Als Dissertation genehmigt von der Naturwissenschaftlichen Fakult¨at der Friedrich-Alexander-Universit¨at Erlangen-Nurn¨ berg

Tag der mundlic¨ hen Prufung:¨ 24.01.2014 Vorsitzender der Promotionsorgans: Prof. Dr. J. A. C. Barth Gutachter: Prof. Dr. K. M. Haase Prof. Dr. J. Koepke Contents

List of Figures 6

List of Tables 8

Abstract 9

Kurzfassung 10

Statement of candidate 12

Full publication list 13

Acknowledgments 15

1 Introduction 16

2 Aims of the study 28

3 Oxygen isotope evidence for the formation of andesitic-dacitic magmas from the fast-spreading Pacific-Antarctic Rise by assimilation-fractional crystallisation 30 3.1 Abstract ...... 30 3.2 Introduction ...... 31 3.3 Geological Background ...... 32 3.4 Results ...... 32 3.4.1 Chemical and O isotopic compositions of the PAR glasses ..... 32 3.4.2 Petrography and mineral chemistry ...... 44 3.5 Discussion ...... 45 3.5.1 Liquid lines of descent and fractional crystallisation processes ... 45 3.5.2 Constraints on assimilation processes in the generation of silicic PAR magmas ...... 46 3.5.3 Evidence from mineral compositions for melt evolution and the link to oceanic ferrogabbros ...... 47 3.5.4 Fractionation of the PAR magma: oxide crystallisation and its effects on the magma composition ...... 48 3.5.5 Thermobarometric constraints on the depth of the mafic magma reservoir ...... 49 Contents 4

3.5.6 A quantitative AFC model for the PAR lavas ...... 50 3.5.7 The nature of the assimilant and the role of amphibole ...... 51 3.5.8 Potential relationship between occurrence of silicic magmas along spreading axes, tectonic and hydrothermal processes ...... 53 3.6 Conclusions ...... 53 3.7 Acknowledgements ...... 55

4 Constraints on the formation of geochemically variable plagiogranite intrusions in the Troodos Ophiolite, Cyprus 56 4.1 Abstract ...... 56 4.2 Introduction ...... 57 4.3 Geological Background ...... 57 4.4 Samples and analytical methods ...... 59 4.5 Results ...... 61 4.5.1 Occurrence of the plagiogranites and field observations ...... 61 4.5.2 Petrology and mineral chemistry ...... 61 4.5.3 Major element composition of the rocks from plagiogranite intrusions 64 4.5.4 Trace element composition of the plagiogranitic intrusives and com- parison with the glassy lava ...... 65 4.6 Discussion ...... 68 4.6.1 Hydrothermal alteration of the samples and stratigraphic context of the plutonic section ...... 68 4.6.2 Relations between evolved intrusive rocks and the lavas ...... 72 4.6.3 The relationship of the plagiogranitic intrusions to the mafic crustal plutonic rocks ...... 73 4.6.4 Generation of the plagiogranite magmas: fractional crystallization versus partial melting ...... 73 4.7 Conclusions ...... 77 4.8 Acknowledgements ...... 77

5 Constraints on the evolution of the crust of the Semail ophiolite (Oman) from the geochemical composition of plagiogranites 97 5.1 Abstract ...... 97 5.2 Introduction ...... 98 5.3 Geological Background ...... 99 5.3.1 Geographical and geological overview ...... 99 5.3.2 Geochemical overview ...... 99 5.3.3 Ophiolite age ...... 101 5.4 Samples and analytical methods ...... 101 5.4.1 Sampling ...... 101 5.4.2 Analytical methods ...... 103 Contents 5

5.5 Results ...... 106 5.5.1 Petrology ...... 106 5.5.2 Geochemistry ...... 106 5.6 Discussion ...... 114 5.6.1 Hydrothermal alteration of the plutonic rocks ...... 114 5.6.2 The geochemical relationship between the plutonic and volcanic rocks114 5.6.3 Formation of the felsic magmas ...... 116 5.6.4 Implications for the tectonic setting during formation of the Oman ophiolite ...... 119 5.6.5 Evidence of source depletion, decoupling effects and the nature of the slab component ...... 120 5.6.6 The influence of slab derived fluids on magma differentiation ... 121 5.6.7 Time interval between the two magmatic events ...... 122 5.7 Conclusions ...... 123 5.8 Acknowledgements ...... 123

6 Findings and Outlook 140

7 Supplementary materials 170 List of Figures

1.1 The Earth´s continental and oceanic crust...... 17 1.2 Composition of continental and oceanic crust ...... 18 1.3 Cross-section through the Oman ophiolite ...... 19 1.4 Processes that probably take place in an axial magma chamber...... 21 1.5 Five life-cycles of a SSZ ophiolite ...... 24 1.6 Tonalite-trondhjemite-granodiorite discrimination diagram ...... 25

3.1 Bathymetric map of the Pacific-Antarctic ridge and the sample locations . 33 3.2 Neodymium isotopes and Nb/Zr ratio of the PAR glasses versus Latitude 36 3.3 Major elements of the PAR glasses ...... 37 3.4 Trace element contents of the PAR ...... 40 3.5 Sulphur and Cl contents of the PAR and EPR glasses ...... 41 3.6 Trace element ratios of the PAR and EPR glasses ...... 42 18 3.7 H2O/Ce, Li/Ce and δ O of the PAR, EPR and GSC glasses ...... 44 3.8 Oxygen isotopes versus Cl/K of the PAR and EPR and assimilation-fractional crystallisation trend ...... 51 3.9 Formation of the basaltic, FeTi basaltic and silica rich lavas at the PAR . 54

4.1 Map of the Troodos Ophiolite in Cyprus with detailed geology of the Mount Olympos area and the distribution of plagiogranite intrusions ...... 60 4.2 Field photographs of plagiogranites in the Troodos Ophiolite ...... 62 4.3 Feldspar (anorthite-albite-orthoclas) composition diagram ...... 63 4.4 Tonalite-trondhjemite-granodiorite discrimination diagram ...... 64 4.5 Major element variations of the plagiogranites, lavas and mafic plutonics . 66

4.6 Incompatible elements and transition metals versus SiO2 from the pla- giogranites, lavas and mafic plutonics ...... 67 4.7 Rare earth element patterns of the plagiogranite groups ...... 69 4.8 Immobile-trace element ratios of the plagiogranite groups and lavas .... 70 4.9 Loss on ignition versus mobile and immobile elements ...... 71 4.10 Fractional crystallization and partial melting models in terms of incompat- ible elements ...... 76

5.1 Maps of the Sumail ophiolite and sample locations ...... 100 5.2 Field photographs of plagiogranites in the Oman ophiolite ...... 102 List of Figures 7

5.3 Normative calculated plutonic rock classification ...... 107 5.4 Major element variations of the plagiogranites and wall-rocks, lavas, sheeted dyke rocks and plutonic rocks ...... 108 5.5 Multi element diagrams of the plagiogranites and lavas ...... 109 5.6 Incompatible element compositions, ratios and Nd isotope ratios of the volcanic and plutonic rock groups ...... 110 5.7 Major and trace elements versus loss on ignition ...... 111 5.8 Incompatible element ratios of the plagiogranite groups and lavas ..... 113 5.9 Isotope ratios and mother/daughter trace element ratios of the plagiogran- ites and lavas ...... 115 5.10 Discrimination between fractional crystallization and partial melting pro- cesses ...... 117 5.11 Magmatic age of plagiogranite sample locations versus incompatible ele- ments and isotope ratios ...... 122 List of Tables

3.1 Representative major and trace elements and isotopic compositions of silicic and basaltic glasses from the Pacific-Antarctic Rise ...... 34 3.2 Details of the AFC calculations of the PAR glasses ...... 38 3.3 Thermobarometry on the basis of clinopyroxene–melt equilibrium calculations 50

4.1 Major and trace elements of Troodos plagiogranites and aphyric dykes .. 78 4.2 Parameters for modeling Troodos plagiogranitic melt evolution ...... 93 4.3 Least squares model and trace element modeling ...... 94 4.4 Melting of amphibolite ...... 96

5.1 Correlation of plagiogranite sample locations with U/Pb age analyses of zircons from early respectively late stage plutonics ...... 103 5.2 Standard measurements ...... 104 5.3 Major elements of the Oman plagiogranites...... 125 5.4 Trace elements and isotope ratios of the Oman plagiogranites ...... 130 Abstract

The magmatic oceanic crust mainly consists of mafic rock types like plutonic gabbros and volcanic basalts. Comparable small amounts of felsic rocks with >60 wt% SiO2 are known in the oceanic crust in forms of intrusive veins and plagiogranite bodies but only rarely in form of erupted lavas. Plagiogranitic rocks are well known from ophiolite complexes, which represent a typical sequence of rocks formed in an oceanic setting obducted and exposed on land. Ophiolites therefore provide an easy and less expensive opportunity to study oceanic crustal rocks and those magmatic evolution processes.

This thesis focuses on the particular information content provided by felsic rocks concern- ing various processes during crustal formation. The combination of parameters like e.g. crustal thickening, hydrothermal alteration and assimilation of crustal rocks during melt fractionation beneath a segment of the Pacific Antarctic Rise is discussed in chapter 3. I used major-, trace elements and isotope ratios (δ18O) of fresh and young glassy lavas (basalts, andesites, dacites) and mineral compositions to model the amount of crystal fractionation and assimilation and to detect the composition of the assimilated material. The result of this study indicates that distinct amounts (30%) of hydrothermal modified crustal rocks must be assimilated to produce the felsic lavas.

I compared several felsic plutonic rock samples (plagiogranites) from two large and well- preserved ophiolite complexes (Troodos and Oman) with each other and with the as- sociated mafic rocks (gabbros, sheeted dykes and lavas) in chapters 4 and 5. Major-, trace elements and isotope ratios (Sr, Nd, Hf) of whole-rocks and mineral analyses (major elements) were used to demonstrate the genetic relation of the felsic intrusives with the majority of the mafic crustal rocks. The result of these studies indicates that the pla- giogranites in the Troodos and Oman ophiolite complexes comprise geochemical different groups. The plagiogranite groups are in each case related to different mafic crustal rocks and lavas indicative for a special magmatic phase. In particular the isotope analysis provide a high information potential including time progressive melt contamination during crust formation.

In summary, both processes, fractional crystallization of mafic melt and assimilation of partial molten crustal rock are important processes to generate felsic melt within oceanic crust settings. Kurzfassung

Die magmatische ozeanische Kruste besteht haupts¨achlich aus mafischen Gesteinen wie den plutonischen Gabbros und den vulkanischen Basalten. Vergleichsweise geringe Men- gen felsischer Gesteine mit > 60 wt.% SiO2, sind auch aus der ozeanischen Kruste bekannt, meist in Form von intrusiven Adern und Plagiogranitk¨orpern, jedoch nur selten als erup- tierte Laven. In Ophiolite Komplexen, welche typische Gesteinssequenzen repr¨asentieren, die in einem ozeanischen Milieu entstanden, obduziert und auf dem Land freigelegt wurden sind Plagiogranitintrusionen vergleichsweise h¨aufig. Ophiolite bieten daher eine einfache und kostengunstige¨ M¨oglichkeit, um ozeanische Krustengesteine und ihre magmatischen Entwicklungsprozesse zu untersuchen.

Diese Promotion untersucht speziell den Informationsgehalt felsischer Gesteine in Bezug auf Krustenbildungsprozesse. Kombinationen verschiedener Parameter, wie z.B. Krusten- verdickung, hydrothermale Ver¨anderung und Assimilation von Krustengesteinen w¨ahrend der Schmelzfraktionierung an einem Segment des Pazifisch-Antaktischen Ruc¨ kens, wer- den im Kapitel 3 diskutiert. Ich nutzte Haupt- Spurenelemente und Isotopenverh¨alt- nisse (δ18O) von frischen und jungen glasigen Laven (Basalte, Andesite, Dacite) und die Mineralzusammensetzungen einerseits, um die Anteile von Kristallfraktionierung und Assimilation zu modelieren und andererseits, um die Zusammensetzung des assimilierten Materials zu ermitteln. Das Ergebnis dieser Studie ergibt, dass gr¨oßere Mengen (30%) hydrothermal ver¨anderten Krustengesteins assimiliert werden mussen,¨ um diese felsischen Laven zu erzeugen.

In Kapitel 4 und 5 vergleiche ich diverse felsische Intrusivgesteine (Plagiogranite) von zwei großen und gut erhaltenen Ophiolite Komplexen (Troodos und Oman) jeweils un- tereinander und mit den assoziierten mafischen Gesteinen (Gabbros, Dike-Komplex und Laven). Haupt-, Spurenelemente und Isotopenverh¨altnisse (Sr, Nd, Hf) des Gesamt- gesteins und Mineralanalysen (Hauptelemente) wurden genutzt, um die Zusammenh¨ange zwischen den felsischen Plutoniten und der mafischen Krustengesteine zu untersuchen. Die Ergebnisse dieser Studien demonstrieren, dass die Plagiogranite innerhalb des Troodos, sowie Oman Ophiolite Komplex geochemisch unterscheidbare Gruppen umfassen. In beiden F¨allen lassen sich die einzelnen Plagiogranitgruppen unterschiedlichen mafischen Krustengesteinen und Laven zuordnen, welche jeweils Indikativ fur¨ eine spezielle mag- matische Phase sind. Insbesondere die Isotopenanalysen beinhalten ein hohes Informa- tionspotential, z.B. in Bezug auf zeitlich voranschreitende Schmelzkontamination w¨ahrend der Krustenbildung. Kurzfassung 11

Zusammenfassend ergibt sich, dass sowohl fraktionierte Kristallisation von mafischer Schmelze als auch Assimilation von partiell geschmolzenem Krustengestein wichtige und nur schwer abgrenzbare Prozesse bei der Entstehung von felsischen Schmelzen innerhalb von ozeanis- cher Kruste sind. Statement of candidate

I certify that the work in this thesis entitled The generation of felsic magmas in the ” oceanic crust: assimilation-fractional crystallization processes versus re-melting of the crust“ has previously not been submitted for any degree nor has it been submitted as part of requirements for a degree to any other university or institution other than the Friedrich-Alexander-University Erlangen-Nurn¨ berg. I also certify that the thesis is a new, original piece of research and it has been written by me. Any help and assistance that I have received during my research work and the preparation of the thesis itself have been appropriately acknowledged. In addition, I certify that all information sources and literature used are indicated in the thesis. This thesis contains material that has been published or accepted for publication in peer- reviewed ISI-journals or is in preparation for publication, as follows:

1) “Oxygen isotope evidence for the formation of andesitic-dacitic magmas from the fast-spreading Pacific-Antarctic Rise by assimilation-fractional crystallisation” has been acctepted for publication in Chemical Geology. My contribution to this publication consisted of parts of analytical work, data analy- ses and interpretation and parts of writing the text, resulting in a total contribution of about 65%.

2) “Constraints on the formation of geochemically variable plagiogranite intrusions in the Troodos Ophiolite, Cyprus” has been revised (after major revision) for publication in Contributions to Mineralogy and Petrology. My contribution to this publication consisted of sampling, parts of analytical work, data analyses and interpretation and parts of writing the text, resulting in a total contribution of about 75%.

3) “Constraints on the evolution of oceanic plagiogranites in the Semail- ophiolite, Oman” is currently in preparation for publication. My contribution to this publication consisted of sampling, parts of analytical work, data analyses and interpretation and parts of writing the text, resulting in a total contribution of about 85%.

Erlangen, 20.10.2013 Full publication list

Publications in peer-reviewed journals

Freund, S., Haase, K. M., Keith, M., Beier, C., Garbe-Sch¨onberg, D. (revised after major revisions) Constraints on the formation of geochemically variable plagiogranite intrusions in the Troodos Ophiolite, Cyprus. Contributions to Mineralogy and Petrology

Freund, S., Beier, C., Krumm, S., Haase, K.M. (2013): Oxygen isotope evidence for the formation of andesitic–dacitic magmas from the fast-spreading Pacific–Antarctic Rise by assimilation–fractional crystallisation. Chemical Geology 347 (271-283)

Conference abstracts

Freund, S., Erdmann, M., Koepke, J., Hauff, F., Haase, K. M. (2013): Crustal evolution and petrogenesis of silicic plutonic rocks within the Oman ophiolite – petrological and geochemical investigations. Goldschmidt, Florenz.

Erdmann M., Fischer L.A., France L., Freund S., Koepke J. (2013): Hydrous partial melting at the dike/gabbro transition at fast-spreading ridges - an experimental study. IODP 335 2nd post-cruise meeting, Corsica.

Erdmann, M., Freund, S., Fischer, L., Haase, K., Koepke, J. (2012): Plagiogranites in the Oman ophiolite – a key to understand the formation of SiO2-rich magmas within fast-spreading oceanic crust. International Conference on the Geology of the Arabian Plate and the Oman Mountains, Muscat.

Erdmann, M., Koepke, J., Freund, S. (2012): Phase relations and distribution coeffi- cients for evolved lavas from the Pacific-Antarctic Rise. EMPG conference, Kiel.

Erdmann M., Freund S., Haase K. M., Koepke J. (2012): Application of techniques to enlarge meltpools and crystals in crystallization experiments: A case study in an andesitic system from the Pacific-Antarctic Rise. EMC, Frankfurt. Full publication list 14

Freund, S. Keith, M., Endres, T., Schmidt, H., Beier, C., Regelous, M., Haase, K.M. (2012): Formation of a silicic upper oceanic crust in the Troodos ophiolite, Cyprus. EMC, Frankfurt.

Freund, S., Haase, K. M., Beier,C. Regelous, M. (2011): Felsic magma generation in the oceanic crust: a geochemical study of Pacific Antarctic Rise lavas. Goldschmidt, Prag.

Freund, S., Haase, K.M., Beier, C. (2010): The petrogenesis of evolved magmas within the oceanic crust (Pacific-Antarctic-Rise) – geochemical and isotopic investigation on volcanic lavas. DMG, Munster.¨ Acknowledgments

First, I would like to thank Prof. Dr. Karsten M. Haase for giving me the reliability to work on this project successfully and supervising me patiently during all highs and lows. Sincerest thanks are also due to Dr. Christoph Beier, Dr. Marcel Regelous and Dr. Stefan Krumm for their support and discussions, their feedback and help during computer problems or softwarecrashes.

The project has greatly benefited from the harmonic collaboration and fruitful discussions at any state of this project with Prof. Dr. Jurgen¨ Koepke and Martin Erdmann.

I would like to acknowledge the scientific cooperations of my colleagues Dr. Dieter Garbe- Sch¨onberg, Prof. Dr. Reiner Klemd, Dr. Helene Br¨atz and Dr. Folkmar Hauff. Special thanks are also directed to the great and essential work of the technical staff at the GeoZentrum Nordbayern: Melanie Hertel, Veronika Kuhnert,¨ Konrad Kunz, Christine Scharf, Bernd, Schleifer, Christian Abe, Gudrun Klein.

Very important were also the discussions, support and some times simple presence of my colleagues at the GeoZentrum Nordbayern: Melanie, Henning, Inga, Maria, Philipp, Felix, Claudio, J¨org, Manja, Manuel, Mario, Andi, Ste, Christoph, Fabian, Lukas and Ansi. Furthermore, I would like to express my gratitude to my friends: you helped me a lot by doing sports and giving me important mental support: Anne, Eva, Gerti, Jonas, Niggi, Christiane, Seb, Felix, Micha, Ludwig, Christina, Heinz, Till, Fritz and Nora.

Special thanks go to my best mates Rike Baiker & Dr. Doreen Zajic: you always did a great job

I deeply acknowledge the great support of my parents Dr. Martina Freund and Joachim Liebe-Freund, my grandmother Marianne Liebe and in thoughts to my grandfathers Erich Freund and Josef Liebe. Thanks is also owed to my family in law.

Most importantly, my heartfelt thank goes to Lukas Pflug for helping with LATEX, com- puter problems, thousands of discussions, feedbacks and his protective presence at any minute.

Finally, I would like to thank Deutsche Forschungsgemeinschaft (DFG) and the gender Buro.¨

I apologise in advance in case that somebody is missing here. I am solely responsible for any remaining errors and omissions of this work. 1 Introduction

Evolution of the Earth’s rock types

The Earth’s crust contains a wide range of different rock types, but can broadly be divided into the continental crust, forming the continents and continental shelf zones and the oceanic crust, which mostly forms the ocean basins (Fig. 1.1). Compared with other planetary crusts in the solar system, the formation process of the Earth’s continental crust appears unique (Taylor and McLennan, 2001). The anorthosites of the Moon’s crust represent a primary crust, which floated on an anhydrous magma ocean as a result of the melting of the moon during its accretion, whereas the lunar maria are the product of secondary partial melting of the Moon’s mantle, similar to the oceanic crust formation of the Earth (Taylor and McLennan, 2001).

The magmatic differentiation of the Earth into the lighter, more evolved continental crust, which is buoying upwards relative to the heavier, denser oceanic crust started shortly after the formation of the Earth at 4.51-4.55 Ga (e.g., Manhes et al. 1980). Neodymium model ages on upper continental crust rocks suggest that about 50% of the continental crust mass was generated by the end of the Archean (Taylor and McLennan, 1985). However, there are clear differences concerning the Archean craton rock composition (e.g. bimodal silica distribution, Na/K, REE, lack of Eu anomaly) compared to younger continental rocks generated along convergent continental margines (Kemp and Hawkesworth, 2003). The oldest K-rich granites, comparable with modern continental granites occur towards the end of the Archean after ∼3.0 Ga (Condie, 1993). The chemical differentiation of the Earth was therefore a process influenced in the Archean only by partial melting of the mantle and post-Archean also by further evolutionary processes like repeated remelting, fractional crystallization, assimilation or mixing of the melts (Taylor and McLennan, 1985; Kemp and Hawkesworth, 2003). By this means, the Earth’s material exsolved itself with the time and the results are geochemically and hence geophysically different rock domains: oceanic and continental crust (e.g. Hofmann, 1988; Wyllie, 1988). The origin of the petrologic differences between the continental and oceanic crusts is a matter of active debate.

The igneous oceanic crust is generally believed to consist of gabbros and basalts, whereas the average continental crust is andesitic in composition. Overall the igneous part of the continental crust contains a wide range of magmatic rock types, e.g. diorites, tonalites, granites and their metamorphic products. Generally, the dense oceanic crust is recycled

1. Introduction 19

Sediments

Layer 2 Lavas sheeted dikes Plagiogranites late varitextured intrusiv Gabbros Gabbros foliated Gabbros Layer 3 Wherlite intrusions

Moho Gabbro sills

Layer 3 Peridotites, Mantle section Pyroxenites Dunites

Harzburgites Metamorphic sole Granat Amphibolites

Figure 1.3: Generalized cross-section through the Oman ophiolite (modified from Nicolas, 1988).

Whalen et al. 2002). The occurrence of felsic rocks in the oceanic crust is of particular importance to better constrain the average composition of the oceanic crust, which plays an essential role during subduction and recycling. In particular, the question arises as to which processes lead to the formation of silica-rich rocks within the oceanic crust?

The overall aim of this thesis is to better constrain the processes and sources of felsic rocks forming along oceanic spreading centers along mid-ocean ridges and back-arc spreading centers.

Structure and geochemistry of the oceanic crust

Seismic models indicate variable velocities of P-waves that are interpreted as different layers within the oceanic crust (e.g. Spudich and Ocrutt, 1980 and references therein). This model is comparable to the petrological model based on field studies on ophiolites (e.g. Boudier and Nicolas, 1985, Nicolas et al. 1988) (Fig. 1.3) and deep drill cores (e.g. Detrick et al. 1994; Alt et al. 1996) confirming the concept of layered sequences in the oceanic crust. The igneous portion of the oceanic crust is commonly subdivided into 1. Introduction 20 three distinct layers (Fig. 1.3): pillow basalts and the sheeted dike complex are part of the volcanic layer (2) that is subdivided into the extrusive layer (2A) and the intrusive layer (2B). The gabbros form the layer 3 and the transition into peridotites of layer 4 represents the petrological Moho. The average total thickness of this igneous oceanic is about 7.1 0.8 km (White et al. 1992).

Incompatible element-depleted tholeiitic basalts are the product of mantle melting beneath a mid-ocean ridge (MORB) and the rocks are transported off-axis to each side of the spreading-center and finally recycled into the mantle by subduction. Additionally, due to subduction pelagic and continental sediments reach the mantle. Subduction of hydrother- mally altered oceanic crust initiates arc volcanism and the modified composition of the slab (including sediments) affects the composition of the arc magmatism (e.g. Kelemen et al. 2003 and references therein). Whereas MORB typically show only small ranges in

SiO2 content (48-52 wt%), arc lavas can reach SiO2 values > 60 wt% over a wide range in MgO content (Kelemen et al. 2003 and references therein). The differences in major element contents between parental MORB and arc melts and the comparable high water content in arc melts (e.g. Anderson, 1974; Sobolev and Chaussidon, 1996; Pichavent et al. 2002) results in a wide range of possible rock types in arcs e.g. (calc-alkaline-) basalts, andesites, adakites, ankaramites, picrites or boninites.

Spreading not only occurs at mid-ocean ridges but also in other decompression settings like e.g. a back-arc basin. The composition of back-arc spreading centres differs from that of normal mid-ocean ridge magmas, especially in terms of incompatible trace elements. Thus, Cs, Rb, Ba, Th, U, K, Sr, Pb and light rare earth elements (LREE) are enriched in most arc-related lavas compared to normal MORB, whereas Nb and Ta are depleted relative to other incompatible elements (Kelemen et al. 2003 and references therein). Nevertheless, in both settings, depleted mantle melts accumulate and generate gabbros, sheeted dikes and lavas.

Ophiolites were long interpreted as representing sections of normal mid-ocean ridges (e.g. Coleman, 1981, Nicolas et al. 1988, Boudier et al. 1988) and therefore most suitable for studying normal ocean crust formation. Recently, several authors suggest that the huge ophiolites (e.g. Semail ophiolite) are subduction zone-related (possible back-arc) oceanic crust (e.g. Rautenschlein et al. 1985, Muenow et al. 1990, Shervais 2001; Pearce and Robinson, 2010). Nevertheless, ophiolites may provide the unique opportunity to compare observations made along recently formed oceanic crust with older oceanic crust segments.

Magma composition and the important role of the magma chamber

Basalts and gabbros represent the volumetric majority of the igneous oceanic crust. In a simple model, partial melting (up to 20 %) of the depleted mantle produces tholeiitic

1. Introduction 22

It is largely accepted that melts are stored within shallow magma chambers of about one to several hundred meters in size below ridge axes (e.g. from geophysical studies: e.g. Morton and Sleep, 1985; Sinton and Detrick, 1992; Dunn et al. 2000), but generally, conclusions about magma chamber processes beneath recent mid-ocean ridges and the overall composition of the oceanic crust are mainly drawn by analyzing volcanic rocks (mid-ocean ridge basalts) and the minerals in these lavas.

Additionally, the presence of oceanic crustal rocks obducted onto continental crust (ophio- lites) offers a view into the oceanic crust via cross-sections and therefore facilitates insights within (frozen) magma chambers presented in form of upper gabbros (e.g. Smewing et al. 1984, Kelemen et al. 1997). Because drilling in recent oceanic crust is relatively difficult and expensive and only provides limited insight into the oceanic crust, a better understanding of the formation of mafic and evolved rocks in ophiolites provides an ideal alternative for studying deep oceanic rocks. Particularly, transects within the lower oceanic crust, even into the mantle lithology and large-scale rock relations and structures are relatively easy to study. Typical examples of ophiolites are: Macquarie Island (Australia), Bay of Island ophiolite (Newfoundland) and the Tethyan ophiolites all around the Mediterranean sea (Fig. 1.5) e.g. the Semail ophiolite (Oman and United Arab Emirates), the Troodos ophiolite (Cyprus), the Ligurian ophilites (Apennines, Italy), the Kizildag ophiolite (Turkey), the Beni Bousera ophiolite (Morocco)) and the ophiolites in the Himalaya. In chapter 2 and 3 of this thesis I have focused on the formation process of the plagiogranites occurring within the two largest and best-preserved complexes, the Semail and the Troodos ophiolites.

A view into the deep oceanic crust: Ophiolites and their formation processes

During the Penrose field conference (1972) a structural oceanic crust model and a relation to the structure and petrology of ophiolites was suggested. Prior to the conference Moores and Vine (1971) had identified the dike complex within the Troodos Mountains as identical to oceanic sheeted dikes.

Firstly spreading and then closing of an ocean must occur prior to obduction of oceanic crust onto continental crust. Mostly smaller ophiolites are interpreted to be tectonically disrupted oceanic crust fragments, intercalated within collisional terranes during a rising orogen. This small occurrence of oceanic lithosphere rocks on land often lacks a full stratigraphic sequence of oceanic crust. For example, the metamorphic tectonic m´elange of ocean crust fragments of the Alps and the Massive Central (e.g. Plankogel formation), which marks an ancient oceanic suture zone (Matte, 2001 and references therein).

Several authors suggested an intermediate to fast spreading “supra-subduction-zone set- ting” (SSZ) as most likely tectonic setting for the formation of large ophiolite complexes 1. Introduction 23 like the Troodos and the Oman ophiolite (e.g. Rautenschlein et al. 1985, Muenow et al. 1990, Shervais 2001; Dilek and Furnes, 2009, Pearce and Robinson, 2010). Shervais (2001) defines five life-cycles of a SSZ ophiolite (Fig. 1.6) as follows:

1. The birth-stage (Fig. 1.6a) includes the formation of the ophiolite in the vicinity of a subduction zone and the initiation of arc volcanism.

2. In the youth-stage (Fig. 1.6b) the partial melting of the previously depleted as- thenosphere mantle continued.

3. During the maturity-stage (Fig. 1.6c) a semi-stable calc-alkaline arc volcanism is often established.

4. Within the death-stage (Fig. 1.6d) the active spreading and the volcanism come to a sudden stop.

5. The resurrection-stage (Fig. 1.6e) is the process of crustal emplacement and obduc- tion upon continental crust.

Most ophiolite complexes do not undergo all of those stages (Shervais, 2001).

The occurrence of several evolved rocks in numerous single intrusions in the crustal (and even in the mantle) section of the Semail Ophiolite and the crustal section of the Troodos compared to the limited amounts of felsic rocks sampled from recent oceanic crust, leads to the question if the special circumstances of the SSZ ophiolite life-cycle possibly facilitate the formation of felsic rocks. According to Shervais (2001) plagiogranites are a product of the mature stage and part of the high-level intrusions feeding the volcanic sequence but the processes leading to the formation of the felsic rocks are still debated.

1. Introduction 26

Processes possibly leading to the formation of felsic melt within the oceanic crust

The formation of felsic magma (> 60 wt% SiO2) within the mafic ocanic crust is a matter of debate since several decades. The earliest interpretations of oceanic plagiogranites explained these by low-pressure fractional crystallization of a basaltic melt (e.g. Coleman and Peterman, 1975; Pallister and Knight, 1981; Lippard et al, 1986). Additionally, Baldridge et al (1973) concluded that the cyclic eruptions of basaltic – andesitic – ice- landitic lavas (SiO2 > 60 wt%) of Hekla volcano (Iceland) are the result of time-dependent fractionation in a sub-surface magma chamber between the eruptions. Similar conclusions were drawn from sequences of basalt – ferrobasalt – ferroandesite to rhyodacite from the Galapagos spreading center (Byerly, 1980). An alternative model suggests partial melting of hydrous altered oceanic crustal rocks, which could be a more important formation process of oceanic felsic rocks especially of plagiogranites. Nicolas and Boudier (1991) proposed that a hydrous partial melt could be produced by a water ingression into young, still hot gabbros. Geochemical studies to date mainly aimed on trace element and isotope studies of the evolved rocks (e.g. Gillis and Coogan, 2002; Stakes and Taylor, 2003; Beard et al. 2005; Wanless et al 2010, 2011; Brophy and Pu 2012, Grimes et al 2013) and suggest that fractional crystallization alone can rarely be the exclusive formation process. O´Nions and Gr¨onvold (1973) discovered Icelandic rhyolites with different initial Sr isotope ratios where several samples have similar Sr isotope ratios to their host basalts indicating a formation by differentiation. Other rhyolites show slightly more radiogenic Sr possibly indicating formation by partial melting of gabbros (O´Nions and Gr¨onvold, 1973).

It is still a matter of active debate whether partial melting of hydrothermally altered crustal rock alone or assimilation-fractional crystallization is the more important process during formation of felsic rocks within the oceanic crust. Gillis and Coogan (2002) described a magma chamber with the tendency of up- and downward movements triggered by melt supply and hydrothermal systems, most likely promoting anatexis of sheeted dikes, resulting in SiO2-rich leucosomes. Some experimental studies characterize the composition and mineral content of the suggested protoliths after a partial melting event (e.g. Koepke et al. 2004, 2007; France et al. 2010). Such protoliths were found in the oceanic crust and described as “granoblastic dikes” (equatorial Pacific: Wilson et al. 2006; Koepke et al. 2008) or “hornfelsic” rocks (Troodos: Gillis and Roberts, 1999; Gillis and Coogan, 2002) and comprise recrystallized gabbros and sheeted dikes. The abundance of such rocks suggests that partial melting events are highly common in oceanic crust and therefore may be an important mechanism for felsic rock generation. 1. Introduction 27

Important tools for studying generation processes of felsic magmas

Whereas experimental studies aim on the reproduction of a natural rock by melting and crystallizing experiments under varying and previously well-defined pressure, temperature and oxygen fugacity conditions (e.g. Koepke et al. 2004, 2007), the geochemical approach considers the composition of the natural sample. Major element variations define typical trends in a rock suite belonging to one or more magmatic processes like the differentia- tion of magma as a result of fractional crystallization. Typical tholeiitic trends during fractionation of olivine, plagioclase and clinopyroxene are defined by an increase in T FeO , Na2O, TiO2 and decrease in Al2O3 and CaO with decreasing MgO (Klein and Langmuir, 1987; Klein et al. 1991), whereas later the residual melt compositions also T shows a decrease in FeO , Na2O, TiO2 during the fractionation of FeTi-oxides and more

Na-rich feldspars. Major and trace element variation diagrams (e.g. TiO2 vs. SiO2 or Ce vs. SiO2) may provide evidence whether a felsic rock is produced by fractional crystallization or by partial melting of oceanic crust (Koepke et al. 2007; Brophy and Pu, 2012). Furthermore, incompatible trace element ratios of a rock sample may provide information on the composition and mineral content of its source, the melting degree of the source and crystallization processes. The partition coefficient of an element for a mineral-melt pair determines incorporation into the mineral respectively the concentration remaining in the melt during crystallization processes as well as during the melting of a rock. Modeling calculations simulating melting and crystallization are commonly used tools for testing rock-forming hypotheses. However, isotope ratios are also a powerful tool to detect source heterogeneities or mixing of magmas from different sources or can indicate assimilation processes e.g. of sediments (e.g. Pb, Nd) or hydrothermally altered crustal rocks (e.g. O isotopes) (e.g. Cox et al. 1999; Eiler et al. 2007; Chauvel et al. 2009; Chekol et al. 2010).

In short, felsic volcanic and plutonic rocks occur within oceanic settings and their for- mation processes, relations to the mafic rocks and particular information content are important with respect to their tectonic setting, physical properties within the crust and contamination processes. Careful geochemical analyses and interpretations are necessary to understand and distinguish the individual processes.

In this study, I will present results of assimilation-fractional crystallization processes in a mid-ocean ridge setting (Pacific-Antarctic Rise) as well as evidence of different magmatic phases in the Troodos and Oman ophiolites involving changes in melt compositions, water content and source contamination. 2 Aims of the study

Oceanic plagiogranites comprise only small volumes of rocks but are common in ophiolites and have been neglected during most geochemical crustal rock studies. Therefore, the main objective of this thesis was to identify the processes that produce felsic magmas in oceanic settings and detect the relations of the plagiogranites with the mafic crustal rocks. Furthermore, the informational value of oceanic plagiogranites concerning crustal generation processes should be tested. With the aim to attain major comparability, this thesis comprises three different sample settings, one in recent oceanic crust and two in ancient and huge, well-preserved ophiolite complexes. In the following each location comprises a single self-contained chapter.

3) “Oxygen isotope evidence for the formation of andesitic-dacitic magmas from the fast-spreading Pacific-Antarctic Rise by assimilation-fractional crystallisation” The third chapter subject fresh, unaltered, glassy lavas from basaltic to dacitic composi- tion from the Pacific-Antarctic Rise. The samples were dredged during two cruises of RV SONNE in 2001 and 2010. The oxygen isotope and mineral analyses are part of this thesis.

4) “Constraints on the formation of geochemically variable plagiogranite intrusions in the Troodos Ophiolite, Cyprus” The fourth chapter presents new geochemical and perological data for tonalites, throndhjemites and some basalts, andesites and gabbros from several locations in the Troodos ophiolite complex. The samples were collected during a fieldtrip in 2010 (3 weeks). The preparation was performed during 2010-2011 and comprises major element, trace element and mineral chemistry analyses.

