we present Cl, Br and I data for a suite of 19 basaltic samples taken from across Iceland. The Icelandic halogen data show no evidence for significant fractionation during degassing or melt generation, allowing constraints to be placed on the as yet uncharacterised halogen composition of Iceland s plume. Source estimates for the Br/Cl and I/Cl ratios for Iceland s plume are found to be (1.56±0.03) x 10-3 and (3.1±0.3) x 10-5, compatible with estimates for the MORB source mantle. Minimum estimates for the halogen source concentrations in central Iceland are 60±1 to 90±2 ppm Cl; 214±2 to 322±3 ppb Br; 7.2±0.1 to 10.8 ±0.1 ppb I. These values are approximately three times higher than estimates for the convecting mantle and correlate with the regions of Iceland that show high 3 He/ 4 He ratios and high source water contents. This may indicate a recycled halogen signature associated with Iceland s proposed mantle plume. 6.2 Introduction The volatile content of the mantle, particularly water content, is an important constraint on models of the mantle, impacting on mantle viscosity, conductivity and dynamics, melting processes and seismic studies [Hilton et al., 2002; Kelbert et al., 181

182 2009; Richard et al., 2002]. The proportion of volatiles recycled through downgoing slabs and the eventual origin of these volatiles will have a significant impact on the mantle s volatile budget and the high halogen concentrations in seawater and oceanic sediments relative to the mantle make them an ideal tracer of such recycling [John et al., 2010; John et al., 2011; M. A. Kendrick et al., 2011; Sumino et al., 2010]. In addition, the distinct I/Cl values found in seawater (9.5 x 10-7 ) and marine pore fluids (~2 x 10-5 up to 2 x 10-3 ), compared to the I/Cl of between 4.4 x 10-6 and 2.0 x 10-4 in the convecting mantle, makes this ratio a potentially excellent tracer of the involvement of recycled components within the mantle [Deruelle et al., 1992]. The mantle halogen data gathered so far have shown some evidence for a recycled halogen signature [R. Burgess et al., 2002; Byers et al., 1985; John et al., 2010; Stroncik and Haase, 2004]. A study of the behaviour of halogens in subducting serpentinite concluded that subduction can strongly affect the halogen budget of the mantle [John et al., 2011]. Laboratory studies of the properties of Br in diamond anvil cells imply that although shallow recycling of Br in arc magmas is extensive, a significant proportion of Br would be expected to be subducted with the down-going slab into the mantle [Bureau et al., 2010]. The halogens are incompatible elements during melt generation and this allows basaltic samples to provide a reliable window onto mantle compositions. Studies of both normal and enriched mid-ocean ridge basalts (MORBs) have constrained the halogen composition of the convecting mantle [Bonifacie et al., 2008; Deruelle et al., 1992; Albert Jambon et al., 1995; Mark A. Kendrick et al., 2012; Schilling et al., 1980].Kendrick et al. (2012) s data from Pacific E-MORBs ( enriched' MORBs), obtained using the sensitive Ar-Ar halogen technique, gave a Br/Cl of (1.64 ± 0.22) x 10-3 and an I/Cl of (4 ± 3) x 10-5 as representative values for the MORB source mantle. Also using the Ar-Ar technique, halogen studies of fluid inclusions from diamonds have been shown to contain a component similar to the MORB halogen composition [R. Burgess and Turner, 1995; R. Burgess et al., 2002; Ray Burgess et al., 2009; Johnson et al., 2000]. Searching for recycled signatures in submarine glasses is often complicated by the potential for the submarine quenching environment to produce a seawater contaminant signature within these samples [Kent et al., 2002; Kent et al., 1999]. 182