5) “Constraints on the evolution of oceanic plagiogranites in the Semail-ophiolite, Oman” The fifth chapter presents data on oceanic plagiogranites from the Oman (Semail) ophiolite complex. The samples are tonalites, trondhjemites, granodiorites some basalts and gabbros from 30 outcrops in eight different tectonic blocks sampled during a three-week fieldtrip in 2011. The sample preparation was carried out from 2011 to 2013 and comprises major element, trace element, isotope analyses and mineral chemistry. 2. Aims of the study 29

Combining the results of the three chapters, there is not a single formation process of felsic melt within oceanic settings. Important processes are the differentiation of more primitive magmas and in most cases assimilation of hydrothermally altered oceanic crustal rocks. Careful approaches are necessary to detect distinct evidence. However, felsic rocks and oceanic plagiogranites are common rocks in recent oceanic crust as well as in ophiolites and require more attention within general models of oceanic crust formation. 3 Oxygen isotope evidence for the formation of andesitic-dacitic magmas from the fast-spreading Pacific-Antarctic Rise by assimilation-fractional crystallisation

Sarah Freund, Christoph Beier, Stefan Krumm, Karsten M. Haase GeoZentrum Nordbayern, Universit¨at Erlangen-Nurnb¨ erg, Schlossgarten 5, 91054 Erlangen, Germany

3.1 Abstract

Andesitic to dacitic lavas occur along a 300 km long portion of the Pacific-Antarctic Rise close to the intersection of the spreading axis with the Foundation seamount chain. The fresh silicic glasses have low δ18O isotope values between 5.6‰ and 5.1‰ whereas basaltic glasses from the same ridge section have normal MORB δ18O values. Additionally, two

FeTi basaltic and all silicic glasses have high Cl (up to 1.1 wt.%) and K2O (up to 1.6 wt.%) contents, indicating assimilation of hydrothermally altered material. Modelling suggests that the fractionating magma assimilated up to 30% of hydrothermally altered material after 57% fractional crystallisation of the basaltic magma in a melt lens at less than 2 km depth. In contrast, the basaltic glasses show little assimilation and clinopyroxene-melt barometry indicates crystal fractionation in deeper melt sills. Relatively low H2O/Ce and Li/Ce ratios as well as the low δ18O values and high Cl and K concentrations in the silicic glasses suggest assimilation of altered crustal rock rather than a brine. While some of the variability in highly incompatible element ratios is best explained by crystal fractionation processes of FeTi-oxides (e.g., decreasing Nb/U) others require a reaction of the melt with residual amphibole and clinopyroxene (Cl/K, Tb/Yb, Hf/Sm, Ce/Yb). The initial onset of FeTi oxide crystallisation is associated with a reduced oxygen fugacity causing sulfide saturation and significant loss of S, Cu, and Co from the evolved melts. This change to more reducing melts as a result of oxide crystallisation is supported by the strong development of a negative Eu anomaly in the andesites and dacites indicating stronger partitioning of Eu into plagioclase. Reverse mineral zoning in the evolved lavas also indicates that replenishment by mafic melts occurs in the shallow melt lens. 3.2. Introduction 31

3.2 Introduction

Mid-ocean ridges typically erupt basaltic lavas that formed by partial melting of depleted peridotites and pyroxenites in the upper mantle and were affected by relatively small amounts of fractional crystallization during ascent (Langmuir et al., 1992). However, rare but significant amounts of felsic magmas with >57 wt.% SiO2 are known from different oceanic settings including mid-ocean ridges (Byerly et al., 1976; Geist et al., 1995; Gunnarsson et al., 1998; Haase et al., 2005; Perfit et al., 1999; Regelous et al., 1999; Smith et al., 2003; Wanless et al., 2010; Wanless et al., 2011). Evolved silicic lavas in the oceanic crust are important in order to understand the generation of the earliest continental crust on earth (Rudnick et al., 2003; Taylor and McLennan, 1985) that must have formed by transformation from either mafic magmas or mafic rocks to more evolved magmas, i.e. by fractional crystallisation or partial melting and assimilation processes or a combination of these processes (Foley et al., 2002; Rollinson, 2008; Smithies et al., 2009).

Although magmas with more than 57 wt.% SiO2 can form by partial melting of hy- drothermally metamorphosed oceanic crust rocks (France et al., 2010; Koepke et al., 2007; O’Nions and Gr¨onvold, 1973), several recent geochemical and experimental studies have shown that extreme crystal fractionation from basaltic liquids appears to be the prevailing process (Berndt et al., 2005; Grove and Juster, 1989; Haase et al., 2005; Wanless et al., 2010). However, pure crystal fractionation of basaltic to andesitic/dacitic melt leads to significantly higher δ18O values in the silicic lavas relative to the basalts because the fractionating minerals have lower δ18O (Bindeman, 2008; Bindeman et al., 2004; Cooper et al., 2004; Eiler, 2001; Muehlenbachs, 1982). Fractional crystallisation processes of magmas in the oceanic crust are often associated with assimilation of hydrothermally altered material and O isotopes can help to understand these processes because the lower sheeted dykes and the uppermost gabbros show δ18O values as low as 3‰ as a result of high temperature hydrothermal alteration (Alt and Teagle, 2000; Alt, 2003; Harper et al., 1988; Ito and Clayton, 1983; Muehlenbachs and Clayton, 1972; Stakes and Vanko, 1986). Latent heating during crystal fractionation of magmas may lead to amphibole breakdown in the hydrothermally altered wall rocks and release of volatiles (Wanless et al., 2010; Wanless et al., 2011). The incorporation of wall rock by either partial melting or assimilation can greatly modify the composition and fractionation pathway of the initial melt and may produce a wide range of rock types (Haase et al., 2005).Thus, understanding the mechanisms of assimilation-fractional crystallisation (AFC) in the oceanic crust is essential for our understanding of the formation of evolved melts in general.

Here, we present new O isotope, geochemical and petrological data for basaltic and andesitic to dacitic glasses from the Pacific-Antarctic Rise near the Foundation Hotspot chain. We extend the AFC model of Haase et al. (2005) and show that the formation of silicic lava occurs in shallow magma chambers and is coupled with the formation of 3.4. Results 32

FeTi-rich basaltic melt. As a consequence of shallow storage the evolved melts indicate sig- nificant assimilation of hydrothermally altered material. In contrast, the basaltic magmas stagnated deeper in the crust (>3.4 1 kbar) and show no evidence for assimilation. The abundance of intermediate to evolved lavas on the Pacific-Antarctic ridge in the vicinity of the Foundation Hotspot may be the result of an increased magma budget and crustal thickness.

3.3 Geological Background

The study area lies at the northern Pacific-Antarctic Rise (PAR), a fast-spreading mid- ocean ridge (MOR) with spreading rates of up to 10 cm/yr at 35◦S and an axial ridge width of 2 to 5 km. In the north the studied PAR segment is bounded by a large overlapping spreading center at 36.5◦S but the neovolcanic zone south of 36.5◦S is linear without significant offsets and deviations (Fig. 3.1). The summit trough is approximately 200 m wide with abundant signs of recent volcanic and hydrothermal activity (Hekinian et al., 1997; Lonsdale, 1994). The time-progressive Foundation seamount chain intersects the PAR between 37 and 38◦S, probably reflecting the interaction of the PAR with a deep mantle plume (Maia et al., 2001; O’Connor et al., 1998). The PAR lavas close to the Foundation seamounts are slightly enriched in incompatible elements but towards the south, the compositions progressively change to depleted MORB with distance from the intersection with the hotspot (Hekinian et al., 1999; Hekinian et al., 1997). This geochemical gradient is believed to reflect the influx of enriched plume material from the Foundation deep mantle plume into the melting zone underneath the PAR. Basaltic as well as andesitic to dacitic glasses were recovered between 36.5◦ and 39.5◦S in water depths between 2190 m and 2250 m (Fig. 3.1) and the evolved melts were shown to have formed by AFC processes in shallow magma lenses (Haase et al., 2005). Both evolved and basaltic samples were recovered within individual dredges, implying close spatial proximity of eruptive vents. The fast-spreading rate of 10 cm/yr requires time intervals of less than 2000 years between mafic and evolved eruptions (Haase et al., 2005). We discriminate the samples in terms of decreasing Foundation plume influence as described in Haase et al. ◦ ◦ (2005), in northern (= 36.5-38.05 S) and southern (= 38.05-40 S) and the SiO2 content (basaltic, andesitic, dacitic) (Fig. 3.2).

3.4 Results

3.4.1 Chemical and O isotopic compositions of the PAR glasses

Major elements

Analytical methods are available online (supplements: analytical methods).

Although the glasses come from a large region of the PAR the major elements (Table 3.4. Results 33

Figure 3.1: Bathymetric map of the Pacific-Antarctic Rise between 36◦S und 40◦S and the ridge locations of the dredged samples. The time-progressive Foundation seamount chain intersects the PAR between 37 and 38◦S. Silicic glasses are marked as circles, basaltic glasses as squares. Using GMT (Wessel and Smith, 1991). 3.4. Results 34

Table 3.1: Representative major and trace elements and isotopic compositions of silicic and basaltic glasses at the Pacific–Antarctic Rise (oxides in wt.%, elements in ppm).

Northern andesites SO100-105DS1 SO157-6DS1 SO157-6DS2 SO157-5DS1 SO157-15DS1 SO157-3DS1 SO157-18DS1 SO157-30GTV1 SO157-32DS1 Lat [S] 37.191 37.562 37.562 37.606 37.618 37.659 37.683 37.791 38.032 Long [E] −110.707 −110.826 −110.826 −110.856 −110.863 −110.872 −110.883 −110.914 −110.980 SiO2 55.5 57.7 57.2 56.4 58.9 61.9 59.0 52.5 57.3 TiO2 2.09 1.55 1.57 1.60 1.42 1.02 1.25 2.29 1.49 Al2O3 12.2 12.7 12.7 13.9 13.7 13.5 13.6 13.5 13.5 FeOT 15.2 12.0 12.0 10.5 9.79 7.97 8.86 13.1 9.75 MnO 0.265 0.207 0.199 0.166 0.165 0.123 0.135 0.221 0.156 MgO 2.03 1.90 2.01 3.12 1.99 1.57 1.87 3.96 2.87 CaO 6.50 5.64 5.71 6.77 5.45 4.47 5.09 8.36 6.46 Na2O 3.35 4.40 4.66 3.98 4.59 4.49 4.63 3.67 4.24 K2O 0.633 0.738 0.736 0.787 1.01 1.27 1.06 0.419 0.786 P2O5 0.813 0.555 0.540 0.384 0.484 0.294 0.377 0.605 0.363 SO2 0.164 0.170 0.161 0.151 0.086 0.060 0.080 0.218 0.125 Cl 0.151 1.15 1.14 0.577 0.972 0.963 0.868 0.128 0.552 H2O 0.510 1.50 0.760 Total 99.02 98.77 98.60 98.43 98.52 97.73 96.86 99.02 99.00 Li 16.8 14.2 14.2 16.4 18.5 18.3 11.4 14.9 Sc 16.2 22.2 17.9 19.8 15.1 11.8 37.1 26.0 Co 20.1 19.9 25.7 18.7 14.5 17.7 29.7 23.1 Ni 1.48 8.33 18.9 6.62 6.05 8.16 24.8 14.7 Cu 16.0 34.0 43.9 35.4 32.9 36.9 37.7 31.6 Rb 10.1 10.8 11.5 16.6 22.1 18.5 6.60 12.5 Sr 144 134 126 117 107 119 150 124 Y 97.6 114 74.0 105 109 102 79.4 94.0 Zr 440 569 498 654 714 765 304 530 Nb 26.8 36.5 23.1 33.7 33.4 33.7 19.0 24.1 Cs 0.107 0.109 0.120 0.177 0.231 0.186 0.071 0.124 Ba 113 129 111 151 179 160 82.0 117 La 25.8 33.4 26.3 37.1 41.3 39.5 17.5 28.0 Ce 64.9 83.0 62.2 89.7 96.8 91.9 46.3 68.6 Nd 48.5 54.7 39.4 54.1 55.2 56.1 33.7 43.6 Sm 14.2 15.2 10.7 14.3 14.2 14.7 10.3 12.2 Eu 4.11 3.60 2.37 3.00 2.66 3.02 3.01 2.67 Tm 1.72 1.85 1.34 1.74 1.80 1.86 1.28 1.56 Gd 17.2 17.3 12.3 15.8 15.4 16.7 12.4 13.9 Dy 19.5 20.3 14.3 18.4 18.4 19.3 14.4 16.6 Er 11.8 12.5 8.86 11.5 11.8 12.3 8.69 10.4 Yb 11.5 12.4 9.04 11.6 12.2 12.6 8.40 10.5 Lu 1.68 1.83 1.34 1.70 1.80 1.87 1.23 1.54 Hf 10.8 14.8 12.1 16.9 18.9 18.3 8.04 14.1 Ta 1.64 2.26 1.42 2.16 2.23 2.07 1.26 1.60 Pb 1.88 1.87 1.74 2.25 2.69 2.39 1.30 1.72 Th 2.26 3.61 3.32 4.74 6.23 5.51 1.43 3.48 U 0.697 1.11 1.04 1.48 1.91 1.70 0.480 1.12 δ18O 5.19 5.08 5.44 5.5 5.43 5.59 5.51 143Nd/144Nd 0.513007 0.513008 0.51300 0.51302 0.51303 3.4. Results 35

Table 3.1: continued.

Southern andesites Dacites Northern basalts Southern basalts SO157-65DS2 SO157-63DS1 SO157-65DS1 SO157-17DS1 SO157-24DS1 SO157-36DS1 SO157-44DS1 SO157-45DS1 Lat [S] 39.505 39.804 39.505 37.667 37.903 38.215 38.493 38.585 Long [E] −111.343 −111.427 −111.343 −110.877 −110.947 −111.050 −111.103 −111.131 SiO2 58.3 57.2 68.0 50.4 49.8 49.4 49.7 50.0 TiO2 1.75 1.64 0.604 1.78 1.52 1.64 1.46 1.62 Al2O3 12.5 13.1 12.4 14.3 14.7 14.7 15.2 14.7 FeOT 12.1 11.9 6.14 11.2 10.2 10.5 9.73 10.5 MnO 0.218 0.191 0.109 0.191 0.134 0.168 0.154 0.203 MgO 2.00 1.99 0.499 6.71 7.57 7.55 7.92 7.72 CaO 5.71 5.78 2.67 11.4 12.0 11.8 12.4 12.0 Na2O 4.36 4.31 4.70 2.64 2.56 2.73 2.79 2.71 K2O 0.711 0.654 1.35 0.194 0.151 0.151 0.059 0.138 P2O5 0.565 0.778 0.130 0.286 0.252 0.266 0.213 0.240 SO2 0.159 0.126 0.025 0.198 0.178 0.216 0.186 0.210 Cl 0.530 0.659 0.869 0.030 0.017 0.019 0.004 0.014 H2O 2.6 1.78 0.300 0.410 0.350 0.300 0.400 Total 98.96 98.30 97.47 99.42 99.14 99.22 99.82 100.08 Li 15.4 19.8 24.8 5.26 5.37 5.93 5.94 4.97 Sc 24.0 22.5 9.91 34.3 29.1 29.0 48.3 35.3 Co 20.2 16.3 5.72 42.3 39.5 39.5 39.4 41.1 Ni 5.44 4.57 1.56 42.8 60.5 64.5 57.2 73.0 Cu 22.2 26.2 10.7 74.2 65.2 60.1 63.6 64.9 Rb 8.87 7.17 18.0 3.04 2.00 1.88 1.68 1.79 Sr 117 118 76.5 145 144 147 141 140 Y 125 152 162 32.6 26.6 29.4 28.1 32.5 Zr 630 762 850 116 86.0 96.6 96.7 109 Nb 18.1 19.8 23.2 7.70 5.91 5.52 4.89 4.43 Cs 0.118 0.096 0.171 0.081 0.015 0.014 0.018 0.048 Ba 85.1 70.4 133 39.0 29.3 26.3 24.5 22.7 La 25.9 26.8 39.6 7.17 5.52 5.49 4.89 5.03 Ce 71.1 78.4 105 18.6 14.2 14.5 13.6 14.4 Nd 52.2 62.6 67.9 13.7 11.3 12.1 10.9 12.0 Sm 15.5 19.1 19.4 4.16 3.57 3.92 3.59 3.92 Eu 3.70 4.58 3.49 1.43 1.26 1.39 1.29 1.36 Tm 1.98 15.9 2.73 0.508 0.466 0.518 0.472 0.517 Gd 18.0 22.3 21.8 5.01 4.54 5.02 4.50 4.97 Dy 21.5 26.3 27.4 5.94 5.31 5.92 5.37 6.00 Er 13.2 15.9 17.8 3.48 3.24 3.59 3.24 3.55 Yb 13.2 15.7 18.3 3.38 3.08 3.42 3.07 3.45 Lu 1.94 2.29 2.68 0.503 0.464 0.508 0.452 0.512 Hf 16.2 18.7 24.7 3.16 2.61 2.79 2.62 3.06 Ta 1.19 1.24 1.60 0.521 0.383 0.358 0.325 0.308 Pb 1.82 1.94 3.07 0.701 0.450 0.467 0.435 0.566 Th 2.38 1.76 4.70 0.598 0.428 0.396 0.328 0.320 U 0.859 0.668 1.65 0.187 0.145 0.136 0.114 0.115 δ18O 5.13 5.43 5.56 5.62 5.51 5.67 5.57 5.63 143Nd/144Nd 0.5131 0.5131 0.51309 0.51302 0.51302 0.51306 0.51307 0.51309 3.4. Results 36

0.51315 Foundation plume - ridge intersection

0.51310

d Figure 3.2: Neodymium isotopes N

144 and Nb/Zr ratio of

/ 0.51305 d the PAR glasses versus

N Northern andesites Southern andesites Latitude. The PAR 143 Northern basalts 0.51300 glasses show a significant Southern basalts range induced by the Dacites A Foundation mantle plume 0.51295 intersection between 37 and 38◦S. Silica rich glasses occur about the 0.08 entire length of the ridge section. a) Nd isotopes and

r 0.06

Z b) Nb/Zr, / b indicating that the N samples are not 0.04 comagmatic but the influence of the Foundation plume - Foundation mantle B ridge intersection 0.02 plume decrease toward 36.5 37.0 37.5 38.0 38.5 39.0 39.5 40.0 the south.

3.1, supplementary Table 1) SiO2, CaO, and Al2O3 show relatively well-defined trends with MgO contents ranging from basalt via FeTi-enriched basalt and andesite to dacite in composition, comparable with published MOR lavas e.g. from the northern East Pacific Rise (EPR) (Fig. 3.3). However, between about 3 and 4 wt.% MgO we find two different T lava types where one is typical FeTi-basaltic with TiO2 of 3.5 wt.% and FeO of 16.5 T wt.% and the other andesitic, i.e. with higher SiO2 and K2O but lower FeO , TiO2 than the FeTi basalts. Calcium constantly decreases with decreasing MgO and Al2O3 shows a shallower decreasing slope for most glasses with less than 4 wt.% MgO and a steeper one for the basalts. As mentioned above the K2O contents in a group of andesites with about 3 to 4 wt.% is significantly higher than that of the FeTi basalts and other primitive andesites but all of these K-rich andesites are from the PAR closest to the Foundation seamount chain, i.e. between 37.1◦S and 38.2◦S.

Trace element variations in the PAR glassses

Both the compatible and incompatible trace elements (Table 3.1 and supplementary Table 1) vary considerably with decreasing MgO contents (Fig. 3.4). For example, Ni, Cu, and Cr (not shown) decrease monotonously with decreasing MgO but the Co contents remain constant in the basalts and FeTi basalts but are much lower and show a decreasing trend in the andesites. Similarly, Sr concentrations are constant in the basaltic glasses and decrease in the silicic glasses (Fig. 3.4D). The incompatible elements like La or Zr increase with 3.4. Results 37

75 20 A D Fractional 70 crystallisation

15 F ]

10% e O % 65 20% t. AFC T [ w w [ 60 10

2 30% mixing t. % O mixing 55 AFC ] Si 20% 10% 5 50 Fractional crystallisation 45 0 5 15 Fractional crystallisation B E 4 10% C

] Fractional crystallisation

10 a % 3 20% O t. AFC 30% AFC [ w 10% w [

30% t. 2 2 20% mixing % O

i 5 ] T 1 mixing Northern andesites Southern andesites 0 Northern basalts 0 16 Southern basalts 2.5 Dacites 30% mixing F 10% MOR lavas published 14 2.0 K

] 20% AFC 2 O % 12 t. 1.5 [ w w AFC [ 10 t. 3

30% %

O 1.0

Fractional crystallisation ] 2 8 mixing

Al 20% 6 10% 0.5 C Fractional crystallisation 4 0.0 0 2 4 6 8 0 2 4 6 8 MgO [wt.%] MgO [wt.%]

Figure 3.3: Major elements of the PAR glasses. Volcanic glasses from the PAR, glasses from the EPR (Wanless et al., 2010) and GSC (Byerly et al., 1976; Perfit et al., 1999) and experimental produced melts (partial melting of hydrothermal altered sheeted dykes at different temperatures) (France et al., 2010) , furthermore assimilation-fractional crystallisation (AFC), fractional crystallisation (FC) (both modelled with MELTS (Ghiorso and Sack, 1995)) and mixing trends of the PAR lavas. For details of model calculations see main text and Table 2a. 3.4. Results 38

Table 3.2: a) Details of the AFC calculations of major and trace elements of the PAR glasses.

MELTS options (Ghiorso and Sack, 1995) Intensive variables T (◦C) 1300-700 Increments (◦C) 5 Liquidus T ◦C (magma) 1195 Pressure (bars) 500 Proportion of assimilant (wt %) 100 Mass assimilant (g) 30 Start of assimilation T (◦C) 1200 Increments (◦C) 20

Composition assimilant Koepke et al. (2008) 309-1256D-80R-2, 92-102 SiO2 51.9 TiO2 1.86 Al2O3 14.9 FeOT 12.3 Mn 0.23 MgO 5.49 CaO 10.1 Na2O 2.97 K2O 1.12 H2O 1.27

Trace element calculation: parental magma: SO157 44DS1 After DePaolo (1981), MgO Mineral Crystallisation content of the remaining melt assemblage of assemblage (result of decrease (7 steps) during the assimilant MELTS AFC calculation) calculation, corresponding mineral assemblage taken from MELTS (AFC calculation) Amphibole 30% Cpx 25% 44.50% Plagioclase 41% 47.50% Magnetite 4% Olivine 2% Ilmenite 5.80% 3.4. Results 39

Table 3.2: b) Parameter of trace element and isotope modelling using EC-RAX FC.

EC-RAX FC options (Bohrson and Spera, 2007) Thermal parameter Magma Liquid T magma (◦C) 1195 Initial T magma (◦C) 1300 Lliquid T assimilant (◦C) 1119 Initial T assimilant (◦C) 700 Solidus T (◦C) 800 Magma isobaric specific heat capacity 1447.1 Assimilant isobaric specific heat capacity 1601.5 Recharge Parameters 0.3 Episodic 926-1050 ◦C Normalized fraction of Wallrock 0.3 Equilibration Temperatur Teq & 918.28 °C Mass of wall rock 3.013

Trace element Parameters Parental magma Element Cl (ppm) K (ppm) 0 Concentration in magma (Cm) 255 1200 Bulk in magma (Dm) 0.003 0.078 Enthalpy in magma 0 0 0 Concentration in assimilant (Ca) 8500 2698 Bulk in assimilant (Da) 0.152 0.384 Enthalpy in asssimilant 0 0 Concentration in recharge 100 1400 Bulk in recharge 0.003 0.078 Enthalpy in recharge 0 0

Oxygen isotope parameters ‰ 18O/16O of magma 5.65 18O/16O of assimilant 3.8 18O/16O of recharge 5.65 3.4. Results 40

50 A D

40 150 Sr m] 30 [ pp pp m] [ o 20 Northern andesites 100

C Southern andesites Northern basalts 10 Southern basalts Dacites MOR lavas published 0 50 80 B 1,000

60 800 Z

600 r m] [

40 pp pp m] [

i 400 N 20 200 E 0 0 50 C F 80 40 60 La m] 30 [ pp pp [ m]

u 40 20 C

20 10

0 0 0 2 4 6 8 0 2 4 6 8 MgO [wt.%] MgO [wt.%]

Figure 3.4: Trace element contents of the PAR. The PAR glasses showing very different behaviour of the different elements. Whereas Ni and Cu behave compatibly and decrease in the basalts, Sr and Co are constant up to 4 wt.% MgO and then decreases. Lanthanum behaves incompatible and increases with decreasing MgO. Zirconium increases less steeply in the basaltic glasses than in the silicic glasses. Published glasses from the EPR (Wanless et al., 2010) resemble the PAR lava trends well for Sr and Zr but have slightly lower La for a given MgO (D, E and F). 3.4. Results 41

2,000 A

1,500 m]

pp 1,000 [ S

500

Figure 3.5: Sulphur and Cl contents of 0 1.5 the PAR and EPR (Wanless B et al., 2011) glasses. A) Sulphur concentrations Northern andesites are significantly lower in Southern andesites 1.0 glasses with less than 2 wt.%

] Northern basalts

% Southern basalts MgO than in more primitive t.

w Dacites samples. [ l MOR lavas published C B) The silicic glasses from 0.5 the PAR show much higher Cl concentrations than the basaltic glasses and also higher than the EPR evolved 0.0 glasses. 0 2 4 6 8 MgO [wt.%] decreasing MgO and show an enrichment by a factor of nine between the primitive basalts and the most evolved dacite. The incompatible element trends show a steepening of the slope at 4 wt.% MgO (Figs. 3.4E, F).

The S contents in the basaltic glasses vary considerably between about 1600 and 700 ppm but appear to be constant between 8 and 4 wt.% MgO (Fig. 3.5A). In contrast, the andesites and dacite show significantly lower S concentrations and most samples indicate a broad decreasing trend from about 1100 ppm S at 4 wt.% MgO to 100 ppm in dacites. Importantly, Cl contents are low in all basaltic and FeTi basaltic glasses (>0.12 wt.%) whereas the evolved glasses with the exception of two samples have high Cl concentrations (0.5 – 1.2 wt.%) (Fig. 3.5B). Interestingly, the dacites have lower Cl concentrations (0.8- 0.9 wt.%) than some andesites, suggesting non-linear enrichment processes.

The incompatible trace element ratios show significantly different behaviour of incom- patible elements during the evolution of the melts from 8 to 0.5 wt.% MgO (Fig. 3.6).

Primitive mantle normalized (Ce/Yb)N are slightly higher (>1.3) in basalts adjacent to the Foundation Seamounts (37-38.0◦S) than in PAR MORB south of 38◦S (Fig. 3.6A) and a similar distinction is obvious in the evolved glasses. The silicic glasses show increasing

(Ce/Yb)N with the northern andesites are more enriched than the southern glasses. In particular andesites with ∼2 wt.% MgO have a wide range (1.3-2) and the dacite have a rather lower value (1.5). Whereas only a slight negative Eu anomaly Eu/Eu∗ =

(Eusample/EuChondrite) / (((Smsample/SmChondrite) + (Gdsample/GdChondrite))/2) develops 3.4. Results 42

2.5 A E 1.6 2.0 1.4 N H )

AFC f/

mixing 1.2 Sm Yb

/ 1.5 e AFC 1.0 (C 1.0 Northern andesites Southern andesites 0.8 Northern basalts Fractional crystallisation 0.5 Southern basalts 0.6 1.0 Dacites 1.4 MOR lavas published

AFC 1.2 (T * 0.8 b / Yb

Eu AFC / 1.0

mixing ) N Eu 0.6 mixing Fractional crystallisation 0.8 B F 0.4 0.6 50 15 C 40 AFC Fractional crystallisation 10 30 Sr/ Pb N / AFC e 20 d

C 5 10 mixing G 0 0 60 2.0 D H 50 Fractional crystallisation 1.5 C U l / 40 / K b AFC 1.0 N 30 AFC mixing 0.5 20 mixing 10 0.0 0 2 4 6 8 0 2 4 6 8 MgO [wt.%] MgO [wt.%]

Figure 3.6: Trace element ratios of the PAR and EPR (Wanless et al., 2011) glasses. Calculated mixing between the most primitive and the most evolved sample, a calculated fractional crystallisation trend and a calculated assimilation-fractional crystallisation trend are shown. The calculated trends suggest that, fractional crystallisation alone could not generate the trace element pattern of the evolved lavas. Mixing between two different end members is also unlikely. 3.4. Results 43 between 8 and 4 wt.% MgO, the Eu/Eu* decreases much steeper at MgO lower than 4 wt.% (Fig. 3.6B) indicating increasing fractionation of plagioclase.

Some incompatible element ratios like Ce/Pb, Nb/U and Hf/Sm were shown to be constant in mafic rocks from the oceans (e.g. Hofmann et al. (1986)). However, we find that the evolved lavas show significantly different ratios than the PAR MORB (Figs. 3.6C, D, E). For example, Hf/Sm is nearly constant in the basaltic glasses but andesites with about 2 wt.% MgO show a wide range of 0.8-1.3, whereas some andesites with higher MgO content have comparatively high Hf/Sm of 1.1-1.2. In terms of the Ce/Pb ratios the PAR basalts (20-33) generally lie within the normal MORB range of 25 5 (Hofmann et al., 1986) whereas the silicic glasses have increasing Ce/Pb ratios up to 44. Nb/U and

(Te/Yb)N generally decrease with decreasing MgO in the silicic glasses (Fig. 3.6F). All basalts have constant (Te/Yb)N ratios of 1.1-1.2 similar to the FeTi basalts. The Nb/U ratios of the PAR basalts and FeTi basalts are constant within the range of global MORB

(Hofmann et al., 1986). In contrast, the andesites show variable Nb/U and (Te/Yb)N and the dacites have the lowest of Nb/U and (Te/Yb)N ratios. Both basalt types, the FeTi basaltic glasses and the silicic glasses have decreasing Sr/Nd with decreasing MgO (Fig. 3.6G).

The highly incompatible element ratio Cl/K increases slightly with decreasing MgO in basaltic glasses similar to EPR lavas whereas the silicic glasses generally have much higher Cl/K (>0.6) (Fig. 3.6H). The dacites have comparatively low ratios (0.73 and 0.77) similar to the silicic lavas from the EPR. Figure 3.6 also shows calculated fractional crystallisation trends (Rayleigh fractionation), potential mixing trajectories between primitive basaltic and evolved dacitic magma and AFC trends (calculated). H2O/Ce and Li/Ce ratios generally appear to decrease with decreasing MgO in the basaltic to dacitic glasses (Figs. 3.7A, B).

Oxygen isotope compositions of the PAR glasses

Unaltered, fresh glasses of recent volcanic rocks can be used to directly infer primary magmatic δ18O values (Bindeman, 2008). The PAR basaltic glasses display a typical narrow MORB range in δ18O from 5.5‰ to 5.7‰ while the andesitic and dacitic glasses generally have lower than expected values between 5.6‰ and 5.1‰ with a mean of 5.4‰ (Fig. 3.7C). Pure crystal fractionation as well as low-temperature hydrothermal alteration of the glasses or contamination with sediment would produce distinctly higher δ18O values in the silicic lavas compared to the basaltic glasses (Bindeman, 2008). However, there is recently no known process to decrease the δ18O within a fractionating melt except contamination with low δ18O material. The decreasing δ18O in andesites and dacites from the PAR differs from increasing δ18O values observed in some evolved lavas of the Galapagos Spreading Center (GSC) (Byerly, 1980; Muehlenbachs, 1982). The PAR silicic glasses resemble the trend found in lavas from the eastern GSC (Michael and Cornell, 3.4. Results 44

400 A

300

e 18 C

/ 2 ‰ 200 Figure 3.7: H O/Ce, Li/Ce and δ O (VSMOW) O 2 of the PAR and published glasses from H 100 the EPR and GSC. A) H2O/Ce and B) Northern andesites Li/Ce in the PAR basaltic glasses remain Southern andesites 0 Northern basalts approximately constant but decrease in Southern basalts 2 Dacites B PAR evolved lavas, whereas H O/Ce and 0.6 MOR lavas published Li/Ce slightly increase in the EPR evolved lavas (Wanless et al., 2011) C) Oxygen

e 0.4 isotopes. The PAR basaltic glasses lie C /

i 18

L within the δ O range of fresh MORB 0.2 glasses (Eiler et al., 2000) whereas the evolved glasses have similar or lower δ18O ‰ 0.0 ratios (5.1 - 5.6 ) than the basalts and

] 7.5 thus significantly lower than expected for

W C silica rich lavas produced by fractional 7.0 crystallisation alone. The evolved samples VSMO [ 6.5 of the PAR and GSC show constant low

‰ 18 and increasing δ O values with decreas- O 6.0 18 ing MgO content (Byerly, 1980; Michael

5.5 MORB range and Cornell, 1998; Muehlenbachs, 1982; Perfit et al., 1999; Wanless et al., 2011). 5.0 18 0 2 4 6 8 The error bars for δ O (PAR) are smaller MgO [wt.%] than symbol size.

1998; Perfit et al., 1999) and the dacitic glasses from the EPR (Wanless et al., 2011) in terms of δ18O. However, the evolved PAR lavas have the lowest δ18O values measured within high-silica glasses from a MOR setting.

3.4.2 Petrography and mineral chemistry

The basaltic and silicic lavas contain few crystals in a glassy matrix only few cm below the pure glassy rim.

Mineral compositions in the basaltic lavas

Phenocrysts in the basaltic lavas (>1 mm) are generally much larger compared to phe- nocrysts in the evolved lavas. Basaltic lavas contain clinopyroxene and plagioclase, but olivine (Fo80) phenocrysts are rare and small (>0.2 mm). The clinopyroxene and plagio- clase grains (average size of 0.4 mm) were measured from core toward the rim (parallel to the optic zoning. The plagioclases are normal Ca-rich (average core composition: An68) and the clinopyroxenes are augites (average En54 Fer12 Wo34). The mineral composition of the PAR basaltic lava is very similar to the minerals within the fractionated tholeiitic basalts from the EPR (Batiza et al., 1977). 3.5. Discussion 45

Mineral compositions of the silicic lavas

The minerals occurring in the glassy to cryptocrystalline matrix are plagioclase and clinopyroxene (maximum size of 0.2 mm) and accessory oxides (>10 µm). Clinopyroxene is mostly euhedral-subhedral and is optically and chemically zoned, frequently displaying a characteristic hourglass texture. A few augites show sieve textures and are inversely zoned, i.e. their Mg# increases from core to rim. These sieve-textured and inversely zoned clinopyroxenes are larger than the normally zoned clinopyroxenes (0.35 – 0.4 mm). The evolved sample SO157 6DS2 (andesite) displays both, normally and inversely zoned augites (e.g. normally zoned: core Mg#79, rim Mg#63; inversely zoned: core Mg#53 versus rim Mg#69). Additionally, a few normally zoned clinopyroxenes show a steep decrease in Mg# from core to rim (e.g. core Mg#81, outer rim Mg#49). Plagioclase occurs as euhedral-subhedral laths and displays a compositional range from labradorite to oligoclase with a normal zoning, (average of various plagioclase measurements of 5 samples (andesites) core: An28−43; rim: An28−43). The compositions of the clinopyroxene and plagioclase crystals of the evolved lavas overlap with those known from ferrogabbros in the southeast Pacific (Constantin et al., 1996; Natland and Dick, 1996).

3.5 Discussion

3.5.1 Liquid lines of descent and fractional crystallisation processes

The PAR lavas show a significant range of Nd isotopes, Nb/Zr and (Ce/Yb)N ratios with the lavas closest to the Foundation Seamount being most enriched in terms of incompatible elements (Fig. 3.2, 3.6). This systematic variation along the PAR reflects the inflow of enriched material from the Foundation plume (Hekinian et al., 1999; Hekinian et al., 1997; Maia et al., 2001). However, although the samples along the 300 km long portion of the PAR have different mantle sources, we find that basaltic and evolved lavas occurring in close proximity have similar sources as reflected by similar Nd isotope ratios. The glasses lie along relatively tight continuous trends in the major elements indicating the processes during melt formation and ascent are largely similar and resemble the experimentally produced crystal fractionation trend of Juster et al (1989). Glasses with 2.5 to 6 wt.% MgO appear to be rare and we only recovered two glasses with FeTi-basaltic composition (Fig. 3.3B, D). In order to test whether solely crystal fractionation processes can generate the silicic glasses we used the MELTS algorithm (Ghiorso and Sack, 1995) starting with a representative, primitive basaltic sample (SO157 17DS1, Table 3.1). The fractional crystallisation model yields a reasonable fit to the basaltic glasses up to the FeTi-rich basalts (samples SO100 106DS2 and SO100 106DS3) and two of the more FeTi-rich northern andesitic glasses (SO157 30GTV1 and SO100 105DS1) with the notable exception of Al2O3 that is too low in the model compared to the glass compositions (Fig. 3.3C). We suggest that the observed variation in the basaltic glasses between 8 and 4 wt.% MgO 3.5. Discussion 46 reflects fractional crystallization during stagnation in crustal magma chambers. Crystal fractionation can also explain the enrichment of incompatible elements like K2O in some evolved glasses but most show higher K contents. Even more extreme is the enrichment of Cl in the andesites and dacites and Cl increases much more than K as observed in the higher Cl/K ratios in the andesites compared to the basalts. Thus, the pure fractional crystallisation model fails to reproduce the more evolved compositions as well as the range in major and trace elements for andesitic glasses with less than 4 wt.% MgO. The most likely process leading to the extreme enrichment of Cl is the assimilation of wall rocks that were hydrothermally altered by seawater (Haase et al., 2005; Michael and Schilling, 1989). Consequently, we agree with previous studies (Haase et al., 2005) that the generation of the silicic PAR melts is a combined process involving assimilation and crystal fractionation of basaltic magmas derived from variably plume influenced mantle sources.

3.5.2 Constraints on assimilation processes in the generation of silicic PAR magmas

Assimilation of wall rocks that reacted with seawater is the most likely process leading to the high Cl contents in the evolved magmas (Fig. 3.5B) and Cl/K (Fig. 3.6H). Michael and Cornell (1998) propose the limit of Cl/K between >0.01-0.08 for pristine MORB and suggest that higher values in MOR lavas are related to assimilation. Furthermore, Haase et al. (2005) observe slightly higher Sr isotope ratios for a given 206Pb/204Pb of some of the evolved PAR lavas compared to the PAR basalts. The decoupling of the Sr isotope system from Pb in the silicic melts reflects assimilation of seawater Sr (seawater 87Sr/86Sr about 0.709, e.g. (Hess et al., 1986; Richter et al., 1992)) and the flux of hydrothermal fluids in the crust increases the Sr isotope ratio in the uppermost 1 km of the oceanic crust (Alt and Teagle, 2003; Alt et al., 1996). The δ18O of fresh MORB glasses range between 5.4‰ and 5.8‰ (Eiler et al., 2000) and the PAR basalt glasses lie within this range whereas the evolved glasses have similar or lower δ18O ratios (5.1 - 5.6‰) than the basalts and thus significantly lower than expected for silica rich lavas (Fig. 3.7C). In a fractionating melt, the δ18O increase from basalt to andesite because the fractionating minerals (olivine and pyroxene) have lower δ18O than the melt and thus drive the evolving melt to up to 0.4‰ higher δ18O values than in the basalts (Bindeman, 2008; Bindeman et al., 2004; Cooper et al., 2004; Eiler, 2001; Muehlenbachs, 1982). However, the lower sheeted dykes and the uppermost gabbros in the oceanic crust show δ18O values decreasing down to 3‰ as a result of high temperature (>200◦C) alteration (Alt and Teagle, 2000; Alt, 2003; Harper et al., 1988; Ito and Clayton, 1983; Muehlenbachs and Clayton, 1972; Stakes and Vanko, 1986). During the ascent the PAR magmas pass through this alteration zone and in combination with the other signs of assimilation (increased Cl contents and Sr isotopes) we propose that the low δ18O of the evolved magmas are due to assimilation of altered material with low δ18O. 3.5. Discussion 47

France et al. (2010) compare experimental melts produced by partial melting of hydrous gabbros with melts formed by partial melting of hydrothermally altered sheeted dykes. Partial melting of the dykes produces silicic melts under low pressure, water-saturated conditions and high f O2 (QFM1-QFM2). These melts are enriched in Si and K but slightly depleted in Mg and Ca (compared to melts formed during pure MORB frac- tionation) whereas hydrous partial melting of oceanic gabbros generates melts depleted in incompatible elements like Ti and K because of their mineralogy and composition. Furthermore, the temperatures for ≤ 30% partial melting of hydrothermally altered ◦ dykes did not exceed 940 C producing melts with SiO2 ≥ 68.5% (France et al., 2010). Consequently, the observed enrichment of K in the PAR silicic glasses relative to the basalts (Fig. 3.3F) cannot be due to assimilation of gabbros, but requires assimilation of relatively K-rich portions of the lower sheeted dyke complex. Hydrothermally altered rocks at 1000 m below seafloor in the lower sheeted dykes have high Cl contents (Sano et al., 2008), 87Sr/86Sr isotope ratios up to 0.706 and comparatively low δ18O (6 to 3‰) (Alt and Teagle, 2000; Alt, 2004; Alt and Teagle, 2003) and we suggest that such material is assimilated by the evolved PAR melts. We conclude that the chemical characteristics of the evolved PAR lavas suggest assimilation of seawater-altered material, most likely at the base of the sheeted dyke complex in a melt lens as described from plutonic rocks (Natland and Dick, 1996; Natland and Dick, 2009). In contrast, the basaltic glasses show no signs of interaction with the hydrothermally altered portion of the crust and thus resided deeper (detailed discussion in chapter 4.4) in the crust.

3.5.3 Evidence from mineral compositions for melt evolution and the link to oceanic ferrogabbros

Clinopyroxene phenocrysts in the basalts and many minerals in the evolved lavas are zoned (decreasing MgO# from mineral core toward the rim) in agreement with fractional crystallisation. The evolved sample SO157 6DS2 (Mg#glass 23) contains both, normally and inversely zoned clinopyroxenes with the inversely zoned augites being larger (0.35 – 0.4 mm) than the normally zoned clinopyroxenes (average 0.2 mm) and most of the inversely zoned clinopyroxenes show sieve textures combined with a euhedral habitus. Sieve- textured plagioclase and clinopyroxene indicate disequilibrium or large oversaturation and rapid crystal growth by extreme undercooling (Nelson and Montana, 1992; Shaw and Dingwell, 2008). The cores of the inversely zoned, sieve-textured clinopyroxenes (sample

SO157 6DS2) are in equilibrium with a melt having Mg#60−54, representing equilibrium crystallisation in a melt with 2.5 to 2 wt% MgO (Putirka et al., 1996) whereas the cores of the normally zoned clinopyroxene in sample SO157 6DS2 are in equilibrium with a melt with Mg#82−71, representing equilibrium with MgOmelt of 5.3 to 3.5 wt%. We suggest that the mineral zoning is most likely due to replenishment of the reservoir by mafic magma. One possible scenario is that minerals crystallised in equilibrium with an andesitic melt remain in the melt lens after an extrusion event and were mixed with more mafic melt as 3.5. Discussion 48 a result of magma recharge before it was incorporated again into the fractionating melt. Such frequent magma mixing events may be typical for the melt lens at fast-spreading ridges as outlined by Natland and Dick (1996; 2009).

The plagioclase and clinopyroxene crystals in the evolved melts from the PAR have similar compositions to those from ferrogabbros sampled from the EPR. These ferrogabbros are believed to indicate extreme fractional crystallisation in the shallow crust of fast-spreading ridges and such rocks appear to be the latest intrusions in one magmatic cycle (Constantin et al., 1996; Natland and Dick, 1996; Natland and Dick, 2009) and thus represent the plutonic counterparts of the evolved PAR lavas. However, whereas Constantin et al. (1996) suggested that evolved melts like FeTi basalts, andesites, and ferrogabbros form close to first-order axial discontinuities as a result of extreme cooling, the PAR lavas occur in an area without large offsets in the spreading axis (Fig. 3.1). On the other hand, Natland and Dick (1996) proposed that the melt lens at fast-spreading axes consists of evolved melts from the low-temperature top of a large more mafic cumulate body. The occurrences of FeTi basaltic and evolved glasses along the PAR indicate that such melts are not restricted to large discontinuities but may also occur in the centres of segments. We conclude that the chemical and petrological similarities between the PAR FeTi basalts, andesites and dacites and evolved plutonic rocks imply a close relationship by fractional crystallisation processes in a very shallow melt lens.

3.5.4 Fractionation of the PAR magma: oxide crystallisation and its effects on the magma composition

Rarely basaltic melts at mid-ocean ridges fractionate so extensively that andesitic and more silicic melts form but tholeiitic melts show a typical evolutionary trend of early Fe and Ti enrichment and decreasing abruptly at about 4 wt% MgO as a result of the onset of crystallisation and fractionation of FeTi oxides (Juster et al., 1989). In a closed system, this onset changes the parameters of fractional crystallisation significantly, for example, by decreasing the f O2 due to oxide formation (Byers et al., 1984). The change of the f O2 affects other elements and, for example, we observe that the evolved melts develop a much more negative Eu anomaly than the basalts in agreement with a decreasing f O2 at about 4 wt.% MgO (Fig. 3.6B). This stronger fractionation of Eu relative to the other rare earth elements is due to the fact that the Eu2+/Eu3+ ratio of a magmatic liquid depends on oxygen fugacity and Eu2+ is more compatible in plagioclase than Eu3+ (Drake, 1974; Drake and Weill, 1975; Philpotts, 1970; Sun et al., 1974). Consequently, more reduced melts have higher Eu2+/Eu3+ leading to stronger incorporation into fractionating plagioclase and a much more pronounced negative Eu anomaly in the silicic than the basaltic melts.

It has been suggested that the crystallisation of oxides leads to a lower solubility of S 2− 2− in silicate melts because SO4 is reduced to S which is then removed from the melt 3.5. Discussion 49 as Cu-Fe-sulfides (Jenner et al., 2010; Sun et al., 2004). Sulphur is significantly lower in most of the silicic glasses (Fig. 3.5A) but similar to other MORB studies (e.g. Jenner and O‘Neil, (2012)) we find that Cu is removed much earlier from the melts compared to the supposed onset of oxide fractionation (Fig. 3.4C), suggesting incorporation into another phase. For example, Cu may also partition into S-bearing volatiles that are degassing at depth from the magmas, (Rubin, 1997; Yang and Scott, 2002). In contrast, Co shows relatively constant contents at MgO higher than 4 wt.% but decreases at lower MgO (Fig. 3.4A) indicating that Co partitions into sulfides.

Hofmann et al. (1986) found that the Nb/U ratios in MORB are constant but in the PAR glasses we find a decrease in Nb/U (and Ta/U) in glasses with less than 4 wt.% MgO similar to the variation observed in the silica-rich EPR lavas (Wanless et al., 2010; 2011) (Fig. 3.6D). This decreasing trend could indicate that either Nb (and Ta) is more compatible than U in the fractionating assemblage or that U is assimilated from the hydrothermally altered rocks (Wanless et al., 2011). Because the Th/U ratio (not shown) increases slightly from the basaltic to the andesitic glasses, there is no evidence for enrichment of U as a possible result of assimilation. We suggest that Nb (as well as Ta) contents decrease relative to the U concentrations due the extensive ilmenite/Ti- magnetite crystallisation and fractionation. These oxide minerals have high distribution coefficients for Nb and Ta (Ewart and Griffin, 1994; Green and Pearson, 1987; Klemme et al., 2006) and fractionation of ilmenite and Ti-magnetite in evolved liquid can lower the Nb/U. Fractionation of apatite may be possible in the most evolved liquids but would lead to an increase in Nb/U because U is more compatible in apatite than Nb (Prowatke and Klemme, 2006). Thus, we conclude that the crystallisation of FeTi-rich oxide minerals is an important process during generation of silicic magmas leading not only T to decreasing FeO and TiO2 but also to decreasing S, to the pronounced Eu anomalies, and to decreasing Nb/U in evolved lavas.

3.5.5 Thermobarometric constraints on the depth of the mafic magma reservoir

In order to constrain the temperature and pressures of the fractionating PAR basaltic melt we use the clinopyroxene – melt thermobarometer of Putirka (2008). We only use values which meet the conditions described in Putirka (2008). Putirka et al. (1996) suggested that error analyses indicate that P estimates are as good as 1.0 kbar if averages of multiple equilibrium pyroxene/liquid pairs are used. Thus, we have compared two independent glasses and their corresponding clinopyroxenes and the pressure estimations are within standard deviation. The analyzed basaltic clinopyroxene cores (n=26) yield temperatures and equilibrium pressures that are similar within error of 1199 ( 8)◦C and 1212 ( 6)◦C and 3.4 ( 0.8) kbar and 3.9 ( 0.5) kbar, respectively (Table 3.3). The clinopyroxenes in the evolved lavas are in disequilibrium with the liquid composition 3.5. Discussion 50

Table 3.3: Thermobarometry (see Putirka, 2008) on the basis of clinopyroxene (cpx)-melt (glass composition) equilibrium calculations.

Sample cpx Oxide in wt.% Calculated according to Putirka et al. 2003 ◦ SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2OK2O Cr2O3 Total P (kbar) T ( C) SO157 17DS 3 1 51.9 0.948 3.60 6.98 0.218 16.5 19.7 0.284 0 0.319 100.5 3.9 1205 2 51.9 0.674 3.24 6.59 0.231 16.6 19.8 0.232 0 0.258 99.6 2.8 1194 SO157 17DS 1 1 51.2 0.897 4.15 6.51 0.195 17.1 19.6 0.277 0.011 0.283 100.3 4.3 1216 2 51.6 0.868 4.13 6.36 0.229 17.7 18.7 0.259 0.01 0.260 100.0 4.0 1219 3 51.2 0.907 4.02 6.93 0.200 16.9 19.5 0.271 0.004 0.214 100.1 4.2 1216 4 51.1 0.908 3.76 6.62 0.211 17.0 19.1 0.237 0.011 0.318 99.2 3.5 1211 5 52.0 0.781 3.38 6.30 0.161 16.8 20.3 0.246 0.006 0.309 100.3 3.4 1205 6 51.4 0.928 3.88 6.65 0.194 16.3 20.4 0.252 0.002 0.289 100.2 3.6 1207 7 51.0 0.882 3.71 6.41 0.168 16.7 19.8 0.264 0 0.383 99.3 4.0 1212 8 50.7 0.994 4.36 7.06 0.214 17.2 18.6 0.246 0.004 0.286 99.5 3.8 1218 9 51.4 0.876 3.52 6.36 0.138 16.8 20.2 0.254 0.014 0.282 99.8 3.6 1207 10 51.2 0.906 3.85 6.19 0.185 16.1 20.6 0.274 0.002 0.197 99.5 4.0 1209 11 50.9 1.04 3.78 6.75 0.177 16.3 20.2 0.254 0 0.098 99.5 3.7 1208 12 50.9 0.984 4.03 6.83 0.143 17.0 19.6 0.252 0 0.288 100 3.8 1212 13 51.4 0.742 3.64 6.24 0.158 16.7 20.0 0.290 0.009 0.393 99.5 4.4 1215 14 51.5 0.731 3.43 6.34 0.144 17.1 19.5 0.247 0 0.327 99.4 3.6 1210 15 50.9 1.01 4.08 6.49 0.176 16.5 20.0 0.283 0 0.390 99.8 4.3 1215 16 51.5 0.663 3.65 6.49 0.177 17.1 18.8 0.273 0.002 0.399 99.1 4.3 1219 17 51.9 0.895 3.89 6.64 0.175 17.1 19.4 0.279 0.012 0.364 100.6 4.3 1218 18 51.4 0.748 3.93 6.32 0.143 16.7 20.3 0.253 0.003 0.403 100.1 3.7 1208 19 51.3 0.789 3.68 6.73 0.186 16.5 20.0 0.250 0.006 0.276 99.7 3.6 1208 20 51.8 0.741 3.44 6.93 0.209 16.7 19.6 0.267 0.002 0.152 99.8 4.0 1212 21 51.0 1.06 4.51 6.71 0.211 16.4 19.4 0.298 0 0.200 99.7 4.7 1222 22 50.5 1.05 4.74 7.04 0.209 16.4 19.2 0.295 0.008 0.072 99.5 4.7 1223 23 51.9 0.768 3.09 6.66 0.238 16.7 19.9 0.229 0 0.169 99.7 3.1 1203 24 51.8 0.871 3.15 6.78 0.237 16.8 19.8 0.217 0.005 0.127 99.9 2.8 1201

(glass) and thus cannot be used for thermobarometry. The barometric results indicate that the basaltic melts started crystallizing at 9 ( 3) km depth which is in agreement with a slightly thickened crust of 9 km at 37◦S due to the increased melt production close to the Foundation Hotspot (Maia et al., 2001). We suggest therefore, that the basaltic liquids stagnate in a deep melt lens, fractionate olivine, plagioclase and clinopyroxene, and periodically inject melts into shallower levels in the crust where AFC leads to the formation of the FeTi basalts towards silicic magmas. Two magma reservoirs in the crust are believed to occur frequently in fast-spreading oceanic crust such as the East-Pacific Rise (Natland and Dick, 2009; Wanless and Shaw, 2012) or the Oman ophiolite (Boudier et al., 1996; Kelemen et al., 1997).

3.5.6 A quantitative AFC model for the PAR lavas

The high Cl and Cl/K but relatively low δ18O isotope ratios of the silicic PAR glasses indi- cate assimilation of hydrothermally altered crustal material (Fig. 3.8). In order to validate and expand the previous model of Haase et al. (2005) we use the Energy-Constrained

Recharge, Assimilation, and Fractional Crystallisation (EC-RAX FC) software (Bohrson and Spera, 2007) to model the evolution of O isotope variation and Cl and K concentration (Fig. 3.8; Table 3.2b). Furthermore, we use AFC calculations of DePaolo (1981) to determine the trace element composition and we use the MELTS algorithm for modelling the effect of AFC on the major element composition. Koepke et al. (2008) suggested that 3.5. Discussion 51

6.5 Northern andesites ) Southern andesites W Northern basalts Figure 3.8: Oxygen isotopes versus Cl/K of Southern basalts the PAR and EPR (Wanless et al., Dacites 6.0 MOR lavas published 2011) and assimilation-fractional

(VSMO crystallisation (AFC) trend. The ‰

10 % PAR evolved lavas show increas- 18 O ing Cl/K with decreasing δ O 18 5.5 20 % suggesting assimilation of up to 30% low (3‰) δ18O rocks, most

30 % likely hydrothermal altered rocks from the gabbro – sheeted dyke 5.0 transition. AFC trend modelled 0.0 0.5 1.0 1.5 2.0 with EC-RAχFC (Bohrson and Cl/K Spera, 2007) granoblastic dykes from the Cocos Plate basaltic dyke complex near the PAR represent residual crustal rocks after partial melting. Thus, we have chosen a sample preferably less affected by prior partial melting events (relatively high K2O and H2O: 209-1256D-

80R-2, 92-102 (Neo et al., 2009); Table 3.2a) for modelling. The high K2O resides either in plagioclase or in amphibole (Koepke et al., 2008). We assume that the trace element composition of the assimilated PAR crustal material resembles an average PAR basalt. To produce the observed trend in the evolved PAR lavas we require a potential crystallisation assemblage of 44.5% clinopyroxene, 47.5% plagioclase, 2.2% olivine and 5.8% oxides (e.g. ilmenite) (result of MELTS AFC calculation) and an assimilant containing 30% amphibole, 25% clinopyroxene, 41% plagioclase and 4% magnetite (AFC calculation of trace elements after DePaolo, 1981) starting with a southern basaltic magma (SO157 44DS1). We conclude that the calculated AFC trends fit the compositional variation of the evolved PAR glasses reasonably well and can best explain the observed enrichment, particularly in K2O and selected trace element ratios. Certainly the amount of assimilation versus fractional crystallization varies between different samples and furthermore, the composition of the parental magma varies with distance from the Foundation mantle plume. However, the results of the trace element and O isotope modelling (EC-RAX FC) imply that the maximal amount of assimilated wall rock must be on the order of 30% relative to the amount of parental magma (Fig. 3.8).

3.5.7 The nature of the assimilant and the role of amphibole

The chemical and mineralogical composition of the assimilant is not known and, for example, both brines and amphibole have been suggested as potential carriers of the high Cl contents (Coogan, 2003; Michael and Schilling, 1989). Oceanic basalts on average have Cl concentrations of less than 0.04 wt.% (Byers et al., 1986; Delaney et al., 1978; Michael and Schilling, 1989; Saal et al., 2002; Wanless et al., 2010) similar to most PAR basaltic glasses presented here (average 0.02 wt.%). In contrast, the FeTi-basalts (>0.1 wt.% Cl) and all silicic glasses show much higher Cl concentrations ranging up to 1.1 3.5. Discussion 52 wt.% (Table 3.1, Fig. 3.5B), higher than the Cl contents (up to 0.8 wt.%) of the GSC and EPR evolved glasses discussed by Byers et al. (1984; 1983) and Wanless et al. (2010). At 30% assimilation of wall rock suggested by our model, the assimilated material needs to be enriched in Cl ( 0.85 wt.%). Sparks (1995) analysed altered basalts from sheeted dykes in ODP/IODP drill sites that have Cl concentrations of 0.02-0.07 wt.% which is too low to explain the enrichment in the evolved PAR glasses. Michael and Schilling (1989) suggested that the increased Cl content of fresh glasses may reflect assimilation of Cl-rich phases like trapped brines or amphiboles. Hydrothermal amphiboles may contain up to 4 wt.% Cl (Ito et al., 1983; M´evel, 1987; Vanko, 1986) and amphibole is an important constituent of oceanic crust altered at high temperatures by hydrothermal fluids (Gardien et al., 2001; Honnorez, 2001). Brines from the hydrothermal systems should be enriched in H2O and Li (Chan et al., 2002) but depleted in Ce compared with crustal rock. We observe that H2O/Ce and Li/Ce ratios decrease in the silica rich lavas compared to the basalts (Figd. 3.7A, B) implying that the assimilated material had relatively low H2O and Li contents compared to Ce unlike the assumed composition of highly saline brines. On the other hand, hydrothermally formed amphiboles appear to have low Li but relatively high Ce contents and water contents based on the totals appear to be generally lower than 2 wt.% (Coogan et al., 2001). We suggest that the most likely source of the high Cl and low H2O/Ce and Li/Ce is hydrothermally formed amphibole in the assimilated rocks. Partial melting of amphibole-bearing altered mafic rocks can also affect the trace element ratios in the evolved lavas (Haase et al., 2005; Wanless et al., 2010) and, for example, Tb and Sm are more compatible in amphibole than Yb and Hf, respectively (Davidson et al., 2007). Indeed we find significant variations of (Te/Yb)N and Hf/Sm between the basaltic and andesitic lavas that may also reflect the presence of amphibole during magma genesis (Figs. 3.6E, F). An influence on the trace elements as a result of crystallisation of amphibole from the PAR evolved magma is unlikely because no amphibole is observed in the samples and because of the relatively low water concentrations (average of 0.32 wt.% H2O in the basaltic glass (Table 3.1). More likely is the assimilation of melts produced by partial melting of high-temperature metamorphic amphibole-bearing rocks, similar to those observed beneath the sheeted dyke complex (France et al., 2010; Koepke et al., 2008). The observed increase in Ce/Pb with decreasing MgO (Fig. 3.6C) contrasts with the distinctly lower Ce/Pb ratios in evolved lavas from the EPR (Wanless et al., 2010) and probably indicates more extreme fractionation of plagioclase during fractional crystallisation of the PAR magma and a higher amount of assimilated material. Excessive plagioclase fractionation is supported by the decreasing trend in Sr/Nd (Fig. 3.6G) and the negative Eu anomaly because Sr and Eu2+ are compatible in plagioclase. In conclu- sion, most of the variations in highly incompatible element ratios reflect a combination of crystal fractionation, e.g. of plagioclase (Ce/Pb, Sr/Nd) and ilmenite/Ti-magnetite (Nb/U), and assimilation processes (Cl/K). 3.6. Conclusions 53

3.5.8 Potential relationship between occurrence of silicic magmas along spreading axes, tectonic and hydrothermal processes

Silicic magmas are relatively rare along the global ocean spreading system and their occurrence may be related to specific tectonic settings or geochemical heterogeneities. Silicic magmas occur at propagating rift tips, for example, close to the East Rift of the Easter Microplate (Naar and Hey, 1991) and also in the Valu Fa backarc (Fretzdorff, 2006). More recently, Wanless et al. (2010) suggested that the silicic magmas formed at propagating rift tips and at overlapping spreading centers as a result of waning magma supply after a dyke intrusion event. However, the PAR between 36.5 and 40◦S does not show significant offsets of the spreading axis (Lonsdale, 1994) (Fig. 3.1) and the relatively high spreading rate requires a high rate of magma replenishment. Thus, the PAR silicic lavas were generated in relatively hot crust and rift propagation into colder oceanic crust cannot cause the extreme fractionation toward andesitic-dacitic magma. Recent studies of seismic and geochemical data observed more evolved lavas on fast- spreading ridge segments with high magma supply rates (Colman et al., 2012; Toomey and Hooft, 2008) and we infer that the PAR close to the Foundation Seamounts must also be an area of relatively high magma supply. Furthermore, the evolved lavas on the PAR are associated with active hydrothermal venting (Stoffers, 2001) similar to other evolved lava occurrences coupled with hydrothermal activity (Galapagos spreading center (Perfit et al., 1999); central volcanoes in Iceland (Gunnarsson et al., 1998; Saemundsson, 1979)). Thus, we speculate that an increased magma supply due to inflow of hot Foundation mantle plume material leads to thicker crust and multiple magma reservoirs in the crust as well as to increased hydrothermal circulation in the shallow PAR crust. This may be the cause for the abundance of evolved lavas on the about 300 km long part of the PAR.

3.6 Conclusions

The generation of andesitic-dacitic low δ18O glasses and FeTi basalts at the PAR most likely is the result of a combination of AFC processes during magma ascent, conducted through an unusually thick, plume-influenced MOR crust (Fig. 3.9). Magma evolution starts with a basaltic melt in a deep crustal sill, indicated by clinopyroxene-melt barom- etry. Relatively evolved FeTi-basaltic melts ascend into the upper crust at perhaps 1 km depth where melt lenses form below the hydrothermally altered sheeted dykes. This wall rock becomes partly assimilated during crystal fractionation processes. EC-RAX FC (Bohrson and Spera, 2007) as well as MELTS (Ghiorso and Sack, 1995) models suggest >57% crystallisation of the basaltic magma and up to 30% assimilation of hydrothermally altered dykes are required to explain the major and trace elements as well as O isotope ratios observed in the andesitic-dacitic glasses from the PAR. The unusually low δ18O values of the silicic PAR glasses (>5.6‰) must be the result of assimilation of low δ18O wall rocks that are typically found at the sheeted dyke-gabbro transition zone (Alt and 3.6. Conclusions 54

hydrothermal venting

pillows crustal cooling sheeted dykes

AFC → andesite-dacite d & FeTi-basalt re lte l a ~ 1 km bsf rma st the ro cru yd ic h an oce

mush zone ~ 2 km bsf gabbros

deep basaltic melt sill ~ 8 km bsf upper mantle

Figure 3.9: Formation of the basaltic, FeTi basaltic and silica rich lavas at the PAR. AFC processes take place within a shallow magma chamber (not to scale). The basaltic magma ascends from a deep crustal magma-sill into the shallow crustal magma chamber at the base of the sheeted dyke complex (∼1000 mbsf). The extensive crystallisation process in the shallow magma chamber is linked to the hydrothermal cooling of the rust. The releasing heat of crystallisation results in assimilation of (hydrothermal altered) wall rock and leads to the formation of silicic melt with the discussed geochemical characteristics e.g. low δ18O values. 3.7. Acknowledgements 55

Teagle, 2000; Harper et al., 1988; Ito and Clayton, 1983; Stakes and Vanko, 1986). The release of crystallisation heat in the shallow crust most likely causes the breakdown and partial melting of amphibole-bearing altered rocks. Because the evolved PAR lavas occur in segments that show a geochemical influence of the Foundation mantle plume we infer that increased magma supply leads to thickened crust, multiple magma systems and increased hydrothermal circulation causing the formation of relatively large volumes of silicic magmas.

3.7 Acknowledgements

We gratefully acknowledge the help of the captains and crews of RV Sonne and RV Atalante and the chief scientists for access to the samples, particularly P. Stoffers. We thank P. Appel and B. Mader for helping with electron microprobe analyses in Kiel and D. Garbe-Sch¨onberg, U. Westernstr¨oer (Bremen) and M. Regelous (Erlangen) for ICP-MS trace element analyses. We also thank the Editor K. Mezger and two anonymous reviews for their constructive comments that considerably improved the quality of this work. We acknowledge the encouraging words and helpful points of M. Meyer, M. Erdmann and F. Genske. We thank P. Brandl for spiritual support during EMP problems and ChB acknowledges frequent delays of German rail allowing him to work on this manuscript. The cruises were funded by the German Ministry of Science and Technology (BMBF) and S.F. was funded by DFG grant HA2568/21. 4 Constraints on the formation of geochemically variable plagiogranite intrusions in the Troodos Ophiolite, Cyprus

Sarah Freund1, Karsten .M. Haase1, Manuel Keith1, Christoph Beier1, Dieter Garbe-Sch¨onberg2 1GeoZentrum Nordbayern, Universit¨at Erlangen-Nurnb¨ erg, Schlossgarten 5, 91054 Erlangen, Germany 2Institute of Geosciences, Universit¨at Kiel, Ludewig-Meyn-Str. 10, 24118 Kiel, Germany

4.1 Abstract

The geochemistry and petrology of tonalitic to trondhjemitic samples (n= 85) from eight different plagiogranite intrusions at the gabbro/sheeted dyke transition of the Troodos Ophiolite were studied in order to determine their petrogenetic relationship to the mafic plutonic section and the lava pile. The plagiogranitic rocks have higher SiO2 contents than the majority of the glasses of the Troodos lava pile, but lie on a continuation of the chemical trends defined by the extrusive rocks, indicating that the shallow intrusions generally represent crystallized magmas. We define three different groups of plagiogranites in the Troodos Ophiolite based on different incompatible element contents and ratios. The first and most common plagiogranite group has geochemical similarities to the tholeiitic lavas forming the lavas and sheeted dyke complex in the Troodos crust, implying that these magmas formed at a spreading axis. The second plagiogranite group occurs in one intrusion that is chemically related to late-stage and off-axis boninitic lavas and dykes. One intrusion next to the Arakapas fault zone consists of incompatible element-enriched plagiogranites which are unrelated to any known mafic crustal rocks. The similarities of incompatible element ratios between plagiogranites, lavas and mafic plutonic rocks, the continuous chemical trends defined by plagiogranites and mafic rocks, as well as incompatible-element modeling results, all suggest that shallow fractional crystallization is the dominant process responsible for formation of the felsic magmas. 4.3. Geological Background 57

4.2 Introduction

The oceanic crust consists dominantly of mafic rocks, because partial melting of mantle peridotite produces a basaltic magma (e.g. Langmuir et al. (1992)). Significant volumes of felsic rocks also occur in the oceanic crust but the processes leading to their formation are not well understood. In particular, the relative roles of fractional crystallization from basaltic melts or partial melting of crustal rocks are debated (Gillis and Coogan 2002; Koepke et al. 2004; Berndt et al. 2005; Koepke et al. 2007; Brophy 2009; Rollinson 2009; France et al. 2010; Brophy and Pu 2012). The composition and structure of the oceanic crust apparently affects the composition of ascending magma and, for example, the high melt supply at fast-spreading mid-ocean ridges leads to shallow melt reservoirs and relatively evolved magmas (Rubin et al. 2009). Extremely evolved magmas ranging to dacites or even rhyolites occur in areas of thickened crust near hotspots like the Galapagos Spreading Center, Iceland and the Foundation Seamount Chain (Byerly et al. 1976; Muehlenbachs 1982; Juster et al. 1989; McBirney 1993; Geist et al. 1995; Gunnarsson et al. 1998; Freund et al. 2013) or close to large segment offsets (Wanless et al. 2010; Wanless et al. 2011). Felsic plutonic rocks such as tonalites, trondhjemites and granodiorites are known from several ophiolites and from active spreading centers, and have generally been termed ’oceanic plagiogranites’ (e.g. (Moores and Vine 1971; Amri et al. 1996; Caselli and Lombardo 2007; Meffre et al. 2012)). Well-preserved ophiolites such as the Troodos Ophiolite in Cyprus offer the possibility to advance our understanding the formation of felsic magmas in the oceanic crust and to examine the relationships between extrusive and intrusive rocks.

Here we present a systematic petrological and geochemical study of most of the large felsic intrusions from the gabbroic and sheeted dyke portion of the Trodoos Ophiolite. Our samples include rocks with tonalitic/trondhjemitic composition and are generally termed plagiogranite in the following text. We show that three geochemically distinct types of felsic intrusions exist in the Troodos Ophiolite, comparable to different plagiogranite types in the Oman Ophiolite (e.g. Rollinson (2009)). The most abundant tholeiitic plagiogranite intrusions in the upper 2 km of the Troodos crust most likely differentiated by fractional crystallization in shallow melt lenses directly below the dykes. The geochemical variability of the plagiogranites indicates parental melts produced by an increasingly depleted mantle source with continuing subduction influence.

4.3 Geological Background

The Troodos Ophiolite of Cyprus formed in the eastern Mediterranean region of Tethys between 90 to 92 Ma (Cenomanium – Turonium) as indicated by U-Pb ages from zircons from a plagiogranite intrusion between Platanistasa and Amiantos (Mukasa and Ludden 4.3. Geological Background 58

1987). The oceanic crust of the Troodos Ophiolite remained in situ until obduction started in the Neogene (Robertson and Hudson 1973). The Troodos Ophiolite represents one of the least altered or tectonically disturbed fragments of oceanic crust on land, and includes the complete sequence of pelagic sediments, pillow lavas, sheeted dykes, felsic to ultramafic plutonic rocks, and mantle peridotites that has been defined in the oceanic crust by seismic data. The intrusive rocks of the Troodos Ophiolite are best exposed between the Solea Graben (thought to represent a former spreading axis) in the north, and the fossil oceanic Arakapas transform fault in the south-east (Abelson et al. 2001). The rocks can be subdivided into several units that show a roughly annular arrangement around its plutonic central part. Ultramafic complexes containing harzburgite, dunite, wehrlite and pyroxenite occur in two separate areas of the main massif: the Mount Olympos and the Limassol Forest (south of the Arakapas transform). Gabbros, diorites, plagiogranites and a tholeiitic sheeted dyke complex containing plagiogranitic intrusions overlie these units. The approximately 1 km thick lava pile consists largely of tholeiitic rocks (Pearce and

Robinson 2010) but boninitic lavas (> 8 wt% MgO, > 52 wt% SiO2 and < 0.5 wt% TiO2 (Le Bas 2000)) occur in some regions overlying these tholeiites (Cameron 1985; Flower and Levine 1987; K¨onig et al. 2010). The tectonic setting in which the Troodos Ophiolite formed is debated but the geochem- ical composition of the lavas and intrusive rocks suggests a subduction-related origin (Miyashiro 1973; Pearce and Cann 1973; Rautenschlein et al. 1985). The well-developed sheeted dyke complex in the main massif of the Troodos Ophiolite implies significant and continuous extension possibly indicating an origin in a back-arc setting, which is consistent with observed high volatile contents in the glasses (Muenow et al. 1990). The Troodos Ophiolite likely was formed at an intermediate (Staudigel et al. 1999) to fast spreading axis (Bednarz and Schmincke 1993; Cann and Gillis 2004). Alternative models suggest that the Troodos Ophiolite rocks formed at a slow-spreading mid-ocean ridge (Varga et al. 1999) possibly influenced by lateral magma plumbing along the axis (Abelson et al. 2001). However, the distinct negative anomalies of the high field strength elements (HFSE) in the Troodos Ophiolite, together with the occurrence of boninites suggest a subduction- related tectonic setting. The boninitic lavas are comparatively rare and spatially restricted e.g. in the Limassol area, south of the Arakapas Fault Zone (Cameron 1985), whereas boninitic dykes also occur in the plutonic section of the main massif (Coogan et al. 2003). Boninitic lavas were believed to be restricted to fore-arc settings (Crawford et al. 1989) and formed during subduction initiation (Pearce and Robinson 2010), but due to a lack of modern examples the specific magmatic and tectonic processes responsible for formation of boninite magmas are poorly known. Recently discovered boninites at active back-arc and island arc volcanoes (Cooper et al. 2010; Resing et al. 2011) indicate that boninites may form in several different tectonic settings. Compositional differences are observed within the plutonic rocks of Troodos, specifically in mineral compositions, and one model suggests that the Troodos lower crust formed 4.4. Samples and analytical methods 59 by two chemically unrelated stages (Thy 1987); the upper gabbros, the sheeted dyke complex, and the lower pillow lavas represent a spreading environment, whereas the lower cumulate gabbros are related to the upper pillow lavas and formed during closed- system crystallization following off-axis magmatic underplating. According to Benn and Laurent (1987) the early gabbros and pyroxenites are intruded by a later suite of poikilitic wehrlites and gabbro-norites. These authors suggest that the two plutonic stages can be explained by a second-stage melting event of a mantle diapir, or by compaction of a zone of crystal mush below the Moho. In contrast, Schouten and Kelemen (2002) suggest that the lavas formed from one evolving source due to different reservoir depths in the crust with primitive basaltic (upper pillow lavas) forming voluminous eruptions whereas the evolved more viscous lower pillow lavas erupted frequently but in small volumes. Coogan et al. (2003) argue that the highly variable clinopyroxene core compositions of the mafic and ultramafic plutonic rocks compared with the considerably less variable lava compositions indicates a filter function of the lower crust. Highly variable primitive parental melt compositions leave the mantle and mix and homogenize within the crustal plumbing system before erupting (Coogan et al. 2003).

4.4 Samples and analytical methods

The felsic plutonic samples (n=85) were collected during a field trip in 2010 from eight different intrusions in the gabbros, sheeted dykes and the transition zone around Mount Olympos (Fig. 4.1). Fresh cores from samples were cut with a rock saw, washed in deionized H2O, crushed and pulverized in an agate mill. We determined the loss on ignition (LOI) by weighing the rock powder before and after drying (1) 12 hours at 105 ◦C in a cabinet dryer and (2) 12 hours at 1030◦C in a muffle furnace. Representative thin sections (n=60) were studied petrographically and mineral contents were analysed by an X-ray diffractometer 2 Siemens D5000 (samples n=10) at the GeoZentrum Nordbayern. The major element and a few trace element concentrations (Cu, Ni, Zn, Cr, V, Rb, Sr, Zr, Nb and Y) of whole rock powders were measured using an XRF spectrometer (Spectro XEPOS plus) at the GeoZentrum Nordbayern. Averages of our measurements and recommended values (Govindaraju and Roelandts 1993; Govindaraju 1994; Imai et al. 1995) of rock standards (BIR-1, BR, JA-3) are given in Table 4.1.

The major element concentrations of minerals (supplementary Table 4.2) in polished thin sections were measured with a JEOL JXA 8200 Superprobe electron microprobe at the

GeoZentrum Nordbayern. SiO2, TiO2, Al2O3, Fe2O3, MnO, MgO, CaO, Na2O, K2O,

Cr2O3 (and Cl) were measured. The EMP was operated with an accelerating voltage of 15 kV, a beam current of 15 nA and a focused beam except for plagioclase, where a slightly defocused beam of 3 µm was used in order to minimize Na loss. Counting times were set to 20 s and 10 s for peaks and backgrounds for most elements, and 40 and 20 s 4.4. Samples and analytical methods 60 for Cl, respectively. Additionally we measured trace element concentrations of minerals (measurement conditions and standards are listed in supplementary Table 4.2).

Trace elements of the whole rock powders were analysed using an Agilent 7500c/s Quadrupole Inductively Coupled Plasma Mass Spectrometer (ICP-MS) at the Institute of Geosciences, Universit¨at Kiel following procedures described previously (Garbe-Sch¨onberg 1993). Ana- lytical precision was monitored by the repeated analysis of one sample yielding <3% RSD for most elements except Zr, Th, U (<7%), and Nb, Ta, (<27%). Results for international rock standards BIR-1 and BHVO-2 are compiled in Table 4.1. The reproducibility of the digestion procedure as monitored by duplicate sample digests is better than 3% for all elements except for Cu, Mo, W (<10%). One sample, however, showed systematic larger errors of <5 to 25% that may reflect inhomogeneity of the sample powder. Comparison of the XRF and ICP-MS data for Zr and (to a lesser extent) Hf imply that during acid digestion zircon crystals had not been completely dissolved in the more evolved samples. Thus, we measured three representative rock samples (and international rock standard BE-N) by laser ablation ICP-MS and compared the data (for further details, methods and comparison, see supplementary material). We will use Zr data from the XRF analyses for diagrams and discussion in this study.

SCALE 1:250000 20 km 33°E 34°E

Troodos ophiolite Cyprus Pedoulas Spilia Lemithou Platanistasa 35°N Mount Larnaka Olympos Palaichori N Amiantos

Zoopigi

Serpentinized harzburgites, minor dunites Plagiogranite intrusions

Tectonized harzburgites, minor dunites Sample locations:

Wehrlites Diabase dykes Zoopigi group

Websterites Basal Group (diabase Main group dykes and pillow lavas) Gabbros Spilia group

Figure 4.1: Map of the Troodos Ophiolite in Cyprus (overwiew) and detailed geology of the Mount Olympos area showing the distribution of plagiogranite intrusions within the gabbros and sheeted dykes. Sample locations are marked with a yellow star (main plagiogranites), diamond (Zoopigi plagiogranite) and circle (Spilia plagiogranite). The Limasol Forest area, separated by the Arakapas Fault is affiliated near the southern edge of the detailed map.). 4.5. Results 61

4.5 Results

4.5.1 Occurrence of the plagiogranites and field observations

The felsic plutons studied here occur in different crustal levels in the gabbro/sheeted dyke transition zone of the Troodos Ophiolite; most are located within the sheeted dykes, but some also occur within the gabbro section (Fig. 4.1). North of the village Zoopigi (SE of Mount Olympos) numerous felsic dykes and plutonic bodies (in the following called Zoopigi plagiogranites) alternate with aphyric dykes and are often intruded by late aphyric dykes. These plagiogranitic dykes are strongly tectonized by faults and shear zones (Fig. 4.2a) possibly because of their location close to the Arakapas Fault zone which may thus indicate a relatively early formation compared to the other plagiogranite bodies in the Troodos Ophiolite. Most of our samples (hereafter called main group) are from felsic intrusions in the gabbros and the sheeted dyke/gabbro transition, for example along road cuts near Pedoulas and Lemithou and from small intrusions NE of Mount Olympos, along the road from Amiantos towards Pelendri, from Alona towards Platanistasa, and from Alona towards Fterikoudi (Fig. 4.1). Near the village of Palaichori we sampled a well- preserved outcrop of felsic, medium-grained dykes and plutonic intrusions alternating with aphyric dykes (Fig. 4.2b). Additionally, aphyric (andesitic - dacitic) dykes frequently cut these intrusions but do not show significant faulting. Some of these plagiogranitic rocks show a mottled texture (Fig. 4.2c) with abundant rounded dark xenoliths up to 15 cm in diameter. Six aphyric andesitic-dacitic dykes crosscutting the Palaichori intrusion were also analysed (Table 4.1).