183 The Icelandic samples used in this study are mainly subglacial in origin which avoids this issue. In addition, Iceland represents not only the MORB source mantle but hotspot activity, which is thought to be due to a plume originating in the deep mantle [Helmberger et al., 1998; Shen et al., 1998; Wolfe et al., 1997]. The presence of subducted slabs and recycled material in this deep mantle region has been invoked to explain some of the geochemical signatures observed in ocean island basalts (OIBs) [C. J. Ballentine et al., 2005b; Brandenburg and van Keken, 2007a; Hofmann, 1997; Holland and Ballentine, 2006]. Geochemical studies across Iceland have shown evidence for a variety of recycled material in Iceland s source mantle, including a HIMU-like and an EM2 component [P Burnard and Harrison, 2005; Macpherson et al., 2005; Peate et al., 2010; Prestvik et al., 2001]. Noble gas helium and neon geochemistry has also shown the influence of a solar-like primordial signature associated with a deep mantle reservoir [E T Dixon et al., 2000; D. Harrison et al., 1999; Mark D. Kurz et al., 1985; Nichols et al., 2002]. This makes Iceland an ideal place to investigate the interplay of these mantle components. Data on the halogen composition of Icelandic samples are sparse: [Unni and Schilling, 1978] provide estimates of 17±2 ppm and 0.06±0.01 ppm for the Cl and Br concentrations of Iceland s MORB component and 61±6 ppm and 0.21±0.02 ppm for the Cl and Br concentrations of the more halogen-enriched component seen in the direction of Iceland s plume. Variations in halogen composition from the Rekyjanes ridge into SW Iceland show that OIBs potentially carry a different halogen signature to MORBs, with higher halogen concentrations associated with Iceland s OIB signature [Unni and Schilling, 1978], but the central Iceland halogen signature has not been characterised. In this study we expand the data on Iceland s halogen signature. We present new results for Cl, Br and I from 19 samples (16 basaltic glasses and three olivine samples) taken from Iceland s western and northern rift zones, central Iceland and the stratovolcano Oræfajökull. This data-set for halogen concentrations across Iceland helps to constrain models of Iceland s mantle, in particular the involvement and interaction of primordial and recycled signatures. 183

184 6.3 Samples The majority of Iceland s current volcanic activity is focused on its eastern, western and northern rift zones (see Figure 43). Iceland s rift zones exhibit a number of subglacial volcanic features (tuyas and tindars; see Chapter 3 for details) whose lowest sequence is often a thick pillow basalt pile formed during the early stages of an eruption [Jones, 1968; Tuffen et al., 2010]. The glassy rims of these pillows are ideal for volatile analysis as the high subglacial eruption pressures and fast quenching rates of the samples minimises volatile loss. The clear subglacial origin of these samples also eliminatesthe issue of seawater contamination often present in MOR glassy samples. Samples ERIK3, HL102 and Hl103 are from the base of classic Icelandic tuyas (Hloðufell - [Skilling, 2009] and Eiriksjökull). JARL8 is from the base of a tindar in the WVZ (a fissure eruption that shows no evidence of having broken through the overlying ice-sheet). LAUN5, HUS2 and MID1 are from pillow piles exposed in quarries on the Reykjanes peninsula. The MID1 pillows are picritic basalts containing large olivine phenocrysts and gabbroic nodules, and are from the same sample area as the DICE10 sample (Figure 25), which has shown unusually high mantle noble gas concentrations and which has been taken as representative for the noble gas composition of Iceland s mantle [D. Harrison et al., 1999; Mukhopadhyay, 2012; Mario Trieloff et al., 2000]. FOG4 is a glassy sample that was collected from pillows found at the base of a sheet sequence found in SE Iceland a thick-ice origin has been proposed for this sequence making the retention of significant mantle volatiles a possibility [Smellie, 2008]. HOF10 is a glassy sample from Oræfajökull, a stratovolcano in SE Iceland that has displayed an enriched geochemical signature [Prestvik et al., 2001]. The remaining samples were not collected by the authors (thanks to Alex Nichols for their provision) but are all glass samples from pillow rinds in the NVZ and central Iceland [Nichols et al., 2002]. 184

185 Eiriksjökull plume centre Hloðufell Figure 43: Map of Iceland showing sample locations collected in this study (solid diamonds) and those analysed in this study donated by Alex Nichols [Nichols et al., 2002]. Map adapted from [Sigmarsson and Steinthórsson, 2007]. 185