One large plagiogranite intrusion (approximately 1 km diameter) is located within the gabbro/sheeted dyke transition zone close to Spilia (called Spilia plagiogranite in the following, Fig. 4.1). The intrusion mainly consists of (high An-) tonalites with few mafic xenoliths up to 1 m in diameter at the rims of the intrusion (Fig. 4.2d). Occasionally, the plagiogranites show a mottled appearance (Fig. 4.2e) with dark-coloured parts of less evolved (64 wt% SiO2, 3 wt% MgO) and light-coloured parts of highly evolved (75 wt%

SiO2, 1.1 wt% MgO) composition. In contrast to the Zoopigi and main group intrusions, the Spilia intrusion apparently has only rarely been intruded by dykes and thus most likely formed relatively late compared to the other evolved plutonic rocks. Based on the field observations of alteration, intrusive contacts, tectonic overprint and geochemical differences, we define three different groups of plagiogranitic plutons namely the Zoopigi plagiogranites, the main group (Fig. 4.1), and the Spilia plagiogranites.

4.5.2 Petrology and mineral chemistry

Most of the plagiogranites are equigranular and fine- to medium-grained rocks even in the larger plutonic bodies. Thin section observations (n=60) and X-ray diffraction analysis (n=10) of the samples show that the main mineral parageneses of all rocks consists of

4.5. Results 64

An

Tonalite

Granodiorite Zoopigi group Main group Spilia group Ab Trondhjemite Or

Figure 4.4: Classification of plagiogranites as tonalites and trondhjemites (CIPW norms calculated from major element compositions of the rocks with more than 65 wt.% SiO2) using the scheme suggested by (Barker and Arth 1976). frequently arranged in clusters and often associated with oxides. Calcic-amphiboles are the predominant amphiboles (magnesiohornblende – ferrohornblende - actinolite – ferroacti- nolite) in all groups (supplementary Tables 4.1 and 4.2). The gabbroic sample TroPed

51 contains orthopyroxene (n=11; Woll2−4, En67−70, Fer27−30; enstatite-ferrosilite) with actinolitic rims. Anhedral clinopyroxene (hedenbergite: Woll54−56, En25−33, Fer19−27) occurs instead of amphibole in the tonalitic sample TroAm 23B (main group).

Titanite is found in most samples and is euhedral to subhedral (<100 µm). Epi- dote/clinozoisite and chlorite are most likely the result of hydrothermal alteration with seawater after solidification and these minerals often occur in veins. Apatite occurs irregularly distributed as fine acicular inclusions within quartz in most samples. Zircons are likely rare and very small in size; no zircons were identified in thin section or by X-ray diffraction, but must be present as indicated by lower Zr contents in the ICP-MS analyses compared to the XRF data.

4.5.3 Major element composition of the rocks from plagiogranite intrusions

Based on the CIPW norms calculated from the major element compositions, the rocks with more than 65 wt.% SiO2 are classified as tonalites and trondhjemites (Fig. 4.4) using the scheme suggested by Barker (1976). Most of the main group and Spilia plagiogranites 4.5. Results 65 are tonalites, whereas the Zoopigi rocks range from trondhjemites to tonalites. The whole rock compositions of the three plagiogranite groups overlap the chemical trends displayed by volcanic glasses from the extrusive rocks (Pearce and Robinson (2010), Regelous et al. in prep.) and sheeted dyke rims (Staudigel et al. 1999), but most are distinctively more

SiO2-rich (Fig. 4.5a). The SiO2 contents of the samples from the plagiogranite intrusions are generally higher than 60 wt% whereas few of the volcanic glass samples have such high SiO2 contents. The most SiO2-rich plagiogranites have the lowest TiO2, CaO and

P2O5 contents and the highest Zr and relatively high Na2O concentrations. The three plagiogranitic groups have similar contents of CaO, FeO, MnO and MgO for a given SiO2. The intrusions belonging to the main group (n=56 + 6 silicic-aphyric dykes) contain rocks of a wide SiO2 range (56 – 71 wt%) whereas the Zoopigi and Spilia intrusions also include more SiO2-rich rocks up to 79 wt%. Generally, all plagiogranitic rocks studied here lie along a steeply increasing trend of SiO2 at MgO lower than 4 wt%, on a continuation of the trend defined by the volcanic glass data (Fig. 4.5a). In general, the TiO2 contents of the SiO2-rich intrusives of the Troodos Ophiolite are low (<1.1 – 0.1 wt%) and decreases with increasing SiO2 (Fig. 4.5b) similar to the steep decrease observed in FeO (Fig. 4.5d).

The CaO and Al2O3 contents also decrease with increasing SiO2 (Fig. 4.5c, e), but the rocks from the Spilia intrusion generally have lower Al concentrations than the other plagiogranite samples. Rocks of the Spilia group have unusually low TiO2 concentrations for a given SiO2 compared to the other groups. The Spilia samples display relatively high K2O concentrations (mean: 0.4 wt%) compared to the mean values of the other two groups (about 0.2 wt%, not shown). The samples of the main group have a range in P2O5 contents of 0.11 to 0.16 wt% at SiO2 of about 69 wt% and then decreasing contents. The

Zoopigi and Spilia plagiogranites generally have lower P2O5 and a slight increase to 0.08 wt% at 75 wt% SiO2 and then show a decrease (Fig. 4.5f).

4.5.4 Trace element composition of the plagiogranitic intrusives and comparison with the glassy lava

All samples show a strong variability of the fluid-mobile large ion lithophile elements (LILE) (e.g. K = 566-7620 ppm; Ba = 1.5-68.4 ppm, Rb = 0.24-3.8 ppm), which is likely due to hydrothermal alteration. The fluid-immobile element Zr indicates some differences between the three groups defined above (Fig. 4.6d). Whereas the Zoopigi plagiogranites have a relatively high range in Zr contents of 64 to 141 ppm at about 77 wt% SiO2, all except one (TroSpi 32 = 305 ppm, not shown) of the Spilia samples tend to relatively low Zr concentrations (<63 ppm). The majority of the main group of plutonic rocks are intermediate between the other two plagiogranite types. Interestingly, the variation of Nb shows the opposite situation; the Spilia samples have higher Nb contents and several of the Zoopigi rocks have lower Nb contents whereas the main group samples are highly variable (Fig. 4.6e). Fluid-mobile elements such as K (not shown), U and importantly

Cu and Zn (Fig. 4.6b, c, f) show significantly lower contents at a given SiO2 than the 4.5. Results 66

Zoopigi group Spilia group main group aphyric dykes 15 15 a) gabbro (this study) d) gabbro published glasses (pubished and Regelous et al., in prep.) 10 10 boninite data published boninite (this study)

5 5 FeO [wt.%] MgO [wt.%]

0 0

2.5 b) e) 15 2.0 1.5 10 [wt.%] 2 1.0 iO CaO [wt.%] T 5 0.5

0.0 0 fractional crystallization 0% H2O 0.3 c) 0.5% H2O f) 20 1.5% H2O boninitic pm 0.2 16 [wt.%] [wt.%] 3

5 O

O

2 0.1 2 12

P Al

8 0.0 45 50 55 60 65 70 75 80 45 50 55 60 65 70 75 80

SiO2 [wt.%] SiO2 [wt.%] Figure 4.5: Major element variations of the plagiogranite groups compared to the Troodos glass data (Staudigel et al. 1999; Pearce and Robinson 2010) (Regelous et al. in prep) and boninite data (glass and whole rock) (Flower and Levine 1987; K¨onig et al. 2010) (Regelous et al. in prep). Gabbros from the Cy-4 bore hole near Palaichori (Laurent and H´ebert 1989) and from the Pedulas area (TroPed 51, this study). Additionally, fractional crystallization trends calculated with MELTS (Ghiorso and Sack 1995) under varying H2O content (for further details see main text).

4.6. Discussion 68 volcanic glasses. In contrast, the Co contents of the plagiogranites follow the same trend as the extrusive rocks, and decrease from 60 wt% SiO2 (Fig. 4.6a). The fluid-immobile rare earth elements (REE) also display significant differences between the three groups (Fig. 4.7). The main group shows gently inclined chondrite-normalised

REE patterns with slightly depleted light REE (LREE). They have (Ce/Yb)N between 0.4 and 0.8 similar to the variation observed in the Troodos tholeiitic lavas, for example from the Akaki Canyon (Regelous et al. in prep.) (Figs. 4.7, 4.8). The main group plagiogranites generally display negative Eu anomalies but some of the xenolith-rich samples have a positive Eu anomaly. The Zoopigi plagiogranites have smooth REE patterns with flat or slightly enriched LREE compared to the other two groups and have (Ce/Yb)N ratios of 0.6 to 1.2 (Figs. 4.7, 4.8) and negative Eu anomalies. Lavas with comparable incompatible element composition to the Zoopigi plagiogranites are not known from the Troodos Ophiolite so far, and thus the relationship of the former to other magmatic rocks in terms of generation and relative age cannot be determined. The rocks from the Spilia plagiogranites are depleted in LREE and MREE (Fig. 4.7) compared to the tholeiitic lavas and the other plagiogranites, having (Ce/Yb)N less than 0.45 (Fig. 4.8). The Spilia samples have comparable REE patterns to the boninitic lavas and late dykes intruding the Troodos plutonics (Flower and Levine 1987; Coogan et al. 2003; K¨onig et al. 2008) showing a typical U-shape.

4.6 Discussion

4.6.1 Hydrothermal alteration of the samples and stratigraphic context of the plutonic section

Minerals such as chlorite, albite, secondary quartz and possibly also epidote/clinozoisite and actinolite within the plagiogranites are the result of a pervasive greenschist to lower amphibolite facies metamorphism (Malpas et al. 1989). Malpas et al (1989) suggest a second, more local type of alteration by intense hydrothermal leaching and epidotisation, for example, spatially limited in the Palaichori area. Based on the CY-4 drill hole Laurent (1990) suggests that the uppermost approximately 1000 m of the Troodos Ophiolite, including the sheeted dykes and the pluton roof consisting of massive amphibole gabbros and plagiogranites, are pervasively altered whereas the deeper cumulate plutonic rocks are mainly altered along tectonized zones. Laurent (1990) attributes the metamorphic alteration to the late intrusion of felsic magmas whereas Schiffman and Smith (1988) and Richardson et al. (1987) propose that the alteration is the result of upwelling hydrothermal fluids. Although we have tried to sample relatively fresh rocks, minor amounts of chlorite, Na-rich plagioclase, secondary quartz, and epidote/clinozoisite occur in our samples (<2%, supplementary Table 4.1). The concentrations of fluid-mobile elements such as Rb and K

4.6. Discussion 71

3

2

1 LOI [wt.%]

a) b) 0 0 1 2 3 4 0 2 4 6 Rb [ppm] Nb [ppm] 3 Zoopigi group Spilia group Main group aphyric dykes 2 gabbro (this study)

1 LOI [wt.%]

c) d) 0 0 100 200 300 0 5 10 15 20 25 Zr [ppm] Ce [ppm]

Figure 4.9: Diagram of loss on ignition (LOI) vs. mobile and immobile elements. a) LOI versus Rb. Note that the plagiogranites are positively correlated, suggesting Rb addition during hydrothermal alteration. b), c), d) LOI versus Nb, Zr, and Ce showing no distinct correlation and group subdivision is visible, suggesting that hydrothermal alteration is not the reason for incompatible element difference between the three groups. are positively correlated with loss on ignition (Fig. 4.9), and thus we suggest that these elements have been enriched by hydrothermal alteration. In contrast, fluid-mobile metals like Zn (Fig. 4.6c) show extreme depletion by factors of up to 90 % compared to the felsic glasses, and this confirms that leaching effectively transports these metals from the uppermost plutonic rocks into the hydrothermal systems and potentially into the sulfide deposits associated with the pillow lavas (Heaton and Sheppard 1977; Richardson et al. 1987). The fluid-immobile incompatible trace elements Ti, Nb, Zr and the REE show systematic variations in the plutonic rock groups, and we conclude that hydrothermal alteration processes did not significantly the concentrations of these elements (Fig. 4.9).

For example, the Spilia group has different contents of TiO2, P2O5, Zr, and LREE compared to the other groups (Figs. 4.5, 4.6, 4.8). The systematic variations in these elements between groups cannot be the result of alteration processes. 4.6. Discussion 72

4.6.2 Relations between evolved intrusive rocks and the lavas

Whereas the lavas and specifically the volcanic glasses in the Troodos Ophiolite represent melt compositions, most of the plutonic rocks likely represent mixtures of melt and accumulated minerals, and thus the chemical composition of intrusive and extrusive rocks may not be directly comparable. However, the isotropic gabbros and plagiogranites from shallow intrusions in the oceanic crust frequently do resemble crystallized liquids, especially if they are relatively fine-grained (Kelemen et al. 1997; Natland and Dick 2009). We suggest that most of the Troodos plagiogranites represent crystallized magmas because they are relatively fine-grained, often occur as dykes, and their major element compositions generally overlap with those of glassy lavas and continue those trends to lower SiO2 (Fig. 4.5). If the plagiogranites represented cumulates then we would expect that they vary less systematically and lie outside of the glass trend towards higher MgO, FeO, or Al2O3 contents. The compositions of the plagiogranite intrusions are generally more SiO2-rich than the glassy lavas sampled from the Troodos Ophiolite (Pearce and Robinson (2010), Regelous et al. in prep.) and the glass rims of the sheeted dykes (Staudigel et al. 1999) but overlap between 58 and 74 wt% SiO2 (Fig. 4.5). Specifically, the main group mostly continues the compositional trend formed by the tholeiitic Troodos glasses suggesting a close genetic relationship to these lavas. This is also supported by the similarity of fluid- immobile incompatible element ratios like Sm/Nd, Nb/Zr and (Ce/Yb)N (Fig. 4.8). For example, the plagiogranites of the main group show the same LREE depletion as the Troodos lavas from the Akaki Canyon and thus probably represent crystallized evolved melts from this tholeiitic series. The similarity in trace element compositions confirms that the main group plagiogranites largely represent melt compositions rather than cumulates and that they are closely related to the tholeiitic magma series forming the bulk of the upper crust (lavas and sheeted dykes) of the ophiolite. Consequently, the main group plagiogranite intrusions most likely represent crystallized shallow melt reservoirs that formed at or close to the spreading axis.

The Spilia intrusives show distinct incompatible element ratios (Fig. 4.8) resembling the boninitic lavas (Flower and Levine 1987; K¨onig et al. 2008) and late dykes (Coogan et al. 2003) found in the Troodos Ophiolite. This implies that these magmas also occurred in shallow magma reservoirs within the Troodos crust and occasionally evolved towards felsic melts (SiO2 60-79 wt%). Because the boninites are late volcanic products and the Spilia plutonic rocks also resemble some late mafic dykes (Coogan et al. 2003) in terms of (Ce/Yb)N and Sm/Nd (Fig. 4.8d), we conclude that the Spilia plagiogranite intrusion formed relatively late, which is supported by field evidence of minor crosscutting dykes, probably reflecting intrusion distant from a spreading axis.

The Zoopigi plagiogranite group shows a tendency towards lower values in Sm/Nd, Nb/Zr and higher in (Ce/Yb)N ratios compared to the two other groups, but overlaps with the main group samples (Fig. 4.8). However, the relative enrichment in (Ce/Yb)N but low 4.6. Discussion 73 ratios in Sm/Nd and Nb/Zr discriminate these plagiogranites from the other two groups. Based on the published data no comparable enriched volcanic lavas exist in Troodos. The Zoopigi plagiogranites represent felsic magmas intruded as dykes, probably during a relatively early stage because the silicic intrusions are cross-cut by many aphyric mafic and silicic dykes and show evidence of significant tectonic faulting (Fig. 4.2). This may indicate a formation close to the spreading axis when mafic dykes were still being actively intruded.

4.6.3 The relationship of the plagiogranitic intrusions to the mafic crustal plutonic rocks

Few geochemical data for the bulk composition of lower crustal plutonic rocks of the Troodos Ophiolite exist, because most authors have concentrated on clinopyroxene com- positions from mafic plutonic rocks. The low Ti contents in clinopyroxenes and gabbros

(Laurent and H´ebert 1989; H´ebert and Laurent 1990) are comparable to the low TiO2 contents of the boninitic lavas and the Spilia plagiogranite group (Fig. 4.5b). Addition- ally, Coogan et al. (2003) studied the Cr, Ti, V, Zr and REE in clinopyroxenes from Troodos mafic and ultramafic rocks and suggested that the lower crust acted as a filter for highly variable (and increasingly depleted) melts. Mixing and homogenisation of magmas prior to eruption results in much less variable lava composition, compared to the highly variable lower crustal rocks (Coogan et al. 2003). In order to determine the potential relationship of the Troodos plagiogranites to the mafic plutonic rocks, we calculated the melt compositions in equilibrium with these highly variable clinopyroxene cores using the REE composition of the minerals (Coogan et al. 2003) and Cpx/basalt distribution coefficients (Sobolev et al. 1996). The calculated liquids cover a wide range in terms of

(Ce/Yb)N vs. Sm/Nd (Fig. 4.8d), and cover the range defined by the main group (and the tholeiitic lavas), and the Spilia group plagiogranites (as well as the boninitic rocks and lavas). One calculated liquid from the gabbro clinopyroxenes also matches the LREE- enriched Zoopigi group. Additionally, Coogan et al. (2003) present data of very depleted late dykes resembling the Spilia plagiogranites and the boninitic lavas. We suggest that the main group and the Spilia plagiogranites have geochemically similar deep plutonic counterparts, which indicates that not only the tholeiitic magmas forming the bulk of the Troodos crust, but also the late boninite magmas have genetically related SiO2-rich magmas.

4.6.4 Generation of the plagiogranite magmas: fractional crystallization versus partial melting

Early models suggested that plagiogranitic melts formed by low-pressure fractional crys- tallization of a basaltic melt (Coleman and Peterman 1975), whereas recent experimental data indicate that silica-rich magmas can also form by hydrous partial melting of gabbros 4.6. Discussion 74

(Koepke et al. 2004; Koepke et al. 2007). We use the MELTS algorithm (Ghiorso and Sack 1995) to determine potential fractionation trends and the most likely conditions during fractionation of the main group plagiogranites. Laurent (1990) estimated the composition of a Troodos parental magma with about 13.5 wt% MgO (Mg#71.2) for the plutonic rocks by using the modal mineral contents and compositions in addition to the crystallization order of these minerals. In order to model the potential fractional crystallization trend

(for the tholeiitic lavas and the main group), we use an average of the lowest SiO2, FeO,

K2O, Na2O, TiO2 and the highest MgO, CaO and Al2O3 values of the Akaki canyon lavas (Regelous et al, in prep.) and add 7% olivine (observed frequency and cumulate composition (Laurent and H´ebert 1989)). The resulting primary mantle melt has an Mg# of 75, 11.0 wt% MgO and 48.9 wt% SiO2 (Table 4.2). Although a Mg# of 75 appears relatively high, it is in agreement with tholeiitic to boninitic Troodos lavas (Mg#78) from a study of Flower and Levine (1987), and observations by Coogan et al. (2003) who argued for a Troodos parental melt with relatively high Mg#, on the basis of clinopyroxenes from mafic plutonic rocks having elevated Mg# for a given Cr2O3 (compared to clinopyroxenes from MOR gabbros from Atlantic and Pacific spreading ridges).

During modelling, we vary the H2O content of the parental melt (0, 0.5, and 1.5 wt%) at constant f O2 conditions (QFM) and pressure (0.5 kbar). The best-fit model is obtained by a melt containing 0.5 wt% H2O (Table 4.2, Fig 4.5) which is comparable to basaltic glasses analysed by Muenow et al. (1990) and comply with a low K2O/H2O ratio (0.15-

0.16). Varying f O2 conditions (QFM, QFM+1, QFM-1) has only minor effects on the melt evolution according to the MELTS model and are not shown. Higher primary water content (1.5 wt%) delays crystallization of plagioclase and produces a steep increase in

Al2O3 content with decreasing MgO, whereas a dry parental melt (0% H2O) results in early decrease in Al2O3 that is not observed in the samples (Fig. 4.5c). In addition, a dry T parental melt model shows a strong increase in FeO and TiO2 unlike that in the Troodos plagiogranites (Fig. 4.5b, d). We suggest therefore that the main group plagiogranites are most likely the result of fractionation of a tholeiitic Mg-rich parental magma with about 0.5 wt% H2O under low-pressure conditions (about 0.5 kbar). The major elements of the Spilia group are modelled using a boninitic glass composition (Kapilio 2, Mg#70.4,

Table 4.2) as parental magma crystallizing at low pressure (0.5 kbar) with 0.5 wt% H2O and at QFM+1 (TiO2 and P2O5 are shown in Fig. 4.5b, f). Using higher water contents, higher pressures and different f O2 conditions (QFM, QFM-1, QFM+2) the model failed to reproduce the observed Spilia plagiogranite trends. Thus, we agree with Coogan et al. (2003) that highly variable parental melts ascend from the mantle into the crust and these melts fractionate under comparable low pressure and water conditions resulting in different plagiogranitic intrusions.

Koepke et al. (2007) compiled a minimum line of TiO2 content for evolved rocks formed by fractional crystallization of MORB, on the basis of experimental data (Dixon-Spulber and Rutherford 1983; Juster et al. 1989; Toplis and Carroll 1995; Thy et al. 1999; Berndt et 4.6. Discussion 75 al. 2005). Hydrous partial melting experiments of Koepke et al. (2004) produced silicic melts (>60 wt% SiO2) plotting below this minimum line, whereas the majority of our samples with SiO2 <68 wt% lie above the minimum line. Only the Spilia group has lower

TiO2 values and thus, an origin of these rocks by partial melting of gabbros is possible. However, the Spilia plagiogranites also show a fractionation trend with firstly increasing

TiO2 with increasing SiO2 (60-65 wt%) and then a decrease in TiO2 (SiO2 = 65-80 wt%), suggesting FeTi-oxide crystallization. Least-squares fractional crystallization calculations were carried out in a two-stage model and yield reasonable results (Table 4.3). The amounts of fractionated phases was determined first by the least squares method from a basalt (ET004-01-01) towards a basaltic andesite (ET013-060-1) and, for the second stage, from the basaltic andesite towards a tonalite (TroPal 19A). Crystallising phases of the first stage are clinopyroxene, plagioclase, and olivine, and during the second stage plagioclase, clinopyroxene, magnetite, olivine and ilmenite. This supports our model that crystal fractionation processes generated the various plagiogranite groups from parental mafic magmas.

The close relationship in terms of incompatible elements between the tholeiitic glasses (Figs. 4.7, 4.8) and the main group plagiogranites supports fractional crystallization as the main process leading to the formation of melts with >60 wt% SiO2. Furthermore, the existence of felsic melts with incompatible trace element signatures comparable to boninitic lavas (and late dykes) implies extreme and efficient crystal fractionation pro- cesses, because it is unlikely that such patterns can form by crustal partial melting. To test the fractional crystallization model further, we use the two-step model discussed above to model the predicted variation in incompatible elements Ce, Yb, Zr and Sm using published distribution coefficients for these elements (Table 4.3b, Fig. 4.10). Crystal fractionation from a basaltic Akaki Canyon glass sample (ET004-01-1) with relatively low REE concentration as parental magma, leads to slightly increasing (Ce/Yb)N and fits the observed compositional variation in the plagiogranites (Fig. 4.10). In contrast, partial melting (Table 4.4) of a hydrothermally altered sheeted dyke rock containing 40 % amphibole produces similar Ce, but significantly higher (Ce/Yb)N to the plagiogranites (Fig. 4.10). Additionally, partial melting of amphibolite facies sheeted dykes would lead to high Zr/Sm ratios that are observed only in a few plagiogranites; most samples have lower and constant Zr/Sm. Thus, both the major element and incompatible trace element variations are consistent with fractional crystallization of mafic melts as the dominant process in forming the plagiogranites. Although some xenoliths are present in the plagiogranite intrusions, assimilation of larger portions of partially molten wall- rock during the differentiation process would change the incompatible element ratios in the evolving magma which is not observed. Consequently, we suggest that no significant assimilation occurs during fractional crystallization of tholeiitic and boninitic parental to felsic melts in the Troodos Ophiolite. 4.6. Discussion 76

2.0 Zoopigi group a) 10% Spilia group Partial melting of Main group amphibolite-facies gabbro sheeted dykes 1.5 boninites 20% Sheeted dykes 30% N )

b 40% Y 1.0 50%

Ce/ Avge. sheeted dyke ( Volcanic glasses

0.5 Fractional crystallization trend

0.0 0 5 10 15 20 25 Ce (ppm) 60 b) Partial melting of 10% amphibolite-facies 50 sheeted dykes 20% 40 30% m S /

r 30 Volcanic glasses Z 50%

20 Fractional crystallization trend

10 Avge. sheeted dyke

0 0 50 100 150 200 Zr (ppm)

Figure 4.10: Fractional crystallization and partial melting models for incompatible elements a) (Ce/Yb)N versus Ce and b) Zr/Sm versus Zr. The fractional crystallization trends are similar to the approximately constant (Ce/Yb)N and Zr/Sm of the tholeiitic main group and glasses. The Spilia group has lower ratios comparable with the boninites (Flower and Levine 1987; K¨onig et al. 2010) and increasing Ce respectively Zr concentrations. The Zoopigi group overlap with the main group but increase towards higher ratios and concentrations. In contrast, the partial melting models indicate significant fractionation of the incompatible elements unlike the observed compositions of the rocks. 4.8. Acknowledgements 77

4.7 Conclusions

On the basis of the incompatible trace element compositions, the various plagiogranitic bodies occurring in the Troodos Ophiolite can be divided into three groups. Two of these are genetically linked to both the mafic lower crustal plutonic rocks, and the extrusive lavas. Fractional crystallization of tholeiitic melt produced the voluminous intrusions of the main group plagiogranites, whereas a single plagiogranite intrusion near Spilia were derived from a more depleted boninitic parental melt. In both cases, fractionation occurred in shallow magma lenses at relatively low water content. Partial melting of hydrothermally modified crustal rocks would fractionate incompatible elements which is not observed. As a result, we can rule out a significant role for partial melting during the formation of the Troodos plagiogranites, and for assimilation of wall rock during crystallisation and cooling. The third chemical group of plagiogranites were intruded in relatively thin dykes close to the Arakapas fault zone, and have no genetic relationship to any known crustal rocks exposed in the Troodos Ophiolite. Our data indicate that highly variable mantle melts ascend into the shallow magma lenses, but do not necessarily mix and homogenize as suggested by Coogan et al (2003). Rather, the mafic tholeiitic and boninitic melts ascend separately into the crust and stagnate over long periods of time, evolving to felsic melts now represented by the plagiogranite intrusions. The different plagiogranite groups reflect multiple melting events with subsequentially increasing depletion of the Troodos mantle source.

4.8 Acknowledgements

We acknowledge the help of U. Westernstr¨oer during trace element analyses. We thank Diplom students T. Endres, C. Weinzierl and H. Schmidt for the help during field work and H. Br¨atz and R. Klemd for timely and fast LA-ICP-MS analyses. We want to acknowledge A. Richter, M. Meyer and P. Brandl for loyal support during EMP problems. We want to express our thanks to M. Hertel for her perennial lab work and Jeanne and Florian for patience sawing the hard rocks. The manuscript benefitted from the very constructive reviews by L. Coogan and an anonymous reviewer and from comments by the editor J. Hoefs. S. Freund was funded by DFG grant HA2568/21. Tables 78

Table 4.1: Major and trace elements of Troodos plagiogranites and aphyric dykes (oxides are given in wt.%, elements in ppm).

Zoopigi group Tro Zoo 37A TroZoo 38c TroZoo 40 Tro Zoo 41 TroZoo 42 Tro Zoo 44 TroZoo 45 h Lat [N] 34°52.437 34°52.423 34°52.376 34°52.369 34°52.343 34°51.928 34°52.127 Long [E] 33°01.907 33°01.893 33°01.812 33°01.771 33°01.746 33°01.701 33°00.741 QAPF PG trondhjemite trondhjemite trondhjemite qz-granitoid PG qz-granitoid SiO2 77.2 76.5 76.9 75.8 75.4 75.9 77.2 TiO2 0.22 0.22 0.16 0.21 0.30 0.17 0.22 Al2O3 12.2 12.1 12.0 12.0 11.5 11.9 11.6 Fe2O3 1.62 2.53 2.85 3.27 3.94 3.36 2.82 MnO 0.02 0.02 0.01 0.03 0.02 0.05 0.02 MgO 0.27 0.35 0.23 0.22 0.75 0.27 0.46 CaO 4.42 3.27 2.27 2.37 3.72 2.71 2.82 Na2O 3.25 3.95 4.79 5.39 2.67 5.10 4.09 K2O 0.07 0.18 0.13 0.14 0.43 0.17 0.11 P2O5 0.06 0.04 0.04 0.05 0.07 0.03 0.05 LOI 0.68 0.85 0.51 0.54 1.15 0.41 0.54 Total 99.95 99.93 99.95 99.95 99.96 99.96 99.94 Li - 0.233 0.321 - 1.23 - 0.279 Sc - 14.9 13.0 - 15.5 - 12.9 V 15.1 14.7 5.21 6.50 10.3 8.60 9.8 Cr LOD 6.08 5.17 LOD 5.73 LOD 7.92 Co - 2.9 0.730 - 5.0 - 3.2 Ni 2.10 5.26 2.91 LOD 3.96 LOD 4.7 Cu 6.50 10.9 2.28 LOD 20.4 13.3 5.23 Zn 5.80 6.79 8.51 6.60 6.08 11.3 12.2 Ga - 13.2 12.0 6.30 13.1 - 14.3 Rb LOD 0.866 0.533 LOD 2.71 LOD 0.628 Sr 184 153 140 LOD 160 117 124 Y 43.1 50.7 46.8 107 36.4 32.1 49.3 Zr (XRF) 103 141 73.0 48.8 64.3 69.3 139 Nb 1.80 1.78 1.79 107 1.38 1.50 2.98 Sn - 0.266 0.178 2.20 0.553 - 0.421 Ba - 23.4 21.1 - 31.2 - 52.1 La - 8.49 4.04 - 6.25 - 6.48 Ce - 20.2 10.7 - 15.3 - 16.9 Pr - 2.90 1.75 - 2.28 - 2.64 Nd - 14.3 9.66 - 11.5 - 13.9 Sm - 4.67 3.73 - 3.75 - 4.86 Eu - 1.17 1.00 - 1.18 - 1.30 Gd - 6.28 5.42 - 4.93 - 6.67 Tb - 1.19 1.06 - 0.905 - 1.22 Dy - 8.23 7.48 - 6.09 - 8.23 Ho - 1.80 1.65 - 1.29 - 1.75 Er - 5.18 4.77 - 3.61 - 4.90 Tm - 0.790 0.726 - 0.543 - 0.732 Yb - 5.27 4.70 - 3.61 - 4.83 Lu - 0.760 0.658 - 0.527 - 0.702 Ta - 0.105 0.107 - 0.075 - 0.186 Pb 1.30 0.414 1.88 1.10 0.286 2.00 0.423 Th - 1.33 0.549 - 0.691 - 0.877 U - 0.191 0.149 - 0.156 - 0.303 Tables 79

Table 4.1: continued.

Spilia group TroZoo 45 d TroSpi 24 TroSpi 25 TroSpi 26 TroSpi 27 TroSpi 28 TroSpi 29 Lat [N] 34°52.127 34°57.933 34°57.921 34°58.015 34°58.264 34°57.713 34°57.018 Long [E] 33°00.741 32°57.470 32°57.558 32°57.608 32°57.785 32°57.786 32°57.604 QAPF qz-granitoid qz-granitoid qz-granitoid PG PG PG PG SiO2 67.8 71.1 75.3 74.5 75.6 73.9 75.0 TiO2 0.52 0.49 0.24 0.25 0.25 0.26 0.24 Al2O3 14.1 11.6 11.3 11.5 11.3 11.3 10.9 Fe2O3 6.02 6.87 4.88 4.90 4.44 3.96 4.65 MnO 0.06 0.08 0.03 0.04 0.03 0.03 0.04 MgO 1.78 1.22 0.49 0.53 0.63 0.87 1.30 CaO 3.95 5.23 2.84 3.70 3.18 5.57 3.59 Na2O 4.51 1.60 3.52 2.60 3.55 0.98 1.12 K2O 0.13 0.47 0.28 0.40 0.28 0.92 0.58 P2O5 0.11 0.06 0.06 0.06 0.06 0.06 0.06 LOI 1.01 1.30 1.06 1.43 0.70 2.11 2.52 Total 99.94 99.97 99.97 99.96 99.96 99.96 99.96 Li 0.691 - 1.19 ---- Sc 21.1 - 15.3 ---- V 65.6 54.2 1.85 15.6 12.5 17.2 14.5 Cr 3.77 LOD 4.30 LOD LOD LOD LOD Co 9.7 - 3.59 ---- Ni 3.03 LOD 2.45 LOD LOD LOD LOD Cu 18.4 7.80 1.61 9.40 5.90 LOD 17.4 Zn 32.9 18.4 18.0 12.1 9.70 LOD 8.00 Ga 16.6 - 12.5 ---- Rb 1.19 1.60 1.53 1.50 1.20 3.80 2.30 Sr 123 94.1 85.7 109 90.9 153 96.3 Y 57.1 14.0 40.7 31.6 35.1 44.8 36.2 Zr (XRF) 103 16.2 44.0 38.8 62.9 43.2 40.7 Nb 2.27 2.0 3.94 3.20 3.80 3.40 4.40 Sn 0.383 - 0.419 ---- Ba 68.4 - 28.48 ---- La 5.90 - 2.57 ---- Ce 15.3 - 6.50 ---- Pr 2.52 - 1.01 ---- Nd 13.8 - 5.83 ---- Sm 5.02 - 2.60 ---- Eu 1.03 - 0.642 ---- Gd 7.14 - 4.25 ---- Tb 1.36 - 0.851 ---- Dy 9.52 - 6.17 ---- Ho 2.09 - 1.42 ---- Er 5.95 - 4.33 ---- Tm 0.907 - 0.713 ---- Yb 6.05 - 5.06 ---- Lu 0.892 - 0.770 ---- Ta 0.119 - 0.224 ---- Pb 0.440 2.20 0.431 1.60 2.70 LOD 2.10 Th 0.741 - 0.478 ---- U 0.185 - 0.288 ---- Tables 80

Table 4.1: continued.

TroSpi 30 TroSpi 31 TroSpi 32 TroSpi 33 TroSpi 34 TroSpi 35 TroSpi 36A Lat [N] 34°57.051 34°57.692 34°57.672 34°57.699 34°57.789 34°57.875 34°57.990 Long [E] 32°57.673 32°57.932 32°57.811 32°57.782 32°57.748 32°57.844 32°57.778 QAPF trondhjemite qz-granitoid qz-granitoid qz-granitoid PG trondhjemite qz-granitoid SiO2 75.1 75.1 79.1 75.1 75.2 59.9 75.0 TiO2 0.22 0.25 0.13 0.23 0.23 0.44 0.24 Al2O3 10.6 11.0 11.4 10.9 11.1 15.0 11.2 Fe2O3 5.25 4.35 1.74 4.12 4.48 9.27 4.09 MnO 0.04 0.01 0.01 0.02 0.03 0.14 0.04 MgO 0.66 1.40 0.21 1.45 0.36 3.86 1.12 CaO 4.41 4.43 1.50 3.57 5.11 6.90 3.68 Na2O 1.39 0.85 4.97 1.23 1.76 2.34 2.46 K2O 0.54 0.32 0.15 0.56 0.44 0.32 0.45 P2O5 0.05 0.07 0.02 0.05 0.07 0.03 0.08 LOI 1.68 2.23 0.65 2.81 1.21 1.77 1.53 Total 99.97 99.97 99.92 99.96 99.96 99.93 99.97 Li - 12.2 - 23.0 - 10.4 2.19 Sc - 19.2 - 15.6 - 35.6 20.9 V 13.2 12.7 6.30 2.33 20.1 239 6.45 Cr LOD 4.58 LOD 4.91 LOD 2.22 8.37 Co - 14.8 - 4.77 - 27.9 5.72 Ni LOD 3.05 1.90 2.62 LOD 10.4 5.47 Cu 7.40 7.82 4.80 0.476 8.30 2.95 1.00 Zn 6.50 2.73 2.90 3.01 1.50 36.5 12.2 Ga - 11.6 - 12.5 - 15.0 11.7 Rb 2.60 2.12 0.700 3.52 1.50 1.88 2.53 Sr 116 103 71.8 125 112 75.3 95.7 Y 30.5 35.0 122 45.0 23.9 31.1 34.0 Zr (XRF) 37.3 35.7 304 55.4 43.5 27.4 47.2 Nb 3.70 2.77 5.10 3.56 4.30 1.85 3.36 Sn - 0.275 - 0.223 - 0.168 0.252 Ba - 1.47 - 4.91 - 29.10 4.05 La - 2.81 - 1.79 - 1.33 1.90 Ce - 6.43 - 4.86 - 3.22 4.84 Pr - 0.95 - 0.809 - 0.519 0.758 Nd - 5.20 - 4.95 - 3.06 4.38 Sm - 2.26 - 2.49 - 1.48 2.02 Eu - 0.53 - 0.718 - 0.421 0.632 Gd - 3.69 - 4.42 - 2.66 3.40 Tb - 0.74 - 0.922 - 0.551 0.691 Dy - 5.42 - 6.85 - 4.09 5.10 Ho - 1.24 - 1.57 - 0.969 1.17 Er - 3.69 - 4.68 - 2.90 3.49 Tm - 0.576 - 0.737 - 0.459 0.565 Yb - 3.90 - 5.01 - 3.09 3.95 Lu - 0.581 - 0.747 - 0.474 0.617 Ta - 0.166 - 0.229 - 0.148 0.211 Pb 1.90 0.057 2.1 0.129 0.80 0.784 0.193 Th - 0.336 - 0.350 - 0.296 0.348 U - 0.133 - 0.225 - 0.195 0.266 Tables 81

Table 4.1: continued.

Gabbro Main group TroSpi 36B TroPed 51 TroPal 1 TroPal 3B TroPal 5A TroPal 6 TroPal 7A Lat [N] 34°57.990 34°57.599 34°57.112 34°54.974 34°54.858 34°54.865 34°54.864 Long [E] 32°57.778 32°49.626 33°02.431 33°05.442 33°05.057 33°05.048 33°05.046 QAPF tonalit gabbro PG trondhjemite trondhjemite trondhjemite qz-granitoid SiO2 63.6 49.1 68.8 70.0 55.9 67.0 69.6 TiO2 0.53 0.20 0.65 0.61 0.99 0.74 0.61 Al2O3 13.0 18.2 13.4 13.2 14.8 14.5 12.9 Fe2O3 9.31 7.68 6.87 6.08 13.6 6.76 6.74 MnO 0.11 0.14 0.09 0.05 0.11 0.09 0.08 MgO 2.88 9.03 1.22 1.02 3.07 1.30 1.20 CaO 6.13 13.8 4.84 4.28 6.51 5.64 4.52 Na2O 1.82 0.80 3.09 3.69 3.57 3.21 3.33 K2O 0.49 0.06 0.15 0.20 0.09 0.11 0.22 P2O5 0.02 0.02 0.11 0.12 0.08 0.15 0.11 LOI 2.04 0.91 0.74 0.69 1.24 0.51 0.65 Total 99.94 99.91 99.95 99.95 99.91 99.95 99.95 Li 6.71 0.33 - 1.17 -- 0.78 Sc 35.4 39.4 - 17.0 -- 17.6 V 152 176 41.7 22.6 35.0 45.8 36.6 Cr 2.15 80.3 LOD 4.42 LOD LOD 4.95 Co 17.8 37.8 - 8.29 -- 9.94 Ni 1.68 74.0 LOD 3.21 LOD LOD 3.66 Cu 1.03 23.9 LOD 3.05 19.4 8.60 2.24 Zn 21.4 39.1 5.60 11.4 5.20 7.30 11.5 Ga 13.3 12.9 - 15.6 -- 14.7 Rb 2.82 0.441 LOD 0.969 0.60 LOD 1.13 Sr 95.8 73.8 121 129 127 122 128 Y 23.6 6.2 40.2 40.5 47.2 39.3 39.3 Zr (XRF) 15.5 0.60 84.6 94.0 112 78.7 85.0 Nb 1.86 0.255 3.40 2.13 3.70 2.50 2.12 Sn 0.431 0.096 - 0.302 -- 0.325 Ba 5.26 2.38 - 14.5 -- 8.56 La 1.01 0.271 - 4.08 -- 3.91 Ce 2.71 0.951 - 11.54 -- 11.04 Pr 0.436 0.178 - 1.90 -- 1.81 Nd 2.59 1.03 - 10.4 -- 9.70 Sm 1.31 0.456 - 3.76 -- 3.47 Eu 0.431 0.253 - 1.17 -- 1.07 Gd 2.29 0.721 - 5.21 -- 4.89 Tb 0.479 0.141 - 0.981 -- 0.928 Dy 3.60 1.02 - 6.72 -- 6.44 Ho 0.84 0.230 - 1.46 -- 1.40 Er 2.53 0.655 - 4.15 -- 3.99 Tm 0.404 0.103 - 0.631 -- 0.606 Yb 2.80 0.70 - 4.20 -- 4.08 Lu 0.424 0.106 - 0.616 -- 0.604 Ta 0.132 0.047 - 0.136 -- 0.135 Pb 0.210 0.194 1.70 0.118 1.90 1.40 0.119 Th 0.219 0.030 - 0.431 -- 0.406 U 0.106 0.016 - 0.136 -- 0.148 Tables 82

Table 4.1: continued.

TroPal 9 TroPal 11 TroPal 12 Tro Pal 14A TroPal 14B h TroPal 14B d Tro Pal 15 Lat [N] 34°54.908 34°54.827 34°54.813 34°54.833 34°54.833 34°54.833 34°54.850 Long [E] 33°05.304 33°04.954 33°04.954 33°04.822 33°04.822 33°04.822 33°04.696 QAPF qz-granitoid trondhjemite PG PG qz-granitoid trondhjemite trondhjemite SiO2 68.8 69.7 68.9 65.1 68.5 62.5 67.9 TiO2 0.66 0.72 0.66 0.95 0.61 0.97 0.71 Al2O3 13.3 13.6 13.5 13.9 13.5 14.0 13.0 Fe2O3 7.00 5.46 6.47 8.61 6.36 9.89 7.24 MnO 0.04 0.08 0.09 0.09 0.05 0.09 0.09 MgO 1.25 1.03 1.02 1.52 1.21 1.99 1.30 CaO 4.70 5.15 5.03 6.16 4.99 6.24 5.40 Na2O 3.24 3.14 3.55 2.88 3.19 2.32 3.13 K2O 0.18 0.17 0.09 0.11 0.36 0.44 0.23 P2O5 0.10 0.15 0.16 0.12 0.16 0.10 0.15 LOI 0.74 0.77 0.46 0.48 1.06 1.37 0.81 Total 99.95 99.95 99.95 99.94 99.95 99.93 99.97 Li ---- 2.30 3.91 - Sc ---- 17.5 31.6 - V 25.2 31.3 20.5 72.4 13.5 86.3 27.8 Cr LOD LOD LOD LOD 6.64 12.2 LOD Co ---- 10.1 17.0 - Ni 4.60 LOD 38.3 1.90 4.27 7.97 LOD Cu 10.4 30.5 3.50 10.3 2.60 9.66 6.60 Zn 4.00 7.40 9.20 7.80 9.93 14.2 11.5 Ga ---- 14.9 15.2 - Rb 0.90 LOD LOD 1.00 1.62 2.03 1.1 Sr 124 134 117 128 119 114 127 Y 49.1 31.9 44.7 33.6 33.9 51.3 33.6 Zr (XRF) 96.1 68.9 74.5 61.5 75.7 76.9 47.8 Nb 4.00 2.90 3.80 2.60 1.87 2.76 2.50 Sn ---- 0.467 0.818 - Ba ---- 11.9 9.20 - La ---- 3.22 3.32 - Ce ---- 8.99 9.88 - Pr ---- 1.50 1.72 - Nd ---- 8.27 9.83 - Sm ---- 3.03 4.02 - Eu ---- 0.879 0.868 - Gd ---- 4.33 6.03 - Tb ---- 0.802 1.19 - Dy ---- 5.51 8.42 - Ho ---- 1.20 1.85 - Er ---- 3.43 5.34 - Tm ---- 0.518 0.822 - Yb ---- 3.44 5.53 - Lu ---- 0.509 0.838 - Ta ---- 0.120 0.168 - Pb 1.70 3.30 1.90 1.90 0.124 0.179 LOD Th ---- 0.352 0.406 - U ---- 0.143 0.170 - Tables 83

Table 4.1: continued.

Tro Pal 16 TroPal 18 TroPal 19A Tro Pal 20 CY05.05.-02 CY-05.05-03 CY05.05.-04 Lat [N] 34°54.847 34°54.858 34°54.903 34°55.018 34°55.490 34°55.648 34°55.662 Long [E] 33°04.664 33°04.419 33°04.328 33°04.247 33°05.330 33°4.859 33°4.835 QAPF trondhjemite trondhjemite trondhjemite trondhjemite trondhjemite tonalit tonalit SiO2 63.7 70.8 73.9 67.3 72.2 65.9 69.8 TiO2 0.81 0.51 0.47 0.65 0.48 0.88 0.61 Al2O3 14.3 12.8 12.7 13.1 12.3 13.0 12.9 Fe2O3 9.37 6.28 2.75 9.27 5.44 8.61 6.88 MnO 0.15 0.06 0.04 0.03 0.04 0.09 0.06 MgO 1.78 0.71 0.65 0.93 0.56 1.15 0.70 CaO 5.75 4.00 4.55 4.40 4.42 5.21 5.32 Na2O 3.12 3.93 3.66 3.07 3.48 3.19 3.01 K2O 0.21 0.19 0.17 0.21 0.19 0.14 0.08 P2O5 0.11 0.12 0.12 0.15 0.11 0.11 0.14 LOI 0.60 0.58 0.96 0.85 0.72 1.76 0.49 Total 99.94 99.95 99.94 99.96 99.97 99.96 99.95 Li - 0.40 0.65 - 0.914 - 0.307 Sc - 17.8 15.8 - 9.76 - 18.0 V 143 8.44 38.1 LOD 1.89 LOD 3.89 Cr LOD 3.89 6.35 3.80 3.90 LOD 4.37 Co - 6.26 7.34 - 8.01 - 9.00 Ni 6.00 2.41 4.62 3.50 0.76 LOD 1.28 Cu 16.7 1.88 2.63 11.0 0.91 LOD 0.75 Zn 38.0 22.2 9.03 12.3 10.6 47.2 19.6 Ga - 15.5 14.0 - 16.3 - 15.9 Rb LOD 1.15 0.91 LOD 1.22 LOD 0.652 Sr 97.4 109 144 138 112 158 137 Y 38.1 47.9 33.1 42.6 21.0 39.1 40.9 Zr (XRF) 58.0 94.2 77.3 71.7 49.5 68.7 79.5 Nb 3.70 2.47 2.01 4.40 0.968 4.00 1.55 Sn - 0.678 0.635 - 0.379 - 0.523 Ba - 46.6 7.70 - 6.24 - 17.9 La - 4.04 2.88 - 1.29 - 3.99 Ce - 11.76 7.71 - 4.22 - 10.7 Pr - 1.98 1.27 - 0.759 - 1.78 Nd - 10.8 6.98 - 4.28 - 9.79 Sm - 4.06 2.60 - 1.67 - 3.62 Eu - 1.19 1.20 - 1.07 - 1.10 Gd - 5.83 3.93 - 2.48 - 5.22 Tb - 1.12 0.75 - 0.476 - 0.982 Dy - 7.80 5.28 - 3.37 - 6.82 Ho - 1.71 1.16 - 0.755 - 1.49 Er - 4.90 3.31 - 2.20 - 4.25 Tm - 0.750 0.500 - 0.346 - 0.649 Yb - 5.01 3.33 - 2.36 - 4.32 Lu - 0.742 0.487 - 0.359 - 0.641 Ta - 0.160 0.128 - 0.068 - 0.115 Pb 1.60 0.278 0.164 - 0.167 0.50 0.155 Th - 0.419 0.303 2.50 0.175 - 0.451 U - 0.177 0.135 - 0.082 - 0.181 Tables 84

Table 4.1: continued.

CY10.05.-01 CY10.05.-04 CY10.05.-05 CY-11.05.-01 CY11.05.-04 CY.12.05.-01 CY-12.05.-02 Lat [N] 34°54.446 34°56.618 34°54.397 34°54.904 34°54.910 34°54.856 34°54.871 Long [E] 33°05.132 33°05.379 33°05.085 33°5.284 33°05.313 33°5.065 33°5.037 QAPF PG tonalit tonalit PG PG PG tonalit SiO2 71.7 59.8 61.8 69.6 69.0 71.7 65.6 TiO2 0.65 1.11 0.96 0.72 0.70 0.55 0.76 Al2O3 11.6 13.8 13.9 13.3 13.4 12.7 13.8 Fe2O3 6.97 12.0 10.3 6.25 6.81 5.27 6.10 MnO 0.04 0.17 0.13 0.03 0.03 0.05 0.09 MgO 0.79 2.67 2.18 1.04 0.99 0.88 1.57 CaO 3.79 4.91 5.39 6.64 5.01 3.18 5.71 Na2O 3.25 3.68 3.75 1.26 3.23 4.29 3.54 K2O 0.14 0.39 0.29 0.05 0.13 0.21 0.58 P2O5 0.11 0.10 0.08 0.12 0.11 0.11 0.11 LOI 0.92 1.27 1.12 0.99 0.48 0.98 2.08 Total 99.94 99.92 99.91 99.95 99.94 99.96 99.95 Li 0.642 1.32 1.36 - 1.46 -- Sc 18.2 29.1 26.3 - 17.4 -- V 17.8 261 192 36.2 18.6 26.3 84.4 Cr 8.65 4.76 5.68 LOD 4.27 LOD LOD Co 7.00 22.6 20.1 - 8.96 -- Ni 6.53 6.61 3.61 2.7 0.21 LOD LOD Cu 1.93 1.35 1.16 LOD 0.21 LOD 7.1 Zn 11.2 63.9 50.0 9.9 10.4 7.9 3.9 Ga 11.9 14.5 15.3 - 16.4 -- Rb 1.33 2.23 1.73 LOD 0.802 0.50 2.30 Sr 138 139 138 127 127 113 154 Y 25.4 30.9 30.7 41.7 40.6 42.5 42.1 Zr (XRF) 66.7 40.0 44.3 94.4 89.3 108 77.3 Nb 1.84 2.11 2.11 2.40 2.14 4.20 3.10 Sn 0.220 0.569 0.490 - 0.454 -- Ba 11.7 15.9 18.3 - 13.2 -- La 1.77 1.62 1.59 - 3.30 -- Ce 5.20 5.91 5.67 - 10.4 -- Pr 0.878 1.08 1.04 - 1.77 -- Nd 4.96 6.22 5.96 - 9.72 -- Sm 1.92 2.42 2.38 - 3.54 -- Eu 0.806 0.800 0.918 - 1.17 -- Gd 2.95 3.68 3.61 - 5.06 -- Tb 0.571 0.711 0.706 - 0.963 -- Dy 4.08 5.09 5.03 - 6.78 -- Ho 0.911 1.12 1.12 - 1.47 -- Er 2.64 3.25 3.26 - 4.25 -- Tm 0.409 0.504 0.503 - 0.653 -- Yb 2.82 3.41 3.43 - 4.37 -- Lu 0.432 0.512 0.523 - 0.641 -- Ta 0.130 0.150 0.151 - 0.138 -- Pb 0.206 0.235 0.192 1.40 0.133 0.50 1.90 Th 0.370 0.253 0.277 - 0.448 -- U 0.142 0.133 0.151 - 0.195 -- Tables 85

Table 4.1: continued.

CY-12.05.-03 Cy-12.05.-04 CY-12.05.-05 CY-12.05.-09 CY-14.05.-01.1CY-14.05.-01.2CY-16.05.-01 Lat [N] 34°54.871 34°54.873 34°54.873 34°54.862 34°54.822 34°54.822 34°54.793 Long [E] 33°5.027 33°5.030 33°5.030 33°5.050 33°4.695 33°4.695 33°4.861 QAPF PG PG PG PG trondhjemite tonalit PG SiO2 73.5 66.3 72.7 70.8 71.5 62.4 61.4 TiO2 0.38 0.7 0.35 0.58 0.41 0.87 0.95 Al2O3 12.2 14.3 12.2 12.9 12.9 14.2 14.2 Fe2O3 6.13 7.58 5.89 5.63 5.29 8.72 9.63 MnO 0.03 0.10 0.03 0.09 0.04 0.12 0.11 MgO 0.44 1.54 0.37 0.96 0.74 2.47 2.50 CaO 2.93 6.11 3.08 4.29 3.97 5.72 6.21 Na2O 4.02 3.06 4.20 3.72 3.94 4.18 3.65 K2O 0.11 0.07 0.06 0.18 0.30 0.27 0.25 P2O5 0.07 0.10 0.08 0.13 0.12 0.08 0.08 LOI 0.13 0.87 0.98 0.72 0.75 0.88 0.97 Total 99.95 100.8 99.96 99.96 99.95 99.93 99.93 Li ------Sc ------V LOD 88.0 11.8 20.4 14.0 197 184 Cr LOD 7.0 LOD LOD LOD LOD LOD Co ------Ni LOD 2.0 LOD LOD LOD 5.4 8.5 Cu 5.8 1.0 3.0 13.9 43.5 5.3 15.8 Zn 6.6 14.0 2.7 11.3 8.6 14.2 10.1 Ga ------Rb LOD LOD LOD LOD 1.70 0.70 0.40 Sr 111 126 112 121 133 115 125 Y 48.0 42.0 46.7 41.9 39.2 39.8 33.5 Zr (XRF) 95.9 79.0 93.8 86.3 93.2 68.4 43.4 Nb 2.30 LOD 4.20 3.30 2.50 2.70 3.40 Sn ------Ba ------La ------Ce ------Pr ------Nd ------Sm ------Eu ------Gd ------Tb ------Dy ------Ho ------Er ------Tm ------Yb ------Lu ------Ta ------Pb 1.50 3.00 1.20 3.90 1.00 1.20 0.50 Th ------U ------Tables 86

Table 4.1: continued.

CY-16.05.-02 CY-16.05.-04 CY-16.05.-05 CY-16.05.-06 CY-16.05.-07 CY-17.05.-06.1CY-17.05.-06.2 Lat [N] 34°54.793 34°54.793 34°54.793 34°54.793 34°54.793 34°54.969 34°54.969 Long [E] 33°4.861 33°4.861 33°4.861 33°4.861 33°4.861 33°5.552 33°5.552 QAPF PG PG PG PG PG PG PG SiO2 66.6 61.1 72.7 61.6 60.2 72.3 71.1 TiO2 0.77 0.97 0.37 1.12 1.12 0.53 0.61 Al2O3 13.6 13.9 12.0 14.7 14.6 12.8 13.5 Fe2O3 7.71 9.53 6.04 8.15 10.2 4.05 3.36 MnO 0.08 0.13 0.01 0.13 0.11 0.05 0.06 MgO 1.39 2.46 0.71 2.72 2.72 1.04 1.39 CaO 4.17 6.02 3.08 6.52 6.15 4.43 4.88 Na2O 4.31 4.44 2.91 3.84 3.60 3.51 3.72 K2O 0.28 0.30 0.74 0.32 0.26 0.11 0.15 P2O5 0.14 0.09 0.08 0.07 0.08 0.13 0.14 LOI 0.92 1.06 1.37 0.74 0.87 1.11 1.06 Total 99.95 99.93 99.96 99.92 99.93 99.95 99.95 Li ------Sc ------V LOD 137 12.2 168 175 50.9 57.5 Cr LOD LOD LOD LOD LOD LOD LOD Co ------Ni LOD LOD LOD LOD LOD 3.2 LOD Cu 17.2 81.2 LOD LOD 4.0 30.1 36.1 Zn 10.1 15.1 2.3 17.0 6.6 3.6 4.7 Ga ------Rb 0.80 1.40 1.70 0.80 1.20 LOD LOD Sr 117 136 109 155 126 150 144 Y 50.8 35.1 47.8 44.6 29.2 19.1 24.6 Zr (XRF) 105 70.3 92.1 63.5 50.4 72.1 85.3 Nb 3.40 4.80 3.60 3.70 3.20 4.30 4.00 Sn ------Ba ------La ------Ce ------Pr ------Nd ------Sm ------Eu ------Gd ------Tb ------Dy ------Ho ------Er ------Tm ------Yb ------Lu ------Ta ------Pb 1.00 0.90 - 0.70 0.50 1.10 1.30 Th ------U ------Tables 87

Table 4.1: continued.

CY-25.05.-01 CY-26.05.-01 CY-17.05.-08 CY31.05.-02 Tro Am 23A Tro Am 23B TroFte 46 Lat [N] 34°54.973 34°56.520 34°54.910 34°54.675 34°54.094 34°54.094 34°56.717 Long [E] 33°5.451 33°4.320 33°4.327 33°06.691 32°56.530 32°56.530 33°03.426 QAPF PG PG PG PG PG trondhjemite trondhjemite SiO2 58.9 67.1 71.2 63.2 76.1 70.5 71.0 TiO2 0.96 0.58 0.47 0.97 0.29 0.45 0.54 Al2O3 14.3 11.4 12.3 13.6 11.4 13.2 12.5 Fe2O3 10.0 9.87 6.43 9.08 2.26 5.44 6.02 MnO 0.07 0.12 0.04 0.09 0.02 0.03 0.02 MgO 3.16 1.47 0.59 2.12 0.46 0.91 0.88 CaO 7.59 1.91 3.88 4.59 6.34 5.81 2.34 Na2O 2.64 5.65 4.03 4.69 1.64 2.82 5.23 K2O 0.09 0.08 0.17 0.24 0.24 0.12 0.19 P2O5 0.08 0.05 0.12 0.10 0.07 0.13 0.11 LOI 2.09 1.76 0.70 1.24 1.15 0.53 1.21 Total 99.91 99.98 99.96 99.92 99.97 99.96 99.96 Li --- 0.556 --- Sc --- 25.3 --- V 270 16.1 LOD 197 0.30 13.6 20.2 Cr 1.9 LOD LOD 10.2 LOD 0.80 LOD Co --- 14.0 --- Ni 9.8 LOD LOD 7.99 LOD 5.10 2.90 Cu 74.3 12.9 14.2 2.30 11.1 20.0 10.7 Zn 2.7 60.5 5.6 23.3 LOD 4.00 6.20 Ga --- 13.3 --- Rb LOD LOD 1.30 1.46 0.80 LOD 0.50 Sr 134 58.3 135 182 113 119 88.7 Y 23.8 37.1 36.3 28.5 44.8 41.9 39.9 Zr (XRF) 45.3 20.9 72.0 68.0 70.8 82.4 66.3 Nb 2.20 2.20 3.30 1.89 3.10 2.00 3.40 Sn --- 0.645 --- Ba --- 36.6 --- La --- 2.34 --- Ce --- 7.68 --- Pr --- 1.31 --- Nd --- 7.19 --- Sm --- 2.61 --- Eu --- 0.845 --- Gd --- 3.64 --- Tb --- 0.680 --- Dy --- 4.71 --- Ho --- 1.02 --- Er --- 2.94 --- Tm --- 0.451 --- Yb --- 3.05 --- Lu --- 0.456 --- Ta --- 0.128 --- Pb 0.80 1.90 2.10 0.228 1.10 1.00 1.60 Th --- 0.509 --- U --- 0.177 --- Tables 88

Table 4.1: continued.

TroFte 47 TroFte 48 TroLem 49 Tro Ped 50 Tro Am 54 Tro Am 55 Tro Am 56 Lat [N] 34°56.795 34°56.839 34°57.162 34°58.166 34°54.675 34°54.483 34°54.521 Long [E] 33°03.381 33°03.431 32°48.732 32°49.447 32°57.372 32°57.603 32°57.494 QAPF qz-granitoid qz-granitoid qz-granitoid PG trondhjemite qz-granitoid trondhjemite SiO2 70.9 73.9 73.3 68.1 64.8 74.0 73.9 TiO2 0.50 0.33 0.50 0.58 0.80 0.31 0.32 Al2O3 11.7 11.0 12.9 12.6 14.6 12.2 12.3 Fe2O3 7.42 6.39 1.71 6.53 7.94 5.13 4.97 MnO 0.04 0.02 0.01 0.03 0.07 0.02 0.05 MgO 0.82 0.44 1.97 1.54 1.30 0.27 0.29 CaO 3.76 2.71 5.90 5.34 6.22 3.12 3.19 Na2O 3.20 3.76 1.10 1.61 3.49 4.13 4.15 K2O 0.21 0.20 0.33 0.73 0.12 0.16 0.11 P2O5 0.08 0.08 0.14 0.11 0.11 0.07 0.07 LOI 1.33 1.11 2.20 2.74 0.50 0.60 0.60 Total 99.95 99.97 99.95 99.95 99.94 99.96 99.95 Li ------Sc ------V 21.0 9.60 23.5 LOD 41.1 1.10 11.4 Cr LOD LOD LOD LOD LOD LOD LOD Co ------Ni LOD 2.80 LOD LOD 4.20 3.20 LOD Cu 15.2 16.0 LOD 7.50 38.3 14.8 20.6 Zn 8.80 5.40 LOD LOD 18.9 3.30 14.4 Ga ------Rb 0.40 LOD 1.80 3.30 0.50 LOD LOD Sr 103 110 111 143 131 103 86.5 Y 34.6 37.6 46.0 42.5 43.2 45.9 40.2 Zr (XRF) 34.6 48.1 87.0 74.4 71.1 93.0 91.4 Nb 3.40 4.30 4.10 2.00 3.30 2.20 2.80 Sn ------Ba ------La ------Ce ------Pr ------Nd ------Sm ------Eu ------Gd ------Tb ------Dy ------Ho ------Er ------Tm ------Yb ------Lu ------Ta ------Pb 0.30 1.60 1.90 1.10 1.80 1.00 0.70 Th ------U ------Tables 89

Table 4.1: continued.

aphyric dikes Tro Am 57 Tro Pla 58A TroPal 3A TroPal 5B TroPal 7B TroPal 19B CY-27.05.-01 Lat [N] 34°54.684 34°56.301 34°54.974 34°54.858 34°54.864 34°54.903 34°54.868 Long [E] 32°57.355 33°02.550 33°05.442 33°05.057 33°05.046 33°04.328 33°4.937 QAPF PG qz-granitoid andesit dacit andesit andesit dacit SiO2 61.5 75.9 59.6 70.4 60.3 60.2 70.0 TiO2 0.92 0.51 0.72 0.56 1.06 1.01 0.51 Al2O3 15.2 12.2 14.7 13.0 13.3 14.0 12.1 Fe2O3 8.39 1.20 9.06 6.15 10.8 8.61 7.02 MnO 0.09 0.02 0.11 0.06 0.11 0.12 0.03 MgO 2.10 1.18 4.20 0.94 3.21 2.91 1.24 CaO 7.66 6.40 6.77 4.31 5.29 6.32 4.06 Na2O 3.43 1.25 2.10 3.51 3.82 5.43 1.87 K2O 0.09 0.16 0.59 0.15 0.27 0.09 0.66 P2O5 0.13 0.12 0.07 0.11 0.08 0.08 0.10 LOI 0.45 1.00 1.96 0.79 1.61 1.12 2.38 Total 99.94 99.94 99.90 99.95 99.94 99.92 99.96 Li -- 5.31 - 1.98 0.335 - Sc -- 31.5 - 32.5 30.3 - V 142 52.8 267 230 242 274 26.5 Cr LOD LOD 10.7 LOD 2.39 7.47 LOD Co -- 26.0 - 26.2 26.0 - Ni 9.70 LOD 16.7 9.9 4.02 9.24 2.0 Cu 9.20 LOD 4.3 21.7 1.7 2.7 LOD Zn 22.0 LOD 23.6 9.80 14.4 37.2 3.1 Ga -- 16.8 - 15.4 16.2 - Rb LOD LOD 2.40 LOD 1.41 0.243 2.70 Sr 130 114 108 123 114 100 107 Y 29.4 43.9 27.1 41.2 28.4 31.6 43.3 Zr (XRF) 47.4 183 36.0 44.0 44.7 47.2 64.8 Nb 1.70 3.30 1.26 4.00 1.33 1.20 3.40 Sn -- 0.363 - 0.729 0.614 - Ba -- 13.2 - 6.50 6.29 - La -- 2.05 - 2.05 2.03 - Ce -- 5.73 - 6.21 6.08 - Pr -- 0.933 - 1.05 1.12 - Nd -- 5.16 - 5.93 6.42 - Sm -- 1.98 - 2.35 2.52 - Eu -- 0.84 - 0.788 0.990 - Gd -- 2.99 - 3.45 3.76 - Tb -- 0.576 - 0.658 0.725 - Dy -- 4.07 - 4.60 5.10 - Ho -- 0.914 - 1.02 1.13 - Er -- 2.67 - 2.92 3.28 - Tm -- 0.411 - 0.447 0.509 - Yb -- 2.79 - 2.98 3.45 - Lu -- 0.421 - 0.452 0.525 - Ta -- 0.091 - 0.094 0.090 - Pb 1.90 1.60 0.130 1.20 0.084 0.122 0.50 Th -- 0.256 - 0.283 0.225 - U -- 0.119 - 0.137 0.117 - Tables 90

Table 4.1: continued.

boninite (glass) CY-27.05.-02 Kapilio 2 Lat [N] 34°54.995 Long [E] 33°4.754 QAPF dacit boninite SiO2 70.7 52.1 TiO2 0.52 0.25 Al2O3 12.4 13.0 Fe2O3 6.90 - MnO 0.04 7.95 MgO 0.66 0.13 CaO 3.14 10.6 Na2O 4.26 0.55 K2O 0.16 0.08 P2O5 0.12 0.01 LOI 1.03 - Total 99.93 96.78 Li - - Sc - - V LOD - Cr 1.2 - Co - - Ni 11.0 - Cu 197 - Zn 8.7 - Ga - - Rb 0.70 - Sr 105 - Y 50.0 - Zr (XRF) 81.6 - Nb 2.10 - Sn - - Ba - - La - - Ce - - Pr - - Nd - - Sm - - Eu - - Gd - - Tb - - Dy - - Ho - - Er - - Tm - - Yb - - Lu - - Ta - - Pb 2.40 - Th - - U - - Tables 91

Table 4.1: continued.

XRF standard measurements, GeoZentrum Erlangen BIR-1 BIR-1 BR BR JA-2 JA-2 JA-3 JA-3 ref. (n=3) ref. (n=3) ref. (n=2) ref. (n=2) SiO2 47.8 47.1 38.2 38.3 56.2 56.8 62.3 61.6 TiO2 0.96 0.98 2.60 2.63 0.67 0.67 0.68 0.67 Al2O3 15.4 15.4 10.2 9.93 15.3 15.8 15.6 15.7 Fe2O3 11.3 11.4 12.9 12.8 6.14 6.29 6.59 6.64 MnO 0.17 0.17 0.20 0.20 0.11 0.11 0.11 0.10 MgO 9.68 9.47 13.3 13.2 7.68 7.74 3.65 4.01 CaO 13.2 13.2 13.8 13.7 6.48 6.38 6.28 6.46 Na2O 1.75 2.00 3.05 3.33 3.08 3.01 3.17 3.01 K2O 0.03 0.04 1.40 1.40 1.80 1.79 1.41 1.43 P2O5 0.05 0.01 1.04 1.05 0.150 0.149 0.11 0.11 LOI 0.12 0.12 3.00 3.00 1.06 1.06 0.12 0.12 Total 100.4 99.82 99.65 99.52 98.67 99.79 99.95 99.86 V 313 316 235 208 130 114 172 161 Cr 382 382 380 368 465 443 67.5 55.8 Ni 166 175 260 273 142 140 35.5 31.1 Cu 126 105 72 92.7 28.6 23.2 45.3 46.0 Zn 71.0 63.9 160 152 62.7 60 67.5 61.3 Ga 16.0 13.1 19.0 15.7 16.4 15.2 17.0 16.1 Rb 0.25 ¡ 0,5 47.0 52.2 68.0 72.7 36.0 35.9 Sr 108 106 1320 1353 252 242 294 267 Y 16.0 12.9 30.0 28.6 18.1 10.4 21.3 18.7 Zr 15.5 4.87 260 250 119 105 119 107 Nb 0.60 1.70 98.0 100 9.80 9.80 3.00 3.80 Pb 3.00 2.77 5.00 5.37 19.3 20.9 6.70 8.30 Tables 92

Table 4.1: continued.

ICP-MS standard measurements, Institut fur¨ Geowissenschaften, Kiel BIR-1 BIR-1 G-2 G-2 BHVO-2 BHVO-2 DR-N DR-N ref. (n=20) ref. (n=20 ref. (n=20) ref. (n=20) Li 3.40 3.24 34.0 29.6 4.60 4.52 40.0 36.9 Sc 44.0 40.5 3.5 3.5 31.8 29.5 28.0 26.3 V 313 307 36.0 35.4 317 315 220 225 Cr 382 245 8.70 6.79 289 196 - 30.3 Co 51.4 50.8 4.60 3.94 45.0 44.3 35.0 36.9 Ni 166 161 5.00 2.06 119 118 15.0 14.6 Cu 126 120 11 10.5 127 127 50.0 46.0 Zn 71.0 69.0 86.0 87.0 103 104 145 147 Ga 16.0 15.0 23.0 23.8 21.7 21.0 22.0 20.6 Rb 0.21 0.20 170 168 9.20 9.16 73.0 75.3 Sr 104 109 478 469 395 395 400 406 Y 16.0 15.3 11.0 9.46 25.5 24.6 26.0 25.3 Nb 0.55 0.53 12.0 11.9 18.0 16.9 7.00 6.89 Sn 0.65 0.70 1.80 1.66 1.80 1.75 2.00 1.84 Ba 5.83 5.87 1882 1800 130 130 385 385 La 0.62 0.50 89.0 88.7 15.2 15.3 21.5 21.6 Ce 1.95 1.83 160 164 38.0 37.8 46.0 46.6 Pr 0.38 0.37 18.0 16.7 5.30 5.30 5.70 5.63 Nd 2.50 2.39 55.0 54.0 25.0 24.4 23.5 23.2 Sm 1.10 1.10 7.20 7.27 6.20 6.04 5.4 5.1 Eu 0.54 0.52 1.40 1.34 2.06 2.05 1.45 1.41 Gd 1.85 1.80 4.30 4.76 6.30 6.13 4.70 5.00 Tb 0.36 0.36 0.48 0.51 0.93 0.93 0.77 0.78 Dy 2.50 2.57 2.40 2.15 5.25 5.28 4.60 4.68 Ho 0.57 0.57 0.40 0.35 0.99 0.97 1.00 0.94 Er 1.70 1.62 0.92 0.87 2.50 2.41 2.50 2.56 Tm 0.26 0.25 0.18 0.11 0.34 0.32 0.39 0.38 Yb 1.65 1.64 0.80 0.63 2.00 1.98 2.50 2.46 Lu 0.26 0.25 0.11 0.08 0.28 0.28 0.40 0.36 Ta 0.04 0.04 0.88 0.73 1.13 1.09 0.60 0.57 Pb 3.08 2.93 30.0 30.0 1.70 1.61 55.0 56.1 Th 0.03 0.03 24.7 24.6 1.21 1.19 5.00 4.74 U 0.01 0.01 2.07 3.21 0.41 0.42 1.50 1.49 Tables 93

Table 4.2: Parameters for modeling Troodos plagiogranitic melt evolution. Parental magma for MELTS calculation (oxides in wt.%).

Parental magma for MELTS calculation plus 7% Ol Akaki Olivine Boninitic canion glass (Laurent suggested parental (Regelous and H´ebert primary magma in prep.) 1989) parental melt (Kapilio 2) SiO2 49.6 39.2 48.9 52.9 TiO2 0.47 0.44 0.25 Al2O3 16.5 15.3 13.0 FeOT 6.74 7.15 Fe2O3 0.67 0.62 FeO 6.07 12.6 6.53 7.95 MnO 0.10 0.15 0.10 0.14 MgO 8.32 47.2 11.0 10.6 CaO 12.7 11.8 12.1 Na2O 1.13 1.05 0.55 K2O 0.09 0.08 0.08 P2O5 0.03 0.03 0.01 Cl 0.02 0.02 0.05 Mg# 65.1 87.0 75.1 70.4 Total 98.41 97.67

best fit T (°C) p (bar) 1700-500 500 Inc. 10°C

f O2 H2O QFM/QFM+1 0.5 wt.% Tables 94

Table 4.3: a) Least Squares Model: Step 1) basalt to basaltic andesite.

Recalculated analysis wt% to 100% OXIDE INITIAL FINAL Ol Cpx Pl SiO2 51.7 57.1 39.7 52.0 49.4 TiO2 0.790 1.59 0 0.53 0 Al2O3 16.4 14.6 0 2.54 31.4 FeOT 8.15 12.4 17.0 8.97 0.870 MnO 0.144 0.197 0.290 0.29 0 MgO 8.53 3.35 42.7 14.6 0.2 CaO 12.6 7.97 0.300 20.7 15.7 Na2O 1.44 2.39 0 0.310 2.47 K2O 0.092 0.280 0 0 0.070 P2O5 0.051 0.114 0 0 0 TOTAL 100 100 100 100 100

RESULTS BULK composition/ differences between magmas observed - add./ sub. calculated material observed calculated residuals OXIDE SiO2 48.8 5.35 5.42 -0.076 TiO2 0.173 0.796 0.926 -0.129 Al2O3 17.3 -1.81 -1.75 -0.06 FeOT 5.9 4.24 4.25 -0.015 MnO 0.137 0.053 0.039 0.014 MgO 11.2 -5.18 -5.14 -0.043 CaO 15.0 -4.65 -4.64 -0.013 Na2O 1.40 0.959 0.653 0.306 K2O 0.037 0.188 0.159 0.028 P2O5 0 0.063 0.075 -0.012 TOTAL 0

SUM of the squares of the residuals 0.1232

Amount as wt.% of substracted PHASE initial magma all phases added phases phases Ol -9.69 14.78 0 14.8 Cpx -21.4 32.71 0 32.7 Pl -34.4 52.51 0 52.5

TOTAL relative to initial magma 0 65.53 Tables 95

Table 4.3: a) Least Squares Model: Step 2) basaltic andesite to dacite

Recalculated analysis wt% to 100% OXIDE INITIAL FINAL Ol Cpx Pl2 Mt Il SiO2 57.1 75.0 39.7 52.0 51.1 0.050 0.030 TiO2 1.59 0.33 0 0.53 0 3.35 50.38 Al2O3 14.6 12.9 0 2.54 30.4 0.420 0 FeOT 12.4 2.51 17.0 8.97 0.740 96.0 47.8 MnO 0.197 0.041 0.290 0.290 0 0.110 1.75 MgO 3.35 0.659 42.7 14.6 0 0.010 0.040 CaO 7.97 4.62 0.3 20.7 14.2 0.020 0 Na2O 2.39 3.71 0 0.310 3.51 0 0 K2O 0.280 0.172 0 0 0.120 0 0 P2O5 0.114 0.122 0 0 0 0 0 TOTAL 100 100 100 100 100 100 100

RESULTS BULK differences between composition/ magmas observed - add./ sub. calculated material observed calculated residuals OXIDE SiO2 39.3 17.9 17.9 -0.03 TiO2 2.8 -1.25 -1.24 -0.011 Al2O3 16.3 -1.78 -1.74 -0.047 FeOT 22.1 -9.88 -9.87 -0.014 MnO 0.168 -0.156 -0.064 -0.092 MgO 6.01 -2.69 -2.69 0.001 CaO 11.3 -3.36 -3.37 0.014 Na2O 1.88 1.32 0.924 0.395 K2O 0.062 -0.107 0.055 -0.163 P2O5 0 0.008 0.061 -0.054 TOTAL 0

SUM of the squares of the residuals 0.1978

Amount as wt.% of PHASE initial all phases added substracted magma phases phases Ol -3.82 7.60 0 7.60 Cpx -9.51 18.9 0 18.9 Pl2 -26.1 51.9 0 51.9 Mt -8.77 17.4 0 17.4 Il -2.11 4.20 0 4.20 Tables 96

Table 4.3: b) Trace element modeling.

start composition basalt Ce Yb Sm Zr ET004-01-01 2.32 1.35 0.92 19.00

Stage 1: Least Squares Model: basalt to basaltic andesite phases Cpx Pl Ol Total Mineral content (%) 32.7 52.5 14.8 100.0 KD cpx KD plag KD ol D Ce 0.1254 0.0278 0.0076 0.057 Yb 0.601 0.0155 0.0468 0.212 Sm 0.4774 0.0132 0.0049 0.164 Zr 0.131 0.0094 0.0047 0.048 Reference 1 1 1

Stage 2: Least Squares Model: basaltic andesite to dacite/rhyolite phases Cpx Pl Ol Mt Ilm Total Mineral content (%) 18.9 51.9 7.60 17.4 4.2 100 KD cpx KD pl KD ol KD mt KD ilm D Ce 0.193 0.2214 0.0076 0.0019 0.0064 0.153 Yb 0.900 0.041 0.0468 0.28* 0.0469 0.246 Sm 0.75 0.1024 0.0049 0.007 0.0104 0.197 Zr 0.29 0.0126 0.0047 0.56 3.002 0.285 Reference 2 3 1 4 1 *KD for Lu

References: 1 (Fujikami et al. 1984), 2 (Klein et al. 2000), 3 (Aigner-Torres et al. 2007), 4 (Klemme et al. 2006)

Table 4.4: Melting of amphibolite.

Ce Sm Yb Zr References Amph, 900°C,1GPa 0.24 1.37 1.15 0.26 1 Cpx, 1050°C, 1.5GPa 0.193 0.75 0.9 0.29 2 Opx 0.0035 0.063 0.39 0.021 3 Plag 0.1 0.052 0.012 0.0009 3 Ilm 0.007 0.009 0.026 4 Mt 0 0.007 0.02 0.06 4,5

References: 1 (Klein et al. 1997), 2 (Klein et al. 2000), 3 (Dunn and Senn 1994), 4 (Nielsen et al. 1992), 5 (Horn et al. 1994) 5 Constraints on the evolution of the crust of the Semail ophiolite (Oman) from the geochemical composition of plagiogranites

Sarah Freund1, Karsten M. Haase1, Christoph Beier1, Jur¨ gen Koepke2, Martin Erdmann2 1GeoZentrum Nordbayern, Universit¨at Erlangen-Nurnb¨ erg, Schlossgarten 5, 91054 Erlangen, Germany 2Institute for Mineralogy, Leibnitz Universit¨at Hannover, Callinstr. 3, 30167 Hannover, Germany

5.1 Abstract

Plagiogranites from about thirty intrusions in seven blocks of the Oman ophiolite display well-defined major and trace element trends and Hf, Nd and Sr isotope ratios, suggesting genetical relations with associated mafic intrusives and with lavas by fractional crystal- lization although assimilation of hydrothermally altered rocks may have occurred. The plagiogranites exhibit two different compositional groups that can be correlated to the V1 and V2 lavas and thus two major magmatic phases exist within the whole crust of the Oman ophiolite. The first phase plagiogranites (P1) have higher TiO2, P2O5, Zr, Hf,

(Ce/Yb)N , εNd, εHf but lower Lu/Hf, Sm/Nd, Th/La and Sr isotope ratios compared to the second stage plagiogranites (P2). The second phase of the V2/P2 rocks indicates a stronger depletion of the mantle source and an increasing influence of a slab component compared to the first phase P1/V1 rocks including the sheeted dykes. Decoupling effects of mother/daughter element ratios from the isotope ratios of the plagiogranitic rocks and correlation of εHf and εNd towards sediment values suggest re-enrichement of the depleted mantle wedge by release of sediment melts during transition of the first toward the second magmatic stage. The geochemical correlation of P1 and P2 sample locations with well-constrained ages and the lavas imply that the magmatic evolution with the mantle depletion and increasing subduction influence occurred within a short time period between 96.5 and 95.5 Ma supporting the assumption that the Oman ophiolite could have formed in the initiation of a subduction zone. 5.2. Introduction 98

5.2 Introduction

Compared to other planets that are clearly dominated by mafic rocks the Earth shows abundant crustal portions consisting of felsic rocks. Felsic intrusives are rarely found in recent oceanic crust but commonly in ophiolites where they are generally called plagiogran- ites (Coleman and Peterman 1975; Malpas 1979). For example, in the Semail ophiolite of Oman, plagiogranite intrusions occur in all lower oceanic crustal regions and also in the mantle but are most abundant at the boundary between the sheeted dikes and gabbros (Lippard et al. 1986; Briqueu et al. 1991). They could reflect either extreme products of fractional crystallization in shallow melt lenses or they may form by re-melting of crustal rocks (Lippard et al. 1986; Amri et al. 1996; Gillis and Coogan 2002; Rollinson 2009). The reason for the formation of felsic melts of variable compositions and volumes in the predominantly mafic ophiolitic crust are still poorly understood but these processes are important for the chemical differentiation of the Earth. Additionally the plagiogranitic rocks yield insights into the formation of the oceanic crust in general. The Semail ophiolite in northern Oman and United Arab Emirates (UAE) is the largest, best-preserved and best-exposed ophiolite complex on Earth and contains abundant plagiogranite intrusions. It is therefore ideally suited to study the origin of plagiogranites in the oceanic crust. Plagiogranites in the Semail ophiolite are abundant in the upper oceanic crust, most commonly at the base of the sheeted dyke complex as dykes and lenses between 0.5 m up to >1 km length but also within the peridotites of the upper mantle and in the lower gabbro section (Lippard et al. 1986; Rollinson 2009; Briqueu et al. 1991). Age- dating of these rocks and the surrounding gabbros suggests that the mafic and felsic magmatism was contemporaneous (Rioux et al. 2013) but two different plagiogranite- formation models have so far been discussed: 1) partial melting of pre-existing mafic rocks and 2) extensive fractional crystallization of a mafic melt (Lippard et al. 1986; Koepke et al. 2004; Tsuchiya et al. 2013; Brophy 2008; Brophy and Pu 2012; Gillis and Coogan 2002; Amri et al. 1996). The mantle-hosted plagiogranites are most likely subduction respectively obduction process related (Briqueu et al. 1991; Cox et al. 1999).

Our aim here is to investigate possible scenarios for the formation of (1) the shallow- level intrusive rocks and specifically the plagiogranites. Systematic sampling of the felsic and associated mafic intrusive rocks in different crustal transects, settings and structural levels will help to decipher relations and differences of felsic and mafic intrusions and lavas within the Semail ophiolite. The focus of this study is the geochemical and isotopic characterization of the felsic lithologies from several settings and different crustal levels, comparing them with each other in order to understand the formation processes of these silica-rich melts and the generation of the Semail ophiolite. 5.3. Geological Background 99

5.3 Geological Background

5.3.1 Geographical and geological overview

The Semail ophiolite (Fig. 5.1) extends almost 600 km along the Gulf of Oman and covers >20,000 km2 of the Arabian Peninsula. It comprises twelve blocks, separated by faults. Each of these blocks extends over several tens of kilometres. The northernmost blocks (Khor Fakkan, Aswad and the northern tip of Fizh) are situated within the United Arab Emirates, the central and southern blocks lie on continental crust belonging to the Sultanate Oman. Each block comprises a complete section from the mantle (harzburgites, dunites), lower crust (wehrlites, pyroxenites, gabbros) and a well developed sheeted dyke complex up to the lava units (Nicolas et al. 2000) with intercalated pelagic sediments. The geodynamic environment during formation of the Oman ophiolite oceanic crust is still a matter of debate. Some authors emphasize the fast-spreading character and suggest a mid-ocean-ridge setting (Boudier et al. 1988; Nicolas et al. 1988; Coleman 1981), whereas others emphasize the geochemical indications of a subduction zone setting. Most authors agree that the Oman ophiolite formed during spreading in a supra-subduction setting (Pearce et al. 1981; Pearce et al. 1984; Searle and Cox 1999; Warren et al. 1005; Alabaster et al. 1982; Lippard et al. 1986; Tsuchiya et al. 2013; Arai et al. 2006; Warren et al. 2005) although the geochemical variations suggest geographically variable tectonic settings, e.g. there appears to be a north-south gradient within the Oman ophiolite between the MOR- related southern blocks and the subduction zone-related northern parts (VanTongeren et al. 2008; Hanghøj et al. 2010).

5.3.2 Geochemical overview

The Semail ophiolite lava sections are commonly divided into three consecutive stages: The V1 (or Geotimes) lavas with MORB-like chemistry are abundant and may represent an early spreading axis stage (Alabaster et al. 1982; Einaudi et al. 2000) whereas the later V2 (or Lasail, Alley, clinopyroxene-rich) lavas have an arc-like chemistry and occur mainly in the northern part of the ophiolite (Alabaster et al. 1982; Ishikawa et al. 2002; Ernewein et al. 1988). However, it was pointed out that although the V2 lavas clearly postdate the V1 lavas both lava types erupted within a very short time interval (Ernewein et al. 1988). In contrast, the alkaline to transitional within-plate basalts of the V3 (Salahi) lavas erupted considerably later (Alabaster et al. 1982; Lippard et al. 1986), and have been interpreted as being associated with the obduction process. These rocks have been observed only in the Salahi area of the Hilti massif (Godard et al. 2003). Comparable different magmatic stages are also observed within the mafic and the felsic plutonic section of the ophiolite (Rollinson 2009; Tsuchiya et al. 2013; Adachi and Miyashita 2003; Koepke et al. 2009; Yamasaki et al. 2006).

A study of plagiogranites in the Oman ophiolite by Amri et al (1996) suggested that 5.3. Geological Background 100

Figure 5.1: Maps of the Sumail ophiolite modified from (Godard et al. 2003; Nicolas et al. 2000). The ophiolite consist of 12 blocks, each massif comprises a complete section from the mantle, lower crust, sheeted dyke complex up to the lava units. The northernmost blocks (Khor Fakkan, Aswad and the northern tip of Fizh) lie within the United Arab Emirates, the others (most of the Fizh, Hilti, Sarami, Haylayn, Nakhl-Rustaq, Sumail, Tayin, Wuqbah, Miskin and Bahlah massif) belong to the sultanate Oman. a) Geographical overview of the Sumail ophiolite. b) Simplified geological map of the Sumail ophiolite with sample locations: red dots mark the plagiogranite sample locations, gray circle: P2 group, empty black circle: P1 group. 5.4. Samples and analytical methods 101 the crustal felsic rocks formed by fractional crystallization whereas the plagiogranite intrusions in the mantle represented partial melts from gabbroic sources. Rollinson (2009) described three different plagiogranite types from the Oman ophiolite; (1) plagiogranites that formed early by partial melting of gabbros with nearly flat Rare Earth Element (REE) pattern, (2) plagiogranites formed late by slab-fluid induced melting of depleted mantle and differentiation with a distinct depletion in Light REE (LREE) compared to the Middle REE (MREE) and Heavy REE (HREE). The third plagiogranite group comprises plagiogranites intruded into the mantle section which formed by mixing melts from the subducting slab and the depleted mantle wedge, emplaced along shear zones and associated with areas of deformation. The mantle plagiogranites have relatively enriched LREE and show an increasing depletion of middle toward the HREE compared to MORB (Rollinson 2009).

5.3.3 Ophiolite age

The oceanic crust of the Semail Ophiolite was formed during the Cretaceous. Several authors have measured different ages. Those have been calculated from zircon analyses. Rollinson (2009) suggested a difference of up to 3 to 4 Ma between early axis stage and the late stage plagiogranites. Tilton et al. (1981) has one sample located within the upper gabbros that has a U-Pb age of 97.3 Ma whereas Warren et al. (2005) determined Ophiolite ages of about 95 Ma respectively earlier than 96.4 (early crustal section). Goodenough et al. (2010) determine a range of ages between 96.4 and 95.2 (younger crustal section), whereas zircon U-Pb ages for “axis-stage” quartz diorites and “late-stage” tonalites from the Lasail complex give 99 2 to 100 2 Ma (Tsuchiya et al. 2013). Rioux et al. (2013) found zircon U-Pb ages of 96.44 0.06 to 95.48 0.06 Ma for axis stage gabbros and plagiogranites but 95.4 0.06 to 95.07 0.06 for the late stage gabbro and plagiogranite samples derived from five different blocks (Fizh, Haylayn, Rustaq, Sumail, Wadi Tayin block). The oceanic crust underwent thrusting followed by transport of up to 400-500 km in a southwestern direction over the Tethyan Hawasina basin before being emplaced onto the Arabian plate during the late Cretaceous (B´echennec et al. 1988).

5.4 Samples and analytical methods

5.4.1 Sampling

During field work in 2011 we sampled mainly felsic intrusive rocks from about thirty intrusions in the Fizh, Hilti, Sarami, Haylayn, Rustaq, Sumail and Wadi Tayin blocks of the Oman ophiolite (Figs. 5.1, 5.2). For simplicity we will call the samples “plagiogran- ites” in the following although they range from quartz-diorites, tonalites/trondhjemites to granodiorites (Fig. 5.3). We sampled seven of the localities that have also been age-dated using U-Pb geochronology by Rioux et al. (2013) and Warren et al. (2005) and thus we 5.4. Samples and analytical methods 102

Wadi Tayin block

P1 P2

Hilti block a) b) (Hi-25) Sumail block Sumail block (Su-8) (Su-23)

P2 P1

c) d) Rustaq block (RU-7) Haylayn block (HA-34)

P2

f) P1 P2

Sarami block (Sa-31)

e) g)

Figure 5.2: Field photographs of P1 and P2 sample locations in the Sumail Ophiolite. Importantly: both groups occasionally contain xenolites within felsic matrix (a, b, g). P1 comprises dikes and net-veined intrusions (a, e) but also large (> 100×100 m) plutons (c) similar to the majority of the P2 group intrusions (d, f, g). 5.4. Samples and analytical methods 103

Table 5.1: Age correlation of plagiogranite sample locations (this study) with U/Pb analyses of zircons from early respectively late stage plutonics studied by Warren et al. (2005) and Rioux et al. (2013).

Sumail ophiolite blocks Sample Age (Ma) (from north to south) Warren et Rioux et this study this study al. (2005) al. (2013) P1 P2 Sarami CW038 SA-33 95.5 Haylayn 8122M02 HA-34 95.5 Rustaq 8121M03 RU-7 96.2 Sumail CW018 SU-23 95.5 Sumail 9127M01 SU-18 95.3 Sumail OM01-05 SU-8 & 9 95.3 Wadi Tayin CW012 TA 13 & 14 95.3 can infer the ages of (some of) our samples (Table 5.1). The mineral content of the plagiogranites comprises plagioclase, quartz and mostly amphibole as primary major phases. Large, isolated plagiogranite bodies occur within the sheeted dike complex in the northern and central blocks of the Oman ophiolite (Fig. 5.1, Fizh, Hilti, Sarami, Haylayn and Nakhl-Rustaq blocks). The intrusive bodies are mostly cutting the upper sequences of the crust and are again intruded by aphyric dikes. We also sampled plagiogranites from the southern blocks (Sumail and Wadi Tayin massif) that occur as large bodies or as dikes intruding the gabbros or the sheeted dike- gabbro transition zone.

In order to detect any potential differences between the plagiogranites we pay atten- tion to several properties during sampling. During sampling we discriminated between plagiogranites intruded into the sheeted dike complex and the sheeted dike-root-zone (SDRZ), relatively small-scale intrusions or dikes (cm- to 10 m) and huge isolated bodies (10 m- to hundreds of m scale, Fig. 5.2) with the aim to be able to distinguish between different sources and generation processes like crustal partial melting and/or fractional crystallization.

5.4.2 Analytical methods

Fresh cores from samples were cut with a rock saw, washed in deionized H2O, crushed and pulverized in an agate mill. We determined the loss on ignition (LOI) by weighing the rock powder before and after drying: 1) 12 hours at 105 ◦C in a cabinet dryer and 2) 12 hours ◦ at 1030 C in a muffle furnace. The major element concentrations (SiO2, TiO2, Al2O3,

Fe2O3, MnO, MgO, CaO, Na2O, K2O, P2O5) and a few trace element concentrations (Cu, Ni, Zn, Cr, V, Rb, Sr, Zr, Nb, and Y) of whole rock powders were measured with a XRF (Spectro XEPOS plus) at the GeoZentrum Nordbayern. Averages of our measurements (and recommended values (Govindaraju and Roelandts 1993; Govindaraju 1994; Imai et al. 1995)) of rock standards (BIR-1, BR, JA-3) are given in Table 5.2. 5.4. Samples and analytical methods 104

Table 5.2: Standard measurements (oxides in wt%, elements in ppm).

standard measurements XRF (GeoZentrum Nordbayern) (n=8) (ref.) (n=8) (ref.) (n=8) (ref.) (n=8) (ref.) sample BE-N BE-N BR BR GA GA GH GH SiO2 38.5 38.2 38.6 38.2 70.0 69.9 75.9 75.8 TiO2 2.66 2.61 2.65 2.60 0.37 0.38 0.08 0.08 Al2O3 9.97 10.1 9.98 10.2 14.6 14.5 12.3 12.5 Fe2O3 13.0 12.8 12.9 12.9 2.69 2.83 1.31 1.34 MnO 0.20 0.20 0.20 0.20 0.09 0.09 0.05 0.05 MgO 13.1 13.2 13.2 13.3 0.93 0.95 0.01 0.03 CaO 14.0 13.9 13.7 13.8 2.44 2.45 0.70 0.69 Na2O 3.15 3.18 3.05 3.05 3.59 3.55 3.91 3.85 K2O 1.40 1.39 1.38 1.40 4.08 4.03 4.81 4.76 P2O5 1.06 1.05 1.05 1.04 0.12 0.12 0.01 0.01 LOI 2.44 2.45 2.81 3.00 0.88 1.00 0.77 0.70 Total 99.47 99.01 99.48 99.65 99.80 99.80 99.85 99.81 V 216 235 222 235 51.5 38.0 9.29 5.00 Cr 332 360 328 380 12.3 12.0 8.2 3 Ni 272 267 272 260 4.03 7.00 9.24 3.00 Zn 119 120 156 160 55.8 80 57.9 55 Ga 15.3 17.0 15.4 19.0 14.8 16.0 23.6 23.0 Rb 47.7 47.0 46.8 47.0 172 175 391 390 Sr 1416 1370 1376 1320 299 310 9.71 8.70 Y 30.1 30 29.7 30 18.8 21 78.4 75 Zr 265 260 262 260 142 150 154 150

The major element concentrations of minerals were measured on a JEOL JXA 8200 Superprobe electron microprobe at the GeoZentrum Nordbayern, Erlangen. The oxides T SiO2, TiO2, Al2O3, FeO , MnO, MgO, CaO, Na2O, K2O, Cr2O3 (and Cl) were measured. The EMP was operated at an accelerating voltage of 15 kV, a beam current of 15 nA and a focused beam diameter (0 µm) for the mineral analyses (except for feldspar where the beam diameter was set to 3 µm). Counting times were set to 20 s and 10 s for peaks and backgrounds for all elements. Trace elements were analysed using a Merchantek 266 LUV (266nm) Laser coupled with an Agilent 7500i (Inductively Plasma Mass Spectrometer: LA-ICP-MS) at the GeoZentrum Nordbayern following the procedure described in Schulz et al. (2006). Each glass disc (produced from sample powder for the XRF) was measured four times and the used values are averages. The external calibration was conducted by using NIST 610 (given values from Pearce et al. (1997)). Average values and recommended values of repeated analyses of international rock standards (NIST 612, BE-N, NIST 614, BIR1A, GH, JA-2, JA-3) are given in Table 5.2. Repeated analyses (n=4) of granitic rock standard GH give a standard deviation for precision 1σ <5.5% for all elements (except Rb <8%), accuracy <10% for all elements (except Rb, Zr, Ba, Tb, Ho, Th <15%, Sr, Dy, Er, Yb, Lu, Hf <19%, Nb <24%). Repeated analyses (n=16) of basaltic rock standard NIST 612 give a standard deviation for precision 1σ <4% for all elements, accuracy <5% for all elements (except Y, Gd, Er, Tm <10%, Nb, Ta <14.5%) and reproducibility of <5% for all elements (except Ta <11%) respectively. 5.4. Samples and analytical methods 105

Table 5.2: continued.

standard measurements LA-ICPMS (GeoZentrum Nordbayern) (n=4) (ref.) (n=4) (ref.) (n=4) (ref.) (n=4) (ref.) (n=16) (ref.) sample BIR1a BIR1a GH GH JA-2 JA-2 JA-3 JA-4 NIST NIST 612 Rb 0.24 0.25 350 390 92.1 68.0 43.2 36.0 30.3 31.6 Sr 112 108 8.52 10.0 237 252 278 294 72.6 76.2 Y 12.5 16.0 79.9 75.0 13.2 18.1 15.4 21.3 35.4 38.3 Zr 12.5 15.5 175 150 91.6 119 92.6 119 35.9 36.0 Nb 0.53 0.60 110 85.0 8.66 9.80 2.98 3.00 33.9 38.1 Ba 7.21 7.00 17.5 20.0 315 317 306 318 39.5 37.7 La 0.79 0.62 27.4 25.0 15.6 16.3 9.21 9.00 36.6 35.8 Ce 1.95 1.95 59.8 60.0 31.3 32.7 20.4 23.3 36.7 38.4 Pr 0.38 0.38 7.78 7.80 3.37 4.38 2.59 2.25 35.7 37.2 Nd 2.42 2.50 31.4 29.0 13.3 13.8 11.6 12.3 34.4 35.2 Sm 1.13 1.10 9.82 9.00 2.90 3.12 2.96 3.14 36.3 36.7 Eu 0.60 0.54 0.13 0.12 0.94 0.94 0.85 0.85 35.7 34.4 Gd 1.65 1.85 10.4 9.50 2.58 3.11 2.71 2.94 35.2 37.0 Tb 0.34 0.36 2.21 1.90 0.42 0.42 0.47 0.52 36.9 35.9 Dy 2.24 2.50 14.6 12.0 2.44 3.01 2.97 2.97 34.3 36.0 Ho 0.52 0.57 3.22 2.90 0.50 0.46 0.65 0.48 37.5 37.9 Er 1.41 1.70 9.47 8.00 1.47 1.37 1.74 1.46 35.4 37.4 Tm 0.22 0.26 1.39 1.30 0.21 0.30 0.26 0.3 35.1 37.6 Yb 1.67 1.65 9.66 8.00 1.64 1.67 1.97 2.18 38.1 40.0 Lu 0.23 0.26 1.35 1.10 0.23 0.27 0.27 0.32 36.1 37.7 Hf 0.52 0.60 7.83 6.60 2.45 2.89 2.89 3.43 35.4 34.8 Ta 0.04 0.04 4.83 4.80 0.58 0.61 0.23 0.14 34.8 39.8 Th 0.02 0.01 19.1 18.0 2.37 2.40 1.04 1.40 38.6 37.2

standard measurements TIMS (Kiel) n=8 n=8 n=16 Isotopes NBS 987 La Jolla SPEX 87 86 Sr/ SrN 0.710250 143 144 Nd/ NdN 0.511850 176Hf/177Hf 0.282170 5.5. Results 106

Due to more available values and therefore better discrimination between the crustal plagiogranites we use XRF measurements of Zr concentrations during following text. For comparing Zr concentrations measured by XRF and LA-ICPMS please see Table 5.1.

Strontium, Nd and Hf isotope ratios were analysed at GEOMAR, Helmholtz-Zentrum fu¨r Ozeanforschung, Kiel. Strontium and Nd isotope ratios were analysed in static mode on a TRITON TIMS and Hf isotopes on a MC ICPMS at IFM-GEOMAR, Kiel. Standard runs of NBS 987 (n=8) gave an average of 87Sr/86Sr of 0.710250 (2σ = 0.000005) (Sr isotope analyses were normalized to a common value of 0.710250 for NBS 987). La Jolla Nd standard measurements (n=8) yielded (normalized to La Jolla = 0.511850) a value of 0.511850 (2σ = 0.000007). The SPEX standard measurements (n=16) gave an average of 176Hf/177Hf of 0.282170 (2σ = 0.000004) respectively. ε values were calculated using the following CHUR (Chondritic Uniform Reservoir) parameters: present-day reference 143 144 147 144 176 177 values are Nd/ NdCHUR = 0.512630 and Sm/ NdCHUR = 0.196, Hf/ HfCHUR 176 177 87 86 = 0.282785 and Lu/ HfCHUR = 0.0336 (Bouvier et al. 2008), Sr/ SrCHUR = 0.7045 87 86 −1 and Rb/ SrCHUR = 0.0827 following Tsuchiya et al (2013) and λSm = 6.54×10–12 yr , λLu = 1.867×10–11 yr−1 and λRb = 1.42×10–11 yr−1 . The measured isotope ratios from plagiogranites and mafic wall rocks were age-corrected to 96 Ma (assumed ophiolite age).

5.5 Results

5.5.1 Petrology

Typical primary minerals observed in the crustally-hosted plagiogranite intrusions are euhedral plagioclase laths, quartz (sometimes granophyric), euhedral to anhedral amphi- bole (calcium-amphiboles, mainly magnesiohornblendes) and accessory apatite zircon, magnetite, ilmenite, titanite, ( clinopyroxene) in variable parageneses and contents. Most rocks show signs of hydrothermal alteration expressed by minor amounts of chlorite, epidote-clinozoisite and by variably amounts of plagioclase alteration.

5.5.2 Geochemistry

Major elements and incompatible elements

The crustally-hosted plagiogranites (n=87) range from (normative calculated) diorite, quartz-diorite, to tonalite/trondhjemite compositions and one granodiorite (Fig. 5.3).

The SiO2 contents of the plagiogranite samples range from 56 to 79 wt% at MgO contents ranging from 4 to 0.1 wt%. All sampled plagiogranites from all blocks and intrusion sizes show comparable trends in SiO2, Al2O3, Fe2O3 and K2O for a given MgO (Fig. 5.4, Table

5.3). However, based on the TiO2, Zr contents, REE ratios and Nd isotopes (Figs. 5.4b, 5.5, 5.6, Table 5.4) we distinguish two groups of the crustal plagiogranites: P1 and P2 where the former have higher Ti and Zr contents and (Ce/Yb)N ratios than the latter.

Compared to P1 rocks the P2 samples also display lower concentrations in P2O5 (0.04 -

5.5. Results 112 larger variability. The P1 group overlaps in incompatible element ratios (e.g. Zr/Hf vs.

(Ce/Yb)N and Lu/Hf vs. Sm/Nd) with the V1 lavas and the P2 group with the V2 lavas (Fig. 5.8).

Neodymium, Hf and Sr isotope ratios

Both crustally-intruded plagiogranite groups (n = 9) and the associated gabbros (n = 3) show high Hf and Nd isotope values resembling MORB but are significantly higher in Sr isotope ratios. Samples from the P1 group have slightly higher Nd and Hf isotope ratios compared to P2 (Fig. 5.9a, b). The P1 group has age-corrected εHf(96 Ma) = 17.1 to 24.2, whereas the P2 group has εHf(96 Ma) = 15.7 – 19.7. Slightly higher values in the P1 group 87 86 samples are also observed in age corrected εNd. The P1 group has lower Sr/ Sr(96 Ma) 87 86 = 0.70367 – 0.70477 compared to P2 ( Sr/ Sr(96 Ma) = 0.70441 – 0.70540).

5.6. Discussion 114

5.6 Discussion

5.6.1 Hydrothermal alteration of the plutonic rocks

Ophiolitic rocks are often significantly altered as a result of hydrothermal circulation of seawater through the upper crustal rocks resulting in modification of both mineralogical and geochemical composition. The occurrence of minerals such as chlorite, albite, sec- ondary quartz and also epidote - clinozoisite and actinolite in the plagiogranites is the result of greenschist to lower amphibolite facies metamorphism (Malpas et al. 1989) of the oceanic rocks. Hydrothermal alteration of the crust is accompanied by increasing water contents in the rocks expressed by correlation of fluid-mobile elements with LOI. On the other hand, some authors suggest partial melting of hydrothermally altered gabbros in the crust leading to felsic magmas or subduction-related fluid-induced mantle melting for the formation of the Oman late-stage crustal rocks (e.g. V2 lavas, wehrlites) (Alabaster et al. 1982; Ernewein et al. 1988; Einaudi et al. 2000; Godard et al. 2003; Koepke et al. 2009; Oeser et al. 2012; Lachize et al. 1996; Borisova et al. 2012; Negishi et al. 2013). These processes can also lead to a high water content in the plagiogranitic rocks. Generally higher LOI can be observed in the P2 group compared to the P1 group with relatively low LOI. Positive correlations of fluid-mobile elements like Rb and CaO with LOI (Fig. 5.7a, b) possibly suggest enrichment during seawater alteration. However, positive correlations of immobile elements (TiO2 and Fe2O3) in P2 (Fig. 5.7c, d) suggest that another processes must be responsible because these elements are not enriched in seawater. The higher 87Sr/86Sr isotope ratios (compared to MORB) could also reflect seawater alteration of the plagiogranites but the co-variation of Hf and Nd istotope ratios (Fig. 5.9b) especially in the P2 samples compared to the P1 group suggests that the Sr, Nd, and Hf isotopes reflect variation in the magma source of the plagiogranite groups rather than hydrothermal alteration. We conclude that hydrothermal alteration after the intrusion has only a minor effect on the rock compositions. Consequently, most major and trace elements as well as the radiogenic isotope ratios reflect magmatic processes. We suggest that the comparatively high water content reflected by the LOI of the P2 sample group indicates a subduction-signature similar to the V1 volcanic rocks rather than stronger hydrothermal alteration than the P1 group rocks.

5.6.2 The geochemical relationship between the plutonic and volcanic rocks

The volcanic evolution of the uppermost crust of the Oman ophiolite is reasonably well understood and it has been shown in several lava successions that the MORB-like V1 (or Geotimes lavas) are overlain by the V2 (or Lasail and Alley lavas) that have a clear subduction signature. The V1 lavas have been interpreted as extrusives at a spreading axis (Alabaster et al. 1982; Ernewein et al. 1988; Einaudi et al. 2000; Godard et al. 2003) whereas the younger, overlying V2 lavas are believed to have formed off-axis (Alabaster et al. 1982; Ernewein et al. 1988; Einaudi et al. 2000; Godard et al. 2003). In agreement

5.6. Discussion 116 with this model most rocks from the sheeted dike studied so far resemble the V1 stage rocks (Fig. 5.4) and thus most of the crust probably formed at a spreading axis. The geochemical differences between early and late-stage magmatism can be observed in terms of TiO2, Zr and (Ce/Yb)N versus MgO where the mafic V1 rocks are more enriched than the V2 rocks. Our new plagiogranite data show that the shallow level intrusives in the whole Oman Ophiolite fit well into these two groups and thus confirm the findings of V1-type and V2-type plutonic rocks in the Abyad (Nakhl-Rustaq block) and Lasail (Fizh block) complexes, respectively (MacLeod et al. 2013; Tsuchiya et al. 2013). The similarity of the shallow level plutonic rocks to the lavas also confirms that the diorites to tonalites indeed represent magma compositions rather than cumulates (Kelemen et al. 1997).

5.6.3 Formation of the felsic magmas

Indicators for partial melting and fractional crystallisation processes

Felsic rocks of both the P1 and P2 groups occur as elongated intrusions with diameters from several meters to kilometers and as dikes of only few cm to meter thickness. The large intrusions most likely represent melt lenses in the shallow crust that are believed to be common at mid-oceanic ridge settings (Detrick et al. 1987). Because we observe continuous trends of major and trace elements with MgO similar to the lavas as well as of incompatible element ratios and Nd isotope ratios with MgO (Figs. 5.4, 5.6) we suggest that most felsic magmas of the P1 and P2 groups form by extensive fractional crystallization from mafic parental melts. These trends closely resemble the fractional crystallization lines from the MELTS model suggested for the Oman lavas by MacLeod et al (2013) also supporting a formation of these trends by crystallization processes. Net- veining of felsic rocks containing basaltic and gabbroic xenoliths is observed frequently (Fig. 5.2, Sumail massif; Wadi Tayin massif) indicating that assimilation likely plays an important role during the plagiogranite genesis. Such assimilation-fractional crystalliza- tion (AFC) processes are typical for ascending magmas and have been described for ocean ridge settings (e.g. (Wanless et al. 2010; Freund et al. 2013)) and has been proposed for the Oman ophiolite plagiogranite on the basis of O isotope data (Stakes and Taylor 2003; Grimes et al. 2013). Borisova (2012) also assumes assimilation of altered rocks by a late-stage water-rich mafic melt to form chromite deposits of the Oman ophiolite and a “dramatically affected MORB magma”.

T Partial melting of mafic rocks leads to low TiO2 and FeO for melts with SiO2 > 60 wt% as inferred from experimental studies (Koepke et al. 2004; Koepke et al. 2007; France et al. 2010) and a negative correlation of the LREEs and SiO2 (> 62 wt%) inferred from model calculations and confirmed by natural samples (Brophy 2008, 2009; Brophy and

Pu 2012). Both Oman crustal plagiogranite groups show low TiO2 trends with increasing

SiO2 content (Fig. 5.10a). But the P1 group starts with a relatively high TiO2 content (>

1 wt%) at SiO2 contents of < 63 wt% and continues the decreasing trend of the V1 lavas.

5.6. Discussion 118

In contrast, the V2 lavas starts with low TiO2 and increases up to 1 wt% (at SiO2 = 57 wt%) and from up there the P2 group decreases in TiO2 with increasing SiO2 content.

The two groups show no negative correlation of LREEs and SiO2 (> 62 wt%) as suggested for plagiogranites produced by hydrous partial melting of MOR gabbros (Brophy 2008, 2009; Brophy and Pu 2012). The P1 group exhibits a strong increase in LREE with increasing SiO2 > 62 wt% (Fig. 5.10b, c) indicating an origin by fractional crystallisation of a MOR magma (Brophy 2008, 2009; Brophy and Pu 2012). The (Ce/Yb)N of the P1 and P2 samples are constant in a large range of MgO of 7 to 0 and 5 to 0 wt% (Fig. 5.6c), respectively. Partial melting processes would lead to increasing ratios of more incompatible elements to less incompatible elements and thus Ce would be significantly enriched relative to Yb. However, the negative Sr anomaly of the P1 group (Fig. 5.5a) in- dicates fractional crystallization of plagioclase from a mantle melt, whereas the positive Sr anomaly of the P2 group (Fig. 5.5b) suggest partial melting or assimilation of plagioclase- rich rocks. The negative Th anomaly of the P2 group displays fractionation of Th and U during plagiogranite formation, which is not typical for simple fractional crystallisation of a mantle melt. But, all except of three (Su 16, 17, 18) P2 intrusives comprise large plagiogranite bodies (> 100 × 100 m, Fig. 5.2), which are unlikely be produced by small degrees of partial melting from gabbros. We suggest fractional crystallization from a basaltic magma is most likely for generation of the P1 group, whereas assimilation fractional crystallization seems more plausible for generation of the P2 group in terms of incompatible elements. We suggest contamination of a mantle-derived melt with crustal rocks is possible.

Melting degrees and relations with crustal rocks

Tsuchiya et al. (2013) suggested that the felsic rocks in the Lasail complex (Fizh block) formed by fractional crystallization of the same magma which produced the massive (late- stage) gabbros including a significant amount of assimilation from the pre-existing layered gabbros. The rocks of the Lasail complex are comparable to our P2 group in terms of their major, trace element signatures and Sr and Nd isotope ratios (Figs. 5.4, 5.6, 5.8, 5.9). The P2 group is still less enriched in terms of most incompatible elements (compared to P1) and shows a “flatter” trace element pattern than the P1 group (Fig 5.5). This suggests a higher melting degree for magmas that formed the P2 group compared to the P1 magma, if both groups have been formed by the same source. Additionally, the V1 lavas overlap in most incompatible element ratios (e.g. Lu/Hf vs. Sm/Nd) with the P1 group and the V2 lavas overlap with the P2 group suggesting increasing melt degrees of the source from “axis-stage” (V1 & P1) toward “late-stage” (V2 & P2) (Fig. 5.8b). But the V2 lavas have slightly higher Lu/Hf ratios than the corresponding P2 group, possibly indicating assimilation distinctly amounts of “axis-stage” mafic plutonics (gabbros or dykes) during the plagiogranite differentiation whereas the corresponding V2 magma remains “original” during ascent and eruption. 5.6. Discussion 119

We suggest therefore that the P1 group is related to the axis-stage V1 lavas and mafic plutonics rather by differentiation of basaltic magma than by partial melting of gabbros and are probably formed by low melting degrees of a mantle with little subduction zone influence. The P2 group is possibly related to the “late-stage” magmatic event forming the late gabbros, wehrlites and V2 lavas but most likely contaminated by assimilation of “axis-stage” plutonics. Thus we suggest an (increasingly) subduction zone related source but different melting degrees for the two plagiogranite groups.

5.6.4 Implications for the tectonic setting during formation of the Oman ophiolite

Different models for the plate tectonic setting of formation of the Oman ophiolite have been suggested and recently MacLeod et al. (2013) suggested that none of the fractionation trends in the different lava suites of the ophiolite resembles dry MORB and rather, all of these lava trends require relatively high contents of water in the primary magmas. The highly precise age-dating of zircons (Warren et al. 2005; Rioux et al. 2013) implies that after the spreading stage the subduction influence of V2 stage commenced but surprisingly there is no transition between the two mantle sources. The increasing sudden influence of a slab component in magma genesis suggests a very fast (< 1 Myrs) and effective replacement of the mantle in the melting region of the Oman ophiolite as indicated by the change of Nd isotopes and (Ce/Yb)N (Fig. 5.6c, d). Some authors assumed the southern blocks as best analogue for a paleo-mid ocean ridge setting, not being subduction zone- related as it is suggested for the northern massifs (VanTongeren et al. 2008; Hanghøj et al. 2010) and additionally, the V2 lavas are more common in the northern blocks (Alabaster et al. 1982). In contrast, our study shows that the P1 group is more common in the southern blocks but the P2 group occurs in all sampled blocks except the Wadi Tayin block (Fig. 5.1). This indicates that the whole ophiolite formed in a subduction environment.

We assume that the P1 group represents magmas that crystallized in shallow melt lenses or intruded into the shallow crust at a spreading axis. Geochemical similarity of the P1 rocks to lavas from the Lau Basin suggests a situation similar to the recent Lau back-arc where variable distance of the spreading axis to the subduction zone results in a variable influx of the slab component in the magmas from the mantle wedge. The geochemical evolution from the V1/P1 to V2/P2 magmas implies that the Oman crust moved closer to a subduction zone from 97 to 95 Ma. Rioux et al (2013) interpreted the “axis-stage” as being a decompression stage above a juvenile subduction-zone due to slab-roll-back. Further development of the subduction zone (start of compression) increases the influence of a slab-component contaminating the depleted mantle wedge (Rioux et al. 2013) which is supported by experimental studies suggesting the importance of water in the formation of the late-stage magmas like the Oman ophiolite wehrlites (Koepke et al. 2009). We 5.6. Discussion 120 conclude that such a model of subduction initiation can explain the geochemical evolution towards more subduction-affected magmas from the V1/P1 to the V2/P2 rocks within a few million years.

5.6.5 Evidence of source depletion, decoupling effects and the nature of the slab component

Mid-ocean ridge basalts usually have variable Hf isotopes at almost constant Nd isotopes (Chauvel and Blichert-Toft 2001) whereas ocean island basalts and island arc basalts (as well as sediments) show well-correlated Hf and Nd isotopes. Although our P1 and P2 samples are within the MORB range (εHf +9 to +25 (Vervoort et al. 1999; Chauvel and Blichert-Toft 2001; Nowell et al. 1998)) the correlation towards lower Hf isotopes with decreasing Nd isotope ratios (and increasing Sr isotope ratios) is clearly visible (Fig. 5.9a, b). Comparable differences in terms of Nd isotope ratios have been observed between axis and late-stage mafic plutonics (Rioux et al. 2012; Rioux et al. 2013) (Fig. 5.11a) supporting the relation and the comparability of the lavas and the plutonic rocks. The trend of decreasing Hf and Nd isotopes and the correlation with significantly increasing Sr isotopes most likely suggests input of sedimentary material into the mantle wedge at the subduction zone where the Oman ophiolite formed. Such a sedimentary input is supported by the increasing Th/La ratios with decreasing εNd (Fig. 5.8c) because sedimentary material typically has high Th/La (GLOSS = 0.26 0.12) (Plank and Langmuir 1998). This sedimentary slab component most likely occurred as a melt because the fluid-immobile Hf must have been mobilized significantly and transported into the mantle wedge.

The plagiogranite Hf and Nd isotope ratios are negatively correlated with their mother/ daughter element ratios (Fig. 5.9d, e), suggesting a decoupling of the isotopes and the incompatible element composition (of the source). Decoupling of the isotope ratios from the parent/daughter element ratios suggest that the P1 samples have a more depleted time-averaged source than the P2 samples implying a relatively recent incompatible ele- ment depletion of the P2 source. Such a depletion of the P2 source is in agreement with the low Zr/Hf ratios and (Ce/Yb)N (Fig. 5.8a). The fact that a few elements like Ba, Th and Sr are increased (Fig. 5.5) but not all of the incompatible elements indicate mixing of the highly depleted mantle wedge with very little sedimentary melt. While the sources of the plagiogranites were probably (isotopically) enriched during emplacement, their trace element ratios rather reflect varying degrees of partial melting of their source. The higher Nd and Hf isotope ratios of the P1 group suggest a more depleted source, while the lower trace element ratios probably display slightly lower degrees of mantle melting compared to the P2 group. The P2 group shows lower isotope ratios suggesting that their source was slightly more enriched and maybe the trace element ratios suggest slightly higher degrees of source melting. Combined this decoupling is consistent with the melting behavior of 5.6. Discussion 121 heterogeneous mantle sources (see Ito and Mahoney (2005) and references therein).

The P1 and P2 rocks show an increasing influence of a subduction component with time and the Nd and Hf isotope ratios indicate that both isotope systems were affected by this component (e.g. enrichment of the source). Partial melt of gabbros or sheeted dykes would not produce silica rich melt with lower Hf and Nd isotope ratios but rather a component with significantly different Hf and Nd isotope composition like sediments is required. In order to change the Hf isotope composition Hf must be mobilized significantly from the slab and this most likely reflects an input of sedimentary melt from the subducting slab (Hanyu et al. 2002; Tollstrup et al. 2010; Woodhead et al. 2001). Furthermore, the P2 group has slightly higher Sr isotope ratios compared to the P1 group, which supports a stronger influence of the slab component. We conclude, that the P2 group (respectively all late stage magmatics) is increasingly influenced by a subduction component enriched mantle source, most likely small amounts of sedimentary material.

5.6.6 The influence of slab derived fluids on magma differentiation

The different magma groups in the Oman ophiolite show clearly distinct fractionation trends that are apparently related to the mantle source compositions. Whereas the V1/P1 rocks show strong Fe enrichment typical of a relatively dry tholeiitic fractionation trend the V2/P2 group of rocks displays much less Fe and also Ti enrichment (Fig. 5.4b, d). According to the MELTS model curves of Macleod et al. (2013) the V2/P2 trend reflects higher water contents in agreement with the relatively strong subduction input of these rocks. Crystallization under higher water pressure (<500 bar) and oxygen fugacity (NNO+2) compared to MORB is also assumed by Lachize et al. (1996) for a suite of calc- alkaline gabbros and norites intruded into still viscous earlier gabbros in the Haylayn block (comparable to similar plutonic relations e.g. in the Rustaq block (Browning 1984)). This calc-alkaline plutonics contain layer rich of Ti-magnetite cumulates, upwards replaced by magmatic Ti-rich amphiboles (Lachize et al. 1996). The oxygen fugacity has a distinct effect on the crystallization assemblage and the fractionation trend of a magma. A high f O2 triggers the precipitation of magnetite and stabilizes orthopyroxene (Grove and Baker 1984; Juster et al. 1989), both is observed in the late plutonics. This would cause increase of SiO2 and decrease of FeO (and TiO2) in the residual melt. Higher water contents during late stage magmatism are also assumed from more recent studies and various authors with different research focuses (e.g. (Oeser et al. 2012)). We suggest that increasing water contents and oxygen fugacities in the mantle wedge due to the increasing influence of a sedimentary melt component from the subducting slab promote earlier FeTi-oxide precipitation in the late-stage magmas, which most likely account for the lower Fe and Ti content in the V2/P2 group (with < 5 wt% MgO) compared to the V1/V2 group.

5.8. Acknowledgements 123 depositions (jasper and sediments) (Fleet and Robertson 1980; Ernewein et al. 1988) suggest a time interval and an increase in hydrothermal activity between the V1 and V2 lava eruptions. Otherwise, the wehrlites of the Haylayn block suggest only short time intervals between early- and late stage plutonism because they intrude discordantly into still very hot, layered on-axis gabbros and deformed them viscously (Koepke et al. 2009). Thus we suggest our plagiogranites groups P1 and P2 indicate that the magmatic evolution with the mantle depletion and increasing subduction influence occurred within a relatively short time period between 96.5 and 95.3 Ma. Interestingly, such a short period of time from spreading to subduction is predicted by models of the initiation of subduction zones (Gerya et al. 2008) which supports the idea that the Oman ophiolite could have formed in such an environment.

5.7 Conclusions

Well-defined major and trace element trends and Hf, Nd and Sr isotope ratios suggest that the shallow plagiogranite intrusions and the associated isotropic gabbros and lavas are genetically related. The shallow-level plagiogranites exhibit two different compositional groups that can be correlated to the V1 and V2 lavas and thus two major magmatic phases exist within the crust of the Oman Ophiolite. The first phase produces the volcanic and intrusive rocks (V1/P1) including the sheeted dyke complex and thus formed by decompression melting in a spreading environment with only little geochemical influence of a subduction zone. The second phase of the V2/P2 rocks indicates a stronger depletion of the mantle source and an increasing influence of a slab component. Importantly, both the P1 and the P2 groups occur in the whole Ophiolite suggesting that the entire Ophiolite formed above a subducting slab. The release of sediment melts re-enriches the depleted mantle wedge causing fluid-induced melting. The close geochemical and isotopic similarities of the P1 and P2 plagiogranites to the V1 and V2 lavas and specifically the mafic rocks, the relatively large volumes of evolved plutonic rocks, and the continuous major element trends of both groups suggest that the plagiogranites formed largely by fractional crystallization of basaltic magmas although assimilation of hydrothermally altered rocks may have occurred. The geochemical correlation of the plutonic rocks with well-constrained ages and the lavas imply that the magmatic evolution with the mantle depletion and increasing subduction influence occurred within a short time period between 96.5 and 95.5 Ma supporting the assumption that the Oman ophiolite could have formed in the initiation of a subduction zone.

5.8 Acknowledgements

We acknowledge the help of H. Br¨atz and R. Klemd (both GeoZentrum Nordbayern) during trace element analyses. We thank A. Richter for patient support with electron 5.8. Acknowledgements 124 microprobe analyses. C. Weinzierl is thanked for his generous helpful corrective advice. We thank the Director General of Minerals, Ministry of Commerce and Industry of the Sultanate of Oman, for allowing us to conduct field work in the Sultanate of Oman. S. Freund thanks L. Pflug for several inspiring discussions and critical questions. S. Freund was funded by DFG grant HA2568/21. Tables 125

Table 5.3: Major elements of the Oman plagiogranites (oxides in wt.%).

P1 sample OM11-RU OM11-RU OM11-TA OM11-TA OM11-TA OM11-SU OM11-Fi OM11-SA no. 7C 7E 15A 15B 20B 23E 28B 32A Lat [N] 23◦29.363 23◦29.422 22◦51.727 22◦51.723 22◦48.990 23◦29.742 24◦37.290 23◦58.443 Long [E] 57◦43.057 57◦43.032 58◦37.688 58◦37.701 58◦20.839 58◦14.450 56◦20.633 56◦42.767 SiO2 57.5 67.6 66.8 74.7 70.9 70.1 63.0 74.6 TiO2 1.09 0.59 0.59 0.24 0.39 0.51 1.20 0.29 Al2O3 14.9 14.2 14.7 12.2 13.8 13.9 14.4 12.2 Fe2O3 9.05 6.10 6.51 3.81 3.93 4.78 8.15 3.94 MnO 0.09 0.03 0.05 0.02 0.03 0.05 0.07 0.02 MgO 2.71 1.03 0.99 0.61 0.71 0.77 1.75 0.20 CaO 6.70 2.09 2.68 0.94 1.70 1.79 3.71 2.17 Na2O 5.68 7.17 6.32 6.34 7.38 7.27 6.31 6.02 K2O 0.23 0.13 0.28 0.09 0.07 0.25 0.14 0.03 P2O5 0.27 0.13 0.14 0.02 0.07 0.11 0.23 0.04 LOI 1.65 0.87 0.82 0.82 0.99 0.30 0.93 0.40 Total 99.85 99.90 99.90 99.88 99.01 99.91 99.91 99.92

P1 sample OM11-SA OM11-TA OM11-TA OM11-TA OM11-TA OM11-TA OM11-TA OM11-TA no. 32B 11A 11G 19B 13A 13B 13Fd 13Fh Lat [N] 23◦58.443 22◦46.955 22◦46.906 22◦51.949 22◦49.120 22◦49.120 22◦49.120 22◦49.120 Long [E] 56◦42.767 58◦26.753 58◦26.801 58◦20.679 58◦31.650 58◦31.650 58◦31.650 58◦31.650 SiO2 75.4 58.4 56.4 61.6 60.9 60.6 58.7 66.3 TiO2 0.24 1.45 1.37 1.15 1.12 1.15 1.53 0.84 Al2O3 12.2 16.6 15.3 15.1 16.2 16.4 15.2 15.4 Fe2O3 3.61 7.46 10.5 7.70 7.11 7.21 8.77 4.88 MnO 0.01 0.08 0.09 0.06 0.06 0.06 0.08 0.04 MgO 0.16 2.66 3.70 2.20 2.02 2.03 2.99 1.36 CaO 1.55 4.89 5.17 4.23 4.76 4.37 4.85 3.57 Na2O 6.17 6.74 5.49 6.03 6.42 6.54 6.49 6.47 K2O 0.14 0.33 0.28 0.45 0.22 0.25 0.25 0.24 P2O5 0.02 0.24 0.13 0.19 0.36 0.35 0.15 0.21 LOI 0.39 1.06 1.46 1.22 0.74 1.02 0.93 0.62 Total 99.91 99.91 99.88 98.78 99.91 99.92 99.90 99.87

P1 sample OM11-HA OM11-HA OM11-HA OM11-RU OM11-TA OM11-TA OM11-TA OM11-TA no. 4A 4B 4C 7F 12D 14A 14B 14C Lat [N] 23◦36.562 23◦36.562 23◦36.562 23◦29.422 22◦49.649 22◦48.593 22◦48.593 22◦48.625 Long [E] 57◦15.375 57◦15.375 57◦15.375 57◦43.032 58◦31.357 58◦28.894 58◦28.894 58◦28.873 SiO2 75.9 58.8 67.9 70.5 76.6 69.0 75.5 74.5 TiO2 0.26 0.44 0.67 0.49 0.28 0.71 0.37 0.70 Al2O3 12.1 18.0 13.8 13.1 11.8 13.8 13.0 14.2 Fe2O3 1.35 1.58 5.79 5.10 2.52 4.64 1.43 0.93 MnO 0.02 0.02 0.12 0.02 0.02 0.04 0.02 0.02 MgO 0.46 0.67 1.16 1.00 0.45 1.78 0.90 0.25 CaO 2.88 6.41 2.69 2.10 1.44 1.65 0.91 1.37 Na2O 5.51 9.21 6.29 6.57 6.16 6.68 6.70 6.83 K2O 0.21 0.07 0.36 0.02 0.16 0.15 0.14 0.26 P2O5 0.05 0.10 0.20 0.11 0.04 0.19 0.00 0.00 LOI 1.54 4.53 0.96 0.90 0.39 1.34 0.98 0.86 Total 99.92 99.90 99.92 99.90 99.89 99.91 99.91 99.91 Tables 126

Table 5.3: continued.

P1 sample OM11-TA OM11-TA OM11-TA OM11-TA OM11-TA OM11-TA OM11-TA OM11-TA no. 14D 15C 11C 11H 11i 11J 19A 20A Lat [N] 22◦48.625 22◦51.754 22◦46.911 22◦46.883 22◦46.883 22◦46.842 22◦51.949 22◦48.990 Long [E] 58◦28.873 58◦37.704 58◦26.792 58◦26.804 58◦26.804 58◦26.820 58◦20.679 58◦20.839 SiO2 69.8 70.2 69.4 73.8 62.3 67.9 60.7 66.7 TiO2 0.58 0.39 0.53 0.49 1.01 0.60 1.28 0.70 Al2O3 13.9 13.0 13.9 12.7 14.5 14.1 14.5 13.9 Fe2O3 3.55 6.02 4.34 2.77 6.92 5.89 8.96 6.55 MnO 0.04 0.05 0.02 0.02 0.06 0.02 0.08 0.09 MgO 1.34 0.70 0.82 0.60 2.35 1.15 2.41 1.39 CaO 2.43 2.41 3.59 2.69 4.72 2.15 4.69 2.34 Na2O 6.71 6.05 6.26 5.48 5.87 6.68 5.20 6.84 K2O 0.24 0.13 0.13 0.24 0.28 0.23 0.38 0.27 P2O5 0.03 0.07 0.16 0.09 0.12 0.14 0.17 0.23 LOI 1.27 0.90 0.79 1.04 1.80 1.04 1.51 0.89 Total 99.90 99.86 99.91 99.89 99.86 99.91 99.89 99.88

P1 sample OM11-SU OM11-SU OM11-SU OM11-SU OM11-SU OM11-SU OM11-Hi OM11-Hi no. 23A 23B 23C 23F 23G 23H 26A 26B Lat [N] 23◦29.742 23◦29.742 23◦29.742 23◦29.719 23◦29.794 23◦29.794 24◦16.139 24◦16.139 Long [E] 58◦14.450 58◦14.450 58◦14.450 58◦14.469 58◦14.442 58◦14.442 56◦21.613 56◦21.613 SiO2 75.4 78.1 76.2 77.9 79.2 67.5 71.4 67.8 TiO2 0.25 0.21 0.34 0.19 0.19 0.75 0.48 0.62 Al2O3 12.8 12.4 13.9 11.4 11.9 14.0 13.0 13.4 Fe2O3 1.69 0.64 0.35 1.45 0.93 5.39 5.29 5.86 MnO 0.02 0.01 0.00 0.01 0.01 0.11 0.07 0.08 MgO 0.52 0.30 0.07 0.25 0.12 1.48 1.04 1.58 CaO 1.24 0.78 0.96 2.56 0.96 3.01 1.24 1.63 Na2O 7.25 6.80 7.67 5.39 5.93 6.01 5.84 6.56 K2O 0.05 0.07 0.06 0.06 0.19 0.40 0.25 0.21 P2O5 0.03 0.02 0.02 0.02 0.29 0.11 0.14 LOI 0.68 0.54 0.31 0.73 0.49 0.94 1.15 2.01 Total 99.91 99.91 99.90 99.89 99.90 99.89 99.92 99.92

P1 P2 sample OM11-Hi OM11-Fi OM11-SU OM11-SU OM11-Hi OM11-Hi OM11-Fi OM11-Fi no. 26D 28C 8C 9D 24A 25A 27A 27C Lat [N] 24◦16.139 24◦41.675 23◦04.810 23◦05.043 24◦08.576 24◦14.515 24◦41.675 24◦41.748 Long [E] 56◦21.613 56◦20.753 58◦07.943 58◦07.340 56◦29.969 56◦26.468 56◦20.753 56◦20.850 SiO2 68.6 61.1 63.2 73.5 68.0 65.8 70.7 70.8 TiO2 0.52 1.11 0.56 0.28 0.72 0.74 0.37 0.37 Al2O3 13.1 14.4 14.9 13.5 13.1 13.1 12.2 11.9 Fe2O3 5.84 7.84 5.94 2.56 7.18 6.35 5.58 6.01 MnO 0.07 0.09 0.05 0.01 0.03 0.03 0.07 0.08 MgO 1.21 1.72 2.46 0.62 1.32 1.98 0.54 0.80 CaO 1.72 4.33 5.30 3.27 1.85 3.67 3.18 3.12 Na2O 6.70 7.95 5.58 5.23 5.94 4.48 4.95 4.81 K2O 0.22 0.14 0.16 0.20 0.27 0.54 0.27 0.33 P2O5 0.11 0.18 0.05 0.06 0.10 0.07 0.06 0.06 LOI 1.84 1.03 1.72 0.73 1.41 3.17 2.09 1.64 Total 99.92 99.89 99.89 99.94 99.92 99.94 99.93 99.93 Tables 127

Table 5.3: continued.

P2 sample OM11-Fi OM11-Fi OM11-Hi OM11-Hi OM11-Hi OM11-SA OM11-SA OM11-SA no. 29B 29C 30A 30B 30C 33A 33C 33F Lat [N] 24◦21.725 24◦21.725 24◦03.254 24◦03.254 24◦03.254 23◦55.414 23◦55.414 23◦55.414 Long [E] 56◦24.433 56◦24.433 56◦32.442 56◦32.442 56◦32.442 56◦48.498 56◦48.498 56◦48.498 SiO2 66.3 67.6 58.8 57.0 66.3 70.9 68.1 70.7 TiO2 0.72 0.72 0.94 0.92 0.86 0.47 0.59 0.50 Al2O3 12.6 12.9 14.1 14.3 13.1 12.3 12.6 12.3 Fe2O3 8.76 6.30 9.65 8.59 6.04 3.73 7.13 6.10 MnO 0.02 0.02 0.13 0.14 0.08 0.04 0.06 0.03 MgO 1.42 1.36 3.87 3.60 1.52 0.90 1.22 0.75 CaO 3.81 3.68 4.53 7.05 4.40 5.45 3.62 3.01 Na2O 4.58 4.64 3.98 3.44 4.04 3.72 4.71 4.72 K2O 0.26 0.30 0.80 0.48 0.38 0.09 0.29 0.28 P2O5 0.07 0.07 0.09 0.10 0.19 0.09 0.07 0.10 LOI 1.35 2.33 2.92 4.26 3.01 2.28 1.55 1.40 Total 99.94 99.93 99.83 99.87 99.90 99.94 99.92 99.93

P2 sample OM11-HA OM11-SU OM11-SA OM11-SA OM11-HA OM11-HA OM11-HA OM11-SU no. 34C 16A 31A 31E 34D 34F 34G 17D Lat [N] 23◦32.067 23◦08.384 23◦58.489 23◦58.489 23◦32.067 23◦32.067 23◦32.067 23◦08.356 Long [E] 57◦26.134 58◦05.628 56◦44.729 56◦44.729 57◦26.134 57◦26.134 57◦26.134 58◦05.566 SiO2 72.8 72.8 74.7 74.9 73.6 73.8 74.3 66.3 TiO2 0.52 0.38 0.26 0.26 0.54 0.49 0.56 0.34 Al2O3 12.2 13.2 11.9 11.9 12.3 12.0 12.7 14.3 Fe2O3 4.44 3.30 4.24 4.29 2.69 3.44 1.18 3.82 MnO 0.02 0.04 0.06 0.07 0.01 0.04 0.02 0.09 MgO 0.87 0.60 0.28 0.31 0.83 1.57 1.18 4.07 CaO 3.17 2.76 1.85 1.52 3.71 1.84 3.51 6.13 Na2O 4.04 5.10 5.73 5.70 5.21 4.88 5.48 3.71 K2O 0.31 0.51 0.31 0.25 0.12 0.07 0.21 0.09 P2O5 0.09 0.08 0.04 0.04 0.12 0.09 0.15 0.05 LOI 1.48 1.24 0.55 0.66 0.75 1.73 0.73 1.08 Total 99.94 99.94 99.94 99.94 99.93 99.93 99.94 99.90

P2 sample OM11-SU OM11-SU OM11-SU OM11-HA OM11-HA OM11-HA OM11-HA OM11-SU no. 18A 18B 18C 34A 34B 34E 34H 8D Lat [N] 23◦08.333 23◦08.354 23◦08.344 23◦32.067 23◦32.067 23◦32.067 23◦32.067 23◦04.810 Long [E] 58◦07.774 58◦07.772 58◦07.761 57◦26.134 57◦26.134 57◦26.134 57◦26.134 58◦07.943 SiO2 68.9 74.8 69.7 72.8 72.7 74.8 73.2 59.5 TiO2 0.31 0.22 0.33 0.51 0.57 0.52 0.55 0.49 Al2O3 15.7 12.8 15.1 12.1 12.7 12.5 12.0 15.3 Fe2O3 1.11 2.12 3.15 4.65 3.32 1.10 3.64 7.14 MnO 0.02 0.03 0.04 0.03 0.02 0.02 0.02 0.10 MgO 1.12 0.61 1.13 0.87 0.96 1.08 0.99 3.47 CaO 6.60 3.83 4.65 2.53 3.59 3.74 3.39 6.04 Na2O 4.40 4.11 4.13 4.05 5.10 5.43 4.80 5.70 K2O 0.06 0.08 0.10 0.31 0.16 0.13 0.17 0.26 P2O5 0.08 0.05 0.10 0.09 0.11 0.06 0.10 0.04 LOI 1.60 1.26 1.54 1.97 0.72 0.63 1.09 1.81 Total 99.94 99.94 99.94 99.91 99.93 99.93 99.92 99.88 Tables 128

Table 5.3: continued.

P2 sample OM11-SU OM11-SU OM11-SU OM11-SU OM11-SU OM11-Hi OM11-Fi OM11-Fi no. 8F 8G 8H 9B 22 24C 27B 29A Lat [N] 23◦04.810 23◦04.810 23◦04.810 23◦04.991 23◦08.413 24◦08.576 24◦41.675 24◦21.725 Long [E] 58◦07.943 58◦07.943 58◦07.943 58◦07.360 58◦06.224 56◦29.969 56◦20.753 56◦24.433 SiO2 64.2 63.4 71.7 73.0 78.2 69.5 71.1 65.8 TiO2 0.58 0.50 0.35 0.30 0.11 0.69 0.37 0.56 Al2O3 15.1 15.2 14.0 13.5 12.0 12.8 12.4 11.9 Fe2O3 5.85 6.51 3.26 0.88 0.63 6.31 5.83 8.47 MnO 0.07 0.09 0.02 0.01 0.02 0.04 0.06 0.04 MgO 1.93 2.39 0.94 0.76 0.09 1.04 0.45 1.29 CaO 5.57 5.27 3.15 4.82 0.74 2.58 3.21 5.55 Na2O 4.72 4.98 5.36 5.20 4.45 5.46 5.00 3.68 K2O 0.22 0.23 0.22 0.15 3.46 0.29 0.34 0.35 P2O5 0.05 0.05 0.06 0.07 0.04 0.12 0.06 0.06 LOI 1.61 1.22 0.85 1.27 0.18 1.14 1.17 2.14 Total 99.89 99.88 99.93 99.93 99.94 99.94 99.93 99.89

P2 wall-rock sample OM11-Fi OM11-SA OM11-SA OM11-SA OM11-SA OM11-RU OM11-RU OM11-RU no. 29D 31B 31D 33D 33G 7D 7A 7B Lat [N] 24◦21.725 23◦58.489 23◦58.489 23◦55.414 23◦55.414 23◦29.359 23◦29.363 23◦29.363 Long [E] 56◦24.433 56◦44.729 56◦44.729 56◦48.498 56◦48.498 57◦42.891 57◦43.057 57◦43.057 SiO2 73.9 75.7 75.3 68.1 69.7 49.6 50.7 52.1 TiO2 0.39 0.25 0.25 0.57 0.58 1.71 1.81 0.74 Al2O3 11.5 11.9 12.0 12.9 12.6 17.6 15.4 15.2 Fe2O3 4.57 3.72 3.83 7.20 6.05 11.5 13.2 9.08 MnO 0.01 0.06 0.05 0.03 0.05 0.18 0.19 0.16 MgO 0.80 0.29 0.26 1.25 1.18 3.92 4.29 7.05 CaO 2.58 1.44 1.95 4.06 3.25 6.65 6.62 9.40 Na2O 4.46 5.64 5.20 4.10 4.29 5.29 4.86 4.09 K2O 0.17 0.32 0.35 0.27 0.26 0.17 0.27 0.23 P2O5 0.06 0.05 0.04 0.14 0.07 0.11 0.16 0.08 LOI 1.53 0.58 0.69 1.31 1.82 3.20 2.36 1.75 Total 99.93 99.93 99.92 99.93 99.91 99.85 99.85 99.87

wall-rock sample OM11-RU OM11-RU OM11-SU OM11-SU OM11-SU OM11-SU OM11-SU OM11-TA no. 7G 7H 8A 8B 8E 9A 9C 11F Lat [N] 23◦29.422 23◦29.307 23◦04.810 23◦04.810 23◦04.810 23◦04.991 23◦05.043 22◦46.905 Long [E] 57◦43.032 57◦43.222 58◦07.943 58◦07.943 58◦07.943 58◦07.360 58◦07.340 58◦26.798 SiO2 47.7 48.1 71.5 72.6 51.4 73.8 63.7 52.6 TiO2 0.58 0.66 0.37 0.38 0.25 0.27 0.46 1.25 Al2O3 15.9 16.0 13.7 13.6 9.0 13.4 15.1 15.4 Fe2O3 7.97 8.67 3.27 2.86 6.85 2.40 7.12 9.80 MnO 0.13 0.17 0.02 0.02 0.15 0.01 0.14 0.08 MgO 9.94 9.60 0.84 0.76 14.1 0.67 2.37 5.69 CaO 12.4 12.4 4.00 3.68 16.2 3.17 4.81 6.24 Na2O 2.19 2.02 4.78 4.87 0.77 5.06 4.72 5.38 K2O 0.18 0.13 0.14 0.14 0.05 0.18 0.33 0.21 P2O5 0.03 0.03 0.06 0.06 0.03 0.06 0.04 0.11 LOI 2.68 2.07 1.18 0.99 0.94 0.89 1.17 3.14 Total 99.77 99.83 99.93 99.94 99.74 99.93 99.88 99.88 Tables 129

Table 5.3: continued.

wall-rock sample OM11-TA OM11-TA OM11-TA OM11-TA OM11-Hi OM11-Hi OM11-Fi OM11-Fi OM11-SA no. 12A 12B 13C 21B 25B 26C 28A 29E 33E Lat [N] 22◦49.649 22◦49.649 22◦49.120 23◦06.852 24◦14.515 24◦16.139 24◦37.290 24◦21.725 23◦55.414 Long [E] 58◦31.357 58◦31.357 58◦31.650 58◦15.499 56◦26.468 56◦21.613 56◦20.633 56◦24.433 56◦48.498 SiO2 48.3 51.5 52.6 53.5 59.3 50.3 50.1 55.2 53.3 TiO2 2.15 1.95 1.31 1.74 0.53 0.16 1.21 1.04 0.54 Al2O3 15.6 15.7 15.7 15.1 14.8 13.7 14.9 14.7 15.7 Fe2O3 13.0 12.7 9.79 12.4 6.70 4.03 11.5 10.6 8.76 MnO 0.12 0.13 0.13 0.16 0.08 0.10 0.15 0.04 0.15 MgO 4.70 4.66 5.87 4.03 5.17 11.4 6.46 5.05 7.16 CaO 8.50 8.30 8.82 7.93 8.14 16.4 9.90 5.58 10.9 Na2O 4.74 4.06 4.29 3.64 3.41 1.25 3.43 3.76 2.01 K2O 0.11 0.10 0.22 0.20 0.46 0.75 0.07 0.24 0.14 P2O5 0.12 0.13 0.09 0.22 0.03 0.03 0.10 0.07 0.04 LOI 2.52 0.73 1.08 0.90 1.23 1.51 2.12 3.63 1.22 Total 99.83 99.85 99.84 99.86 99.89 99.58 99.86 99.87 99.90 Tables 130

Table 5.4: Trace elements and isotope ratios of the Oman plagiogranites (elements in ppm).

P1 OM11- OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample RU - 7C RU - 7E TA - 15A TA - 15B TA - 20B SU - 23E Fi - 28B SA - 32A XRF V 106 15.6 11.0 7.80 12.0 14.4 80.6 10.2 XRF Cr 19.6 13.0 10.8 14.6 14.2 9.60 16.1 15.2 XRF Ni 33.9 16.8 8.90 17.1 11.7 23.0 21.6 12.9 XRF Cu 20.0 2.70 14.6 10.5 9.80 16.1 16.0 10.8 XRF Zn 22.0 11.7 11.4 9.40 10.5 19.2 16.6 4.20 XRF Ga 23.3 20.1 19.0 18.5 21.9 23.2 20.2 15.4 XRF Rb 0.90 0.60 1.50 0.90 0.40 1.00 1.00 LOD ICPMS Rb 1.68 0.49 1.36 0.42 0.18 0.78 0.96 0.13 XRF Sr 349 109 150 55.0 38.9 94.7 142 89.6 ICPMS Sr 314 109 149 55.2 42.1 98.2 135 92.6 XRF Y 75.9 83.1 64.3 118 64.5 116 52.1 62.5 ICPMS Y 50.1 78.2 56.7 121 67.3 119 41.6 65.7 XRF Zr 230 366 288 486 208 198 193 245 ICPMS Zr 155 377 293 592 247 220 172 296 XRF Nb 5.70 7.40 7.20 9.00 3.50 10.1 4.80 5.00 ICPMS Nb 3.60 7.06 6.33 10.7 4.75 11.0 3.43 3.18 XRF Ba 221 39.1 67.1 45.1 7.27 150 27.0 7.84 ICPMS La 7.90 16.9 16.5 22.6 8.98 22.0 13.8 11.9 ICPMS Ce 20.2 36.5 38.7 52.0 22.0 52.6 25.2 26.8 ICPMS Pr 3.45 5.55 5.53 7.93 3.75 8.31 3.43 4.35 ICPMS Nd 19.5 29.2 27.0 40.2 22.0 43.4 17.6 23.4 ICPMS Sm 6.22 9.00 7.50 12.2 7.65 14.0 5.35 7.43 ICPMS Eu 2.00 2.69 2.52 1.79 2.53 2.17 1.80 1.91 ICPMS Gd 7.41 10.8 8.24 14.9 9.95 17.0 6.34 9.26 ICPMS Tb 1.41 2.11 1.52 2.93 1.89 3.27 1.20 1.77 ICPMS Dy 9.25 14.1 10.0 20.4 12.7 22.2 7.73 11.7 ICPMS Ho 2.04 3.16 2.17 4.71 2.86 4.90 1.78 2.72 ICPMS Er 5.79 9.03 6.17 13.8 8.06 14.1 4.91 7.74 ICPMS Tm 0.84 1.39 0.92 2.13 1.17 2.14 0.73 1.12 ICPMS Yb 6.18 10.2 6.79 15.15 8.45 15.4 5.59 8.15 ICPMS Lu 0.87 1.54 1.00 2.20 1.23 2.23 0.79 1.16 ICPMS Hf 4.47 9.87 7.45 15.4 7.38 9.98 4.98 8.47 ICPMS Ta 0.26 0.54 0.41 0.77 0.31 0.88 0.21 0.32 XRF Pb 2.00 LOD LOD 0.70 LOD LOD 1.30 LOD ICPMS Th 0.58 1.20 1.38 1.98 1.20 2.46 0.59 1.06 ICPMS U 0.29 0.38 0.38 0.65 0.36 0.67 0.25 0.30

96 Ma 87Sr/86Sr 0.704767 0.704169 96 Ma 143Nd/144Nd 0.512919 0.512925 96 Ma 176Hf/177Hf 0.283209 0.283285 96 Ma εNd 8.05 8.16 96 Ma εNf 17.12 19.83 Tables 131

Table 5.4: continued.

P1 OM11- OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample SA - 32B TA - 11A TA - 11G TA - 19B TA - 13A TA - 13B TA - 13Fd TA - 13Fh XRF V 7.20 88.9 271 81.1 57.0 68.7 215 116 XRF Cr 12.8 11.9 12.9 14.8 15.9 13.7 13.8 14.4 XRF Ni LOD 21.0 38.3 12.5 LOD 9.50 19.5 7.10 XRF Cu 57.0 23.9 8.70 10.3 69.3 8.50 8.10 22.5 XRF Zn 7.70 12.8 10.3 9.90 15.4 14.0 15.2 11.0 XRF Ga 17.5 22.8 17.0 15.8 20.5 20.4 17.3 22.6 XRF Rb 1.10 1.40 0.70 2.30 1.60 0.40 1.30 1.00 ICPMS Rb 0.39 1.18 0.79 2.35 0.81 1.14 1.31 0.82 XRF Sr 100 178 201 229 194 204 169 163 ICPMS Sr 101 170 193 218 170 192 162 171 XRF Y 89.8 47.5 36.5 42.2 57.5 54.3 55.7 41.6 ICPMS Y 91.8 38.9 29.1 35.3 42.3 44.4 43.0 39.2 XRF Zr 302 126 111 136 155 159 104 376 ICPMS Zr 354 112 96 121 141 148 86.7 415 XRF Nb 6.80 4.40 2.20 3.80 5.60 4.10 6.40 4.80 ICPMS Nb 6.26 3.06 1.80 2.70 3.93 4.33 4.40 3.30 XRF Ba 40.7 31.5 37.2 21.1 40.5 39.9 37.3 49.1 ICPMS La 13.9 5.53 4.82 5.37 8.87 9.24 8.11 7.00 ICPMS Ce 32.6 13.6 12.3 14.7 22.7 22.9 17.6 16.6 ICPMS Pr 5.37 2.25 1.92 2.47 3.54 3.53 2.88 2.73 ICPMS Nd 29.7 13.0 10.9 14.0 18.6 19.0 16.4 15.0 ICPMS Sm 9.43 4.50 3.60 4.78 5.72 5.72 5.18 4.65 ICPMS Eu 1.72 1.76 1.42 1.88 2.06 2.23 1.61 1.84 ICPMS Gd 12.4 5.44 4.35 5.59 6.61 6.97 6.19 5.65 ICPMS Tb 2.44 1.05 0.81 1.08 1.25 1.26 1.22 1.05 ICPMS Dy 16.5 6.94 5.41 7.19 7.96 8.41 8.19 6.96 ICPMS Ho 3.76 1.53 1.19 1.57 1.79 1.83 1.79 1.59 ICPMS Er 10.8 4.35 3.26 4.23 4.91 5.09 5.07 4.61 ICPMS Tm 1.65 0.64 0.51 0.63 0.73 0.71 0.75 0.72 ICPMS Yb 11.9 4.80 3.65 4.75 5.22 5.06 5.78 5.35 ICPMS Lu 1.76 0.71 0.49 0.63 0.74 0.70 0.83 0.83 ICPMS Hf 10.5 3.10 2.60 3.78 3.87 3.69 3.00 10.5 ICPMS Ta 0.45 0.21 0.13 0.18 0.24 0.24 0.29 0.26 XRF Pb 0.90 LOD LOD LOD 1.70 LOD LOD LOD ICPMS Th 1.19 0.36 0.36 0.58 0.48 0.45 0.67 0.67 ICPMS U 0.38 0.14 0.17 0.24 0.16 0.19 0.18 0.20

96 Ma 87Sr/86Sr 0.703670 0.704289 0.704131 96 Ma 143Nd/144Nd 0.512930 0.512925 0.512919 96 Ma 176Hf/177Hf 0.283410 0.283269 0.283318 96 Ma εNd 8.25 8.15 8.04 96 Ma εNf 24.24 19.26 20.98 Tables 132

Table 5.4: continued.

P1 OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample HA - 4A HA - 4B HA - 4C RU - 7F TA - 12D TA - 14A TA - 14B XRF V 16.1 21.3 19.2 37.2 28.9 43.0 9.20 XRF Cr 12.5 13.6 10.7 16.2 20.6 11.8 17.8 XRF Ni 10.9 7.70 15.4 20.8 17.0 15.6 14.1 XRF Cu LOD 9.50 1.80 LOD LOD 8.80 LOD XRF Zn 8.00 6.60 25.0 6.10 4.80 7.60 5.10 XRF Ga 14.5 20.6 15.5 16.8 15.3 18.0 14.5 XRF Rb 1.30 0.40 1.90 1.10 0.50 0.40 0.90 XRF Sr 163 73.8 171 73.8 113 106 73.1 XRF Y 112 95.0 59.5 86.3 46.6 37.3 17.0 XRF Zr 277 333 209 327 340 247 299 XRF Nb 4.50 6.20 3.60 6.30 5.90 4.10 5.10 XRF Ba LOD LOD LOD LOD LOD LOD 53.4

P1 OM11- OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample TA - 14C TA - 14D TA - 15C TA - 11C TA - 11H TA - 11i TA - 11J TA - 19A XRF V 13.4 12.5 12.9 35.5 48.0 131 27.0 123 XRF Cr 13.9 14.4 12.4 12.6 26.2 16.3 14.1 15.8 XRF Ni 11.6 12.9 18.0 17.2 21.4 25.3 26.6 21.6 XRF Zn 4.40 6.60 18.3 6.40 4.80 9.20 6.30 9.40 XRF Ga 17.1 17.9 21.0 17.7 17.1 18.5 16.3 17.8 XRF Rb 1.20 1.50 LOD LOD 0.7 1.1 0.8 1.4 XRF Sr 159 188 123 129 280 524 157 217 XRF Y 10.1 41.1 95.3 49.9 24.3 40.0 68.3 47.9 XRF Zr 275 279 439 221 206 196 265 144 XRF Nb 6.40 5.80 9.50 3.30 3.80 5.40 5.20 4.00 XRF Ba 47.9 62.6 57.3 25.2 LOD LOD LOD LOD XRF Pb LOD LOD LOD LOD 1.10 LOD LOD LOD

P1 OM11- OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample TA - 20A SU - 23A SU - 23B SU - 23C SU - 23F SU - 23G SU - 23H Hi - 26A XRF V 13.8 10.6 8.10 4.90 7.20 6.70 25.8 9.60 XRF Cr 15.4 16.2 14.8 18.0 18.7 13.8 15.7 10.3 XRF Ni 16.0 14.9 6.90 13.0 15.0 11.6 15.6 18.3 XRF Cu LOD 1.50 LOD LOD LOD LOD LOD LOD XRF Zn 28.6 8.40 3.90 3.90 6.20 7.50 29.7 18.0 XRF Ga 19.2 20.1 16.4 17.6 21.6 17.4 20.0 14.2 XRF Rb 1.2 LOD 0.3 0.6 0.4 0.9 1.7 1 XRF Sr 108 54.7 67.3 60.3 101 87.2 137 87.3 XRF Y 83.9 69.3 18.5 83.0 88.4 60.7 64.2 53.7 XRF Zr 282 396 385 314 336 248 247 173 XRF Nb 4.80 8.60 5.70 6.70 7.80 6.30 7.40 3.20 XRF Ba LOD LOD 80.5 86.0 53.9 87.7 190 50.7 XRF Pb 0.70 0.70 LOD LOD LOD LOD LOD LOD Tables 133

Table 5.4: continued.

P1 P2 OM11- OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample Hi - 26B Hi - 26D Fi - 28C SU - 8C SU - 9D Hi - 24A Hi - 25A Fi - 27A XRF V 31.5 24.6 118 202 14.1 65.0 18.9 12.5 XRF Cr 14.7 10.7 13.0 13.9 11.9 11.8 10.4 15.9 XRF Ni 17.7 15.9 25.7 20.0 7.10 LOD 14.6 16.0 XRF Cu LOD LOD LOD 9.30 1.80 81.3 13.6 4.40 XRF Zn 21.9 19.1 16.3 12.6 4.30 13.6 15.6 46.0 XRF Ga 20.1 18.5 23.4 13.1 9.90 17.6 15.2 14.0 XRF Rb 1.5 1.3 1.2 1.20 0.60 1.70 2.30 2.20 ICPMS Rb --- 0.54 0.46 1.64 1.95 1.59 XRF Sr 40.9 58.9 176 276 173 110 261 123 ICPMS Sr --- 266 191 104 270 130 XRF Y 53.9 62.2 48.0 33.3 21.8 42.3 32.0 36.8 ICPMS Y --- 27.0 23.2 36.6 31.1 32.0 XRF Zr 210 203 256 61.9 56.1 98.6 67.1 83.8 ICPMS Zr --- 60.8 64.1 97.0 73.4 75.1 XRF Nb 3.30 4.20 3.90 1.70 0.90 1.70 2.30 2.70 ICPMS Nb --- 0.81 0.77 2.15 1.33 1.98 XRF Ba 48.4 LOD LOD 106 102 36.1 65.4 51.9 ICPMS La --- 2.49 2.99 4.01 3.07 3.43 ICPMS Ce --- 6.03 6.28 9.14 7.08 7.57 ICPMS Pr --- 1.05 1.07 1.47 1.21 1.23 ICPMS Nd --- 6.23 6.22 8.78 7.14 7.10 ICPMS Sm --- 2.67 2.45 3.21 2.73 2.65 ICPMS Eu --- 0.68 0.67 1.06 1.01 0.81 ICPMS Gd --- 3.41 3.23 4.32 3.94 3.72 ICPMS Tb --- 0.68 0.64 0.86 0.80 0.77 ICPMS Dy --- 4.81 4.30 6.09 5.65 5.74 ICPMS Ho --- 1.11 0.99 1.41 1.32 1.33 ICPMS Er --- 3.26 2.78 4.09 3.82 3.96 ICPMS Tm --- 0.51 0.44 0.64 0.58 0.60 ICPMS Yb --- 3.90 3.29 4.81 4.39 4.60 ICPMS Lu --- 0.59 0.46 0.72 0.63 0.66 ICPMS Hf --- 2.11 2.43 3.10 2.53 2.49 ICPMS Ta --- 0.06 0.05 0.14 0.09 0.14 ICPMS Th --- 0.25 0.35 0.35 0.29 0.38 ICPMS U --- 0.21 0.26 0.23 0.18 0.41 Tables 134

Table 5.4: continued.

P2 OM11- OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample Fi - 27C Fi - 29B Fi - 29C Hi - 30A Hi - 30B Hi - 30C SA - 33A SA - 33C XRF V 23.1 26.6 28.6 LOD 292 37.2 30.4 30.4 XRF Cr 13.4 13.2 13.4 LOD 14.6 37.6 13.1 15.7 XRF Ni 13.3 13.6 20.0 LOD 35.0 33.8 14.7 13.2 XRF Cu 3.20 7.10 LOD 117 LOD 13.6 8.20 8.90 XRF Zn 47.0 6.40 8.70 LOD 44.6 28.1 4.90 10.1 XRF Ga 15.1 12.4 13.2 LOD 14.8 14.9 13.7 13.1 XRF Rb 2.80 1.70 1.30 LOD 2.50 2.20 LOD 2.80 ICPMS Rb 2.16 1.44 1.24 4.37 2.72 1.73 0.35 1.69 XRF Sr 109 212 293 LOD 368 278 119 296 ICPMS Sr 108 205 313 196 363 294 134 296 XRF Y 35.6 38.4 35.8 LOD 30.4 44.8 36.8 31.5 ICPMS Y 31.2 31.0 34.8 25.2 24.3 42.0 37.8 26.5 XRF Zr 84.6 62.2 69.1 LOD 79.3 121 76.7 76.6 ICPMS Zr 77.6 58.0 68.9 73.4 68.5 122 86.1 73.2 XRF Nb 2.50 1.20 2.10 LOD 3.40 5.30 2.00 2.20 ICPMS Nb 1.88 1.09 1.26 1.90 1.58 3.37 1.65 1.46 XRF Ba 47.2 30.8 21.7 54.6 40.9 63.5 3.23 37.2 ICPMS La 3.16 2.86 3.77 3.18 4.09 5.39 3.67 2.26 ICPMS Ce 6.58 8.05 6.42 8.29 7.64 12.2 7.71 5.20 ICPMS Pr 1.06 1.46 1.16 1.35 1.26 2.08 1.28 0.91 ICPMS Nd 6.20 8.50 7.15 7.31 7.00 11.5 7.86 5.53 ICPMS Sm 2.37 3.06 2.94 2.38 2.52 4.05 3.04 2.26 ICPMS Eu 0.80 1.02 0.79 0.90 0.91 1.39 1.12 0.89 ICPMS Gd 3.47 4.08 4.27 3.21 3.19 5.30 4.68 3.28 ICPMS Tb 0.73 0.84 0.85 0.66 0.60 1.05 0.94 0.66 ICPMS Dy 5.10 5.69 5.94 4.63 4.28 7.39 6.60 4.66 ICPMS Ho 1.21 1.30 1.40 1.06 1.03 1.65 1.48 1.08 ICPMS Er 3.55 3.85 4.01 3.06 2.87 4.90 4.41 3.13 ICPMS Tm 0.56 0.59 0.61 0.47 0.45 0.73 0.65 0.47 ICPMS Yb 4.23 4.44 4.41 3.56 3.39 5.41 4.89 3.79 ICPMS Lu 0.63 0.64 0.63 0.49 0.46 0.79 0.72 0.55 ICPMS Hf 2.49 2.04 2.46 2.26 2.07 3.61 2.77 2.41 ICPMS Ta 0.14 0.06 0.07 0.12 0.11 0.18 0.11 0.09 XRF Pb 1.10 LOD 1.10 LOD LOD 1.40 LOD LOD ICPMS Th 0.45 0.29 0.33 0.27 0.23 0.44 0.32 0.28 ICPMS U 0.37 0.24 0.19 0.15 0.17 0.23 0.19 0.21

96 Ma 87Sr/86Sr 0.705402 96 Ma 143Nd/144Nd 0.512904 96 Ma 176Hf/177Hf 0.283169 96 Ma εNd 7.74 96 Ma εNf 15.70 Tables 135

Table 5.4: continued.

P2 OM11- OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample SA - 33F HA - 34C SU - 16A SA - 31A SA - 31E HA - 34D HA - 34F HA - 34G XRF V 11.3 21.1 9.20 5.20 8.30 25.2 21.8 22.1 XRF Cr 15.8 11.7 11.1 13.0 13.5 15.8 27.7 20.8 XRF Ni 11.9 13.9 7.00 10.4 14.8 1.70 16.3 9.90 XRF Cu 12.0 1.90 6.20 6.80 5.70 48.2 12.7 5.50 XRF Zn 8.20 12.0 31.9 33.8 34.6 9.70 28.7 6.10 XRF Ga 12.0 12.9 13.7 11.7 13.7 15.4 15.7 13.7 XRF Rb 1.90 0.90 4.00 1.40 1.40 0.50 0.60 1.50 ICPMS Rb 1.45 0.97 3.58 1.04 1.07 0.43 0.26 0.72 XRF Sr 290 140 115 89.5 76.6 115 96.2 108 ICPMS Sr 305 148 124 92.2 79.3 120 96.8 109 XRF Y 31.4 43.4 41.5 39.0 40.3 47.6 40.0 50.6 ICPMS Y 31.6 43.6 44.2 40.1 40.7 45.3 39.4 47.6 XRF Zr 78.2 113 128 91.9 96.4 127 102 117 ICPMS Zr 81.4 122 150 103 108 131 113 128 XRF Nb 3.50 2.60 3.10 3.30 2.20 3.50 3.30 1.90 ICPMS Nb 1.69 1.70 2.70 2.30 2.33 1.68 1.54 1.71 XRF Ba 20.8 26.3 123 70.2 53.9 31.2 22.0 27.0 ICPMS La 2.75 4.35 5.27 5.41 3.80 5.19 3.91 6.27 ICPMS Ce 6.24 11.1 12.0 8.97 8.75 13.0 9.04 16.6 ICPMS Pr 1.19 2.04 2.06 1.56 1.55 2.42 1.62 2.95 ICPMS Nd 7.08 11.7 12.2 9.65 9.48 13.8 9.84 17.0 ICPMS Sm 2.87 4.50 4.78 3.82 3.78 5.10 3.61 5.65 ICPMS Eu 1.01 1.39 1.92 1.24 1.23 1.52 0.96 1.41 ICPMS Gd 4.12 5.83 6.15 5.09 5.43 6.39 4.99 7.25 ICPMS Tb 0.82 1.10 1.22 1.07 1.09 1.26 0.98 1.26 ICPMS Dy 5.62 7.83 7.97 7.57 7.44 8.31 6.71 8.78 ICPMS Ho 1.28 1.78 1.85 1.71 1.73 1.97 1.52 1.98 ICPMS Er 3.78 5.29 5.53 4.98 5.06 5.50 4.50 5.72 ICPMS Tm 0.56 0.76 0.86 0.77 0.75 0.85 0.70 0.90 ICPMS Yb 4.33 5.39 6.51 5.60 5.49 6.15 4.92 6.30 ICPMS Lu 0.65 0.78 0.96 0.82 0.80 0.88 0.77 0.96 ICPMS Hf 2.77 4.11 4.86 3.60 3.48 4.56 3.76 4.26 ICPMS Ta 0.11 0.12 0.18 0.15 0.15 0.14 0.11 0.13 XRF Pb 1.20 LOD LOD LOD LOD LOD 2.10 LOD ICPMS Th 0.35 0.60 0.42 0.39 0.36 0.63 0.51 0.61 ICPMS U 0.19 0.31 0.26 0.16 0.17 0.22 0.20 0.22

96 Ma 87Sr/86Sr 0.704407 96 Ma 143Nd/144Nd 0.512910 96 Ma 176Hf/177Hf 0.283281 96 Ma εNd 7.87 96 Ma εNf 19.67 Tables 136

Table 5.4: continued.

P2 OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample SU - 17D SU - 18A SU - 18B SU - 18C HA - 34A HA - 34B HA - 34E XRF V 85.2 26.8 15.5 30.5 20.1 12.1 23.0 XRF Cr 134 10.5 13.9 12.6 15.8 10.0 16.1 XRF Ni 48.0 11.8 8.10 15.3 19.3 6.2 11.0 XRF Cu LOD 2.60 3.40 9.20 LOD LOD LOD XRF Zn 32.6 8.90 9.10 15.7 19.5 7.00 6.30 XRF Ga 12.3 15.0 11.5 14.1 14.8 12.5 13.3 XRF Rb 0.40 0.70 0.90 LOD 1.00 0.90 0.80 XRF Sr 114 282 212 204 154 116 103 XRF Y 18.4 9.30 11.5 16.2 43.1 45.9 55.8 XRF Zr 89.2 15.2 30.7 43.0 126 143 126 XRF Nb 1.90 1.00 1.50 1.80 1.80 1.50 1.80 XRF Ba 75.4 LOD LOD LOD 48.9 63.6 LOD XRF Pb LOD 3.10 1.80 LOD LOD LOD LOD

P2 OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample HA - 34H SU - 8D SU - 8F SU - 8G SU - 8H SU - 9B SU - 22 XRF V 30.0 219 166 198 44.9 16.9 4.90 XRF Cr 16.6 15.5 13.9 8.40 11.7 13.5 17.2 XRF Ni 21.3 27.4 20.5 13.4 10.5 14.3 7.70 XRF Cu LOD LOD LOD LOD LOD LOD LOD XRF Zn 9.20 41.3 30.1 34.8 8.30 5.30 7.50 XRF Ga 14.5 13.7 12.3 15.4 12.3 11.4 12.0 XRF Rb 0.90 1.10 1.30 0.90 1.40 0.50 36.3 XRF Sr 116 219 250 274 177 287 18.7 XRF Y 38.3 18.2 36.3 23.0 27.1 21.7 22.3 XRF Zr 105 53.5 53.4 62.2 39.6 53.7 92.4 XRF Nb 1.90 2.40 1.40 1.70 1.80 2.20 3.80 XRF Ba 38.0 86.3 64.8 96.0 LOD LOD LOD XRF Pb LOD LOD LOD LOD LOD 1.40 5.00

P2 OM11- OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample Hi - 24C Fi - 27B Fi - 29A Fi - 29D SA - 31B SA - 31D SA - 33D SA - 33G XRF V 17.8 14.0 141 16.4 4.60 3.70 21.3 24.0 XRF Cr 9.80 10.7 17.3 12.4 11.8 18.6 11.6 18.1 XRF Ni 4.30 7.90 20.7 12.4 7.60 14.1 9.90 19.8 XRF Cu LOD LOD LOD LOD 32.4 LOD LOD LOD XRF Zn 9.40 51.2 10.5 7.90 29.9 30.9 8.20 13.1 XRF Ga 11.6 14.8 17.6 13.9 13.3 12.8 10.9 12.5 XRF Rb 1.80 2.90 1.70 0.90 2.40 1.60 1.90 1.60 XRF Sr 114 122 216 165 87.0 85.5 193 269 XRF Y 48.7 36.1 40.6 20.7 47.7 42.5 34.3 36.4 XRF Zr 113 86.3 79.4 76.5 118 143 69.3 83.3 XRF Nb 3.10 3.70 1.20 0.70 2.20 5.30 1.20 1.90 XRF Ba LOD LOD 54.4 LOD LOD LOD 44.8 41.6 XRF Pb LOD LOD LOD LOD 1.80 1.90 LOD LOD Tables 137

Table 5.4: continued.

wall-rocks OM11- OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample RU - 7D RU - 7A RU - 7B RU - 7G RU - 7H SU - 8A SU - 8B SU - 8E XRF V 269 305 226 195 150 51.6 46.1 226 XRF Cr 27.9 26.8 109 631 473 11.5 10.5 1074 XRF Ni 28.3 42.2 74.1 235 169 LOD 9.00 190 XRF Cu 17.1 6.80 3.20 65.0 40.9 67.1 9.90 37.6 XRF Zn 17.9 34.7 54.5 71.8 52.3 8.80 4.70 31.3 XRF Ga 17.9 17.8 15.3 11.3 10.2 14.0 12.1 7.00 XRF Rb 0.90 1.50 0.90 0.80 0.60 0.50 0.70 LOD ICPMS Rb 1.34 1.10 1.17 0.88 0.98 0.25 0.36 0.26 XRF Sr 394 320 220 165 134 206 185 54.9 ICPMS Sr 385 314 224 158 125 220 187 63.0 XRF Y 34.5 49.3 26.7 14.3 13.4 13.8 14.1 6.10 ICPMS Y 28.9 40.1 20.5 9.38 9.02 12.7 12.8 6.76 XRF Zr 98.7 120 74.5 32.4 22.4 38.3 25.3 9.60 ICPMS Zr 87.2 102 56.4 23.4 14.4 40.1 23.7 6.41 XRF Nb 2.60 5.90 3.00 1.10 1.20 1.90 0.60 0.90 ICPMS Nb 2.04 3.51 1.40 0.64 0.55 0.71 0.64 0.11 XRF Ba 219 103 38.4 153 27.1 64.9 74.3 11.9 ICPMS La 4.57 6.48 3.31 3.14 1.20 1.80 2.06 3.36 ICPMS Ce 11.5 15.0 7.97 7.53 2.21 3.27 3.30 0.84 ICPMS Pr 1.79 2.51 1.23 1.03 0.40 0.54 0.54 0.17 ICPMS Nd 9.59 14.1 6.58 4.36 2.32 3.35 3.37 1.15 ICPMS Sm 3.35 4.84 2.17 1.12 0.93 1.17 1.21 0.54 ICPMS Eu 1.49 1.74 0.93 0.59 0.54 0.72 0.73 0.27 ICPMS Gd 3.96 5.98 2.71 1.25 1.27 1.59 1.68 0.88 ICPMS Tb 0.76 1.14 0.55 0.25 0.25 0.30 0.30 0.19 ICPMS Dy 5.36 7.68 3.76 1.73 1.76 2.07 2.13 1.35 ICPMS Ho 1.20 1.72 0.82 0.39 0.40 0.50 0.47 0.30 ICPMS Er 3.35 4.66 2.37 1.05 1.06 1.34 1.40 0.83 ICPMS Tm 0.49 0.70 0.35 0.15 0.16 0.24 0.22 0.12 ICPMS Yb 3.78 5.01 2.45 1.23 1.26 1.73 1.65 0.93 ICPMS Lu 0.53 0.71 0.37 0.17 0.16 0.26 0.24 0.12 ICPMS Hf 2.53 2.88 1.67 0.71 0.52 1.48 0.95 0.25 ICPMS Ta 0.16 0.27 0.11 0.07 0.06 0.06 0.05 0.01 XRF Pb LOD LOD LOD LOD LOD 0.90 1.60 LOD ICPMS Th 0.30 0.34 0.28 0.13 0.12 0.28 0.22 0.04 ICPMS U 0.16 0.13 0.13 0.03 0.03 0.16 0.15 0.02

96 Ma 87Sr/86Sr 0.705375 96 Ma 143Nd/144Nd 0.512897 96 Ma 176Hf/177Hf 0.283191 96 Ma εNd 7.61 96 Ma εNf 16.50 Tables 138

Table 5.4: continued.

wall-rocks OM11- OM11- OM11- OM11- OM11- OM11- OM11- OM11- sample SU - 9A SU - 9C TA - 11F TA - 12A TA - 12B TA - 13C TA - 21B Hi - 25B XRF V 16.0 258 289 355 437 332 365 233 XRF Cr 9.40 11.8 46.0 17.0 28.6 203 16.5 107 XRF Ni 6.70 12.6 47.4 22.3 36.7 70.8 13.1 47.6 XRF Cu LOD 49.7 4.30 1.90 10.9 6.60 13.2 LOD XRF Zn 5.30 46.2 7.90 15.8 18.9 19.0 29.3 16.4 XRF Ga 11.9 15.7 14.9 19.2 17.4 13.9 18.1 11.6 XRF Rb 1.30 1.80 1.60 LOD 0.50 1.90 2.20 4.00 ICPMS Rb 0.52 1.39 0.62 0.32 0.28 0.72 0.84 2.79 XRF Sr 228 265 178 551 183 183 208 119 ICPMS Sr 239 264 180 531 185 186 227 118 XRF Y 23.4 20.6 35.0 36.7 36.2 37.3 50.9 23 ICPMS Y 23.8 17.0 26.8 26.7 27.9 29.0 37.9 18.4 XRF Zr 58.7 42.1 78.6 115 92.2 74.3 140 45.5 ICPMS Zr 65.7 36.6 72.8 77.4 77.6 63.2 121 39.7 XRF Nb 1.90 1.80 1.60 4.30 4.30 2.00 4.60 LOD ICPMS Nb 0.74 0.50 1.49 2.15 2.79 1.58 3.33 0.60 XRF Ba 76.1 93.0 23.4 32.3 25.7 30.5 22.3 67.8 ICPMS La 2.15 1.68 4.61 4.29 4.32 6.37 6.30 3.97 ICPMS Ce 4.26 3.67 11.8 10.5 12.1 11.4 16.5 2.99 ICPMS Pr 0.72 0.63 1.94 1.65 1.91 1.84 2.64 0.57 ICPMS Nd 4.52 3.66 10.9 9.18 10.4 10.1 13.9 3.45 ICPMS Sm 1.89 1.51 3.53 3.18 3.36 3.38 4.61 1.47 ICPMS Eu 0.63 0.56 1.19 1.42 1.38 1.35 1.67 0.46 ICPMS Gd 3.05 2.31 4.14 4.01 4.16 4.22 5.49 2.04 ICPMS Tb 0.63 0.46 0.78 0.79 0.75 0.79 1.03 0.45 ICPMS Dy 4.24 3.26 5.09 5.07 5.45 5.42 6.85 3.20 ICPMS Ho 0.98 0.74 1.11 1.13 1.17 1.21 1.49 0.73 ICPMS Er 2.87 2.10 2.98 3.29 3.26 3.44 4.25 2.13 ICPMS Tm 0.46 0.32 0.42 0.47 0.52 0.50 0.64 0.32 ICPMS Yb 3.41 2.55 3.14 3.46 3.54 3.85 4.63 2.27 ICPMS Lu 0.50 0.37 0.42 0.51 0.47 0.51 0.65 0.34 ICPMS Hf 2.52 1.44 2.08 2.08 2.14 1.84 3.07 1.36 ICPMS Ta 0.05 0.04 0.11 0.14 0.16 0.11 0.20 0.04 XRF Pb LOD 3.80 LOD LOD LOD LOD LOD LOD ICPMS Th 0.36 0.19 0.37 0.39 0.26 0.28 0.39 0.15 ICPMS U 0.23 0.18 0.11 0.15 0.12 0.09 0.15 0.09

96 Ma 87Sr/86Sr 0.705970 0.705996 0.703521 96 Ma 143Nd/144Nd 0.512918 0.512911 0.512934 96 Ma 176Hf/177Hf 0.283174 0.283241 0.283261 96 Ma εNd 8.03 7.88 8.33 96 Ma εNf 15.90 18.26 18.96 Tables 139

Table 5.4: continued.

wall-rock OM11- OM11- OM11- OM11- sample Hi - 26C Fi - 28A Fi - 29E SA - 33E ppm XRF V 134 275 382 259 XRF Cr 2017 93.3 13.4 105 XRF Ni 217 61.0 26.2 53.1 XRF Cu 76.3 3.20 23.1 54.5 XRF Zn 20.3 18.6 9.60 33.3 XRF Ga 7.20 14.4 16.2 14.1 XRF Rb 19.0 1.30 1.20 0.60 ICPMS Rb 15.9 0.44 1.06 0.76 XRF Sr 138 210 317 104 ICPMS Sr 153 208 290 100 XRF Y 3.90 34.3 28.1 14.6 ICPMS Y 5.22 25.3 20.6 10.9 XRF Zr 6.40 79.4 61.9 26.3 ICPMS Zr 4.57 62.9 49.2 21.2 XRF Nb 2.10 2.40 1.70 0.80 ICPMS Nb 1.31 1.43 1.18 0.54 XRF Ba 75.4 31.1 17.0 14.7 ICPMS La 1.92 3.03 2.19 1.98 ICPMS Ce 3.95 9.50 5.11 2.78 ICPMS Pr 0.42 1.67 0.87 0.47 ICPMS Nd 1.67 9.72 5.07 2.56 ICPMS Sm 0.56 3.29 2.00 1.05 ICPMS Eu 0.31 1.20 0.75 0.47 ICPMS Gd 0.84 3.94 2.72 1.46 ICPMS Tb 0.15 0.74 0.55 0.32 ICPMS Dy 1.02 5.07 3.77 2.14 ICPMS Ho 0.22 1.10 0.86 0.46 ICPMS Er 0.58 2.90 2.38 1.32 ICPMS Tm 0.09 0.44 0.35 0.23 ICPMS Yb 0.69 3.17 2.72 1.57 ICPMS Lu 0.08 0.45 0.38 0.21 ICPMS Hf 0.24 1.82 1.52 0.73 ICPMS Ta 0.11 0.10 0.07 0.05 XRF Pb LOD LOD LOD LOD ICPMS Th 0.79 0.31 0.18 0.11 ICPMS U 0.33 0.07 0.12 0.05

96 Ma 87Sr/86Sr 0.707521 96 Ma 143Nd/144Nd 0.512914 96 Ma 176Hf/177Hf 0.283144 96 Ma εNd 7.95 96 Ma εNf 14.84 6 Findings and Outlook

Ophiolite complexes represent oceanic crust obducted on land and provide therefore a comparable easy way to study oceanic crust processes. Small amounts of felsic rocks can be found in recent oceanic crust and are also common in ophiolite complexes. The majority of previous studies dealing with felsic intrusive rocks from ophiolites confine the sampling to a spatial constricted area or just one single sampling locality (e.g. Cox et al. 1999, Gillis and Coogan, 2002; Tsuchiya et al. 2013). Others handle few plagiogranite data within studies mainly dealing with mafic crustal rocks or mention them only ancillary (e.g. Juteau et al. 1988; Bickle and Teagle, 1992; Nicolas et al. 2000; Yamaoka et al. 2012; MacLeod, 2013; Oeser et al. 2013). In contrast, the main aim of this study deals with the petrography and geochemistry of felsic rocks. In chapter 3, fresh, young glassy lavas sampled from the top of the PAR ridge-axis imply that felsic (andesitic to dacitic) magmas are not only restricted to arc-settings, segment-ends (e.g. Wanless et al. 2010, 2011) or crustal intrusive zones (e.g. Dick et al. 2000; Koepke et al. 2007) but also exist beneath normal mid-ocean-ridges and even have the potential to erupt at the ocean floor. Chapters 4 and 5 aim at investigating several plagiogranites from two large and well-preserved ophiolite complexes in Cyprus and in the Oman. I show that there are distinct differences between the felsic rocks from these complexes and their relations with their mafic crustal rocks but also that the felsic intrusives provide important information about mantle contamination and the temporal evolution of magmatic phases.

Evolution of felsic melts within oceanic crust: observations at a recent MOR

My data presented in chapter 3 support the idea that the evolved samples may be explained by assimilation-fractional crystallization processes in the upper magma chamber of the Pacific Antarctic Rise, recharged by basaltic magma ascending from a deeper melt pool and thitherto have been unaffected by assimilation. The felsic lavas have low δ18O isotope values between 5.6‰ and 5.1‰ and high Cl contents indicating assimilation of hydrothermally altered material whereas basalts from the same ridge section have normal MORB δ18O values and low Cl contents. Clinopyroxene-melt barometry indicates crystal fractionation of the basaltic melt in deep sills, whereas the hydrothermally altered material that is assimilated is located in the upper crust. Fractional-crystallization of the melt 6. Findings and Outlook 141 in the upper melt lens is accompanied by a reduction of the oxygen fugacity causing sulphide saturation and significant loss of S, Cu and Co from the felsic melt. The release of heat during crystallization in the shallow crust magma chamber most likely causes the breakdown and partial melting of amphibole-bearing altered rocks. The evolved lavas along the PAR occur at a ridge-segment that is significantly influenced by the Foundation mantle plume. Thus, I infer that increased magma supply leads to a thickened crust, multiple magma systems and increased hydrothermal circulation cause the formation of relatively large volumes of felsic magmas within this MOR crust system.

Chapter 3 is the first study that deals with a representative number of fresh, basalts next to fresh felsic glassy lava samples from a complete ridge segment of the Pacific Antarctic Rise. I can show that assimilation of significant amounts of hydrothermally altered crustal rock must be an essential process during the evolution of the evolved mid- ocean ridge lavas. The unusual thickened crust next to the Foundation plume intersection is essential for comparatively long magma storing periods, more intense hydrothermal cooling and therefore advanced crystal fractionation and assimilation processes of the melt in a shallow magma chamber.

In the future, more seismic studies of recent ocean-ridge settings possessing felsic lavas can provide further information about the structure of the crust and spatial positioning of melt pools promoting advanced melt evolution. Furthermore, a detailed comparisons of geochemical and isotope compositions of crustal rocks and lavas can give important evidence for the understanding of magma evolution during its residence time within the crust.

Evolution of felsic melts within oceanic crust: plagiogranites in ophiolites

In chapter 4 and 5, I present high quality data of plagiogranites sampled from the entire Troodos and the Oman ophiolite complexes and compare them with mafic crustal rocks to discover genetic relations.

Chapter 4 deals with eighty-five felsic samples from eight different intrusions in the Troodos ophiolite and investigates the genetic relations with other crustal rocks studied so far. The plagiogranitic samples from the Troodos ophiolite have higher SiO2 contents than the majority of the Troodos glasses but extend the lava trends (in major element variation diagrams) indicating that the felsic intrusions generally represent crystallized magmas. I find that the various plagiogranite bodies in the Troodos ophiolite can be divided into three groups based on their incompatible element compositions. Two of these felsic intrusions are genetically linked to the mafic crustal rocks, including the sheeted dyke complex, implying that the magmas formed at a spreading axis. Ascending tholeiitic melts produced a (basaltic-andesitic) sheeted dyke complex and further fractional crystallization of this 6. Findings and Outlook 142 magma produced the most voluminous intrusions of the first group of plagiogranites, whereas the second group occurs in one intrusion that is chemically related to off-axis and late-stage boninitic lavas and dykes. The third plagiogranite group has apparently no genetic relations to other crustal rocks of the Troodos ophiolite. Partial melting of hydrothermally altered crustal rocks would fractionate the incompatible elements which is not observed in our data and thus I rule out significant partial melting during the formation of the Troodos plagiogranites. Our data indicate that highly variable mantle melts (mafic tholeiites and boninites) ascend separately into shallow magma lenses in the upper crust and stagnate to evolve to felsic melts that now form plagiogranite intrusions.

Comparable results yielded the data of eighty-five felsic samples from about thirty in- trusions in seven blocks of the Oman ophiolite. In chapter 5 I present the plagiogranite data displaying well-defined major and trace element trends and Hf, Nd and Sr isotope ratios, suggesting that these are genetically related with their associated mafic intrusives and lavas by fractional crystallization although assimilation of hydrothermally altered rocks may have occurred. The plagiogranites display two different compositional groups, both intruded into the crust of the entire length of the ophiolite. The two groups can be correlated to the two major magmatic phases expressed by the V1 and V2 lavas. The second group (plagiogranites and lavas) indicates a stronger depletion of the mantle source and an increasing influence of a sedimentary slab component compared to the first group rocks including the sheeted dykes. A correlation of the Hf and Nd isotope ratios imply a re-enrichment of the depleted mantle wedge by release of sediment melts during transition of the early towards the late magmatic stage of the subduction zone related Oman ophiolite crust. Correlation of well-constrained age-dated plagiogranitic sample locations compared with the two magmatic phases represented by the two plagiogranite and lava groups imply that the magmatic evolution, including the mantle depletion and an increasing subduction zone influence occurred within a short time period between 96.5 and 95.5 Ma.

To summarize, the plagiogranite intrusions in the two studied ophiolite complexes are related to the mafic crustal rocks by having the same parental magmas. In both cases the ophiolite complexes comprise several magmatic phases reflected by subsequent depletion in incompatible element compositions not only in the mafic rocks but also in the felsic intrusions. In the case of the Oman ophiolite plagiogranites the isotope ratios demonstrate a sediment melt contamination of the mantle source and therefore give evidence for a subduction zone component in the entire crust. In the Troodos ophiolite one plagiogranite group is genetically related with boninitic melts and demonstrates that these melts can fractionate toward felsic composition.

This is the first study dealing with complete geochemical data sets (bulk rock major-, trace elements, isotope ratios and mineral analyses) of numerous felsic intrusions from the entire length of two of the largest and best-preserved ophiolite complexes, which allow to 6. Findings and Outlook 143 better determine the evolution and relations with other crustal rocks.

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170 Title: Oxygen isotope evidence for the formation of andesitic-dacitic magmas from the fast-spreading Pacific-Antarctic Rise by assimilation-fractional crystallisation

METHODS

The samples were recovered by dredging and wax coring along the neovolcanic zone of the

PAR between 36°S and 40°S during expeditions SO100 (cruise in year 1995; 11 evolved lavas 3 basaltic lavas), SO157 (cruise in year 2001; 25 evolved lavas, 26 basaltic lavas),

SO213 (cruise in year 2010/2011, 4 evolved lavas, 4 basaltic lavas) with the German research vessel RV SONNE and Atalante FH (cruise in year 1997, 1 basaltic lava) (Fig. 1, Table 1).

All samples are from pillow or sheet flows with fresh glassy outer rims without evidence of hydrothermal alteration. Thin hydrothermal sediments covering the lavas were observed frequently in camera tows but were avoided during sample preparation. The freshest glasses were crushed, handpicked, washed with deionised water and used for geochemical analyses.

Eight representative thin sections (same pillows but some cm below the glassy outer rim) were studied petrographically.

The major element concentrations of the SO100, SO157 and Atalante FH glass samples were measured on a JEOL JXA 8900 Superprobe electron microprobe (EMP) at the

University of Kiel following the methods and conditions outlined by Fretzdorff et al. (2004).

The major element concentrations of the SO213 glasses and the minerals (thin sections of six evolved lavas and two basalts) were measured on a JEOL JXA 8200 Superprobe electron

T microprobe at the GeoZentrum Nordbayern, Erlangen. SiO2, TiO2, Al2O3, FeO , MnO, MgO,

CaO, Na2O, K2O, P2O5 and, additionally for glasses, SO2, Cl and F (Table 1 and

Supplementary table 1) were measured and data quality was checked by running standards

VG-2 and A-99 together with the samples (for further details see Brandle et al. (2012)). The EMP in Erlangen was operated at an accelerating voltage of 15 kV, a beam current of 12 nA and a defocused beam diameter (12 µm) for the glasses and a focused beam for mineral analyses. Counting times were set to 20 s and 10 s for peaks and backgrounds for all elements, and 40 and 20 s for Cl and F, respectively. Water (5 evolved glasses, 7 basaltic glasses) was analyzed in the glass rims at the University of Kiel using a Bruker IFS 66v/S

FTIR spectrometer following the methods outlined by Fretzdorff et al. (2002).

Trace elements of 17 evolved glasses and 13 basaltic glasses (SO100, SO157, Atalante

FH) were analyzed using an Agilent 7500c/s Quadrupole Inductively Coupled Plasma Mass

Spectrometer (ICP-MS) at the Institut für Geowissenschaften, Universität Kiel following procedures described previously from Garbe-Schönberg (1993). Repeated analyses of international rock standards (e.g. BHVO and BIR, see supplementary table 1) give a standard deviation for precision and accuracy of <5 % and <8 % (2!), respectively.

Trace element analyses of 2 evolved glasses and 2 basaltic glasses (SO213) were analyzed using a Thermo X-Series 2 quadrupole ICP-MS at the GeoZentrum Nordbayern in

Erlangen. Approximately 50 mg of sample powder was accurately weighed into a Teflon beaker, and digested in 3 ml of a 3:1 mixture of 12M HF 15M HNO3 at 100°C overnight. The sample solution was evaporated to incipient dryness, treated with 1ml 15M HNO3; this step was repeated 3 times until the sample was fully in solution. One drop of 12M HF, 4 ml 15M

HNO3 and 4 ml H2O were then added to the sample, which was quantitatively transferred to a

250 ml bottle and diluted with water to obtain a dilution factor of approximately 4000 and a final HNO3 concentration of 2 % containing trace HF. All steps were carried out using ultrapure reagents in Class 1000 filtered air.

Multi-element standards containing all the elements of interest, with concentrations spanning the range in sample solutions (0.05 - 50 ppb) were prepared immediately before each analytical session, and used to determine calibration curves for each element. Rare earth oxide correction factors were determined before each analytical session; the CeO/Ce ratio was always in the range 0.012-0.015, and other interferences were found to be insignificant.

Instrument drift was monitored using a 10 ppb Be-Rh-In-Bi internal standard which was mixed with sample solutions online, and a drift monitor which was measured after every 5 samples. A blank correction was applied using a procedural blank which was processed together with each batch of samples, blank corrections were negligable for most elements.

Accuracy and precision for most elements, determined by repeated analysis of rock standards over a period of several months, are in the range 2-5 %, except for Sc, Cs, Th, Ta (5-8%).

Repeated analyses of international rock standards (BHVO and BIR) are given in the supplementary table 1.

Strontium (Haase et al., 2005) and Nd isotope ratios were analyzed at GEOMAR,

Helmholtz-Zentrum für Meereswissenschaften following the methods described in Fretzdorff et al. (2004). Handpicked glass chips were leached for one hour in hot ultrapure 6N HCl before dissolution to avoid any potential influence from seawater alteration. The dissolution and separation techniques are described in detail in Hoernle et al. (1991). Strontium and Nd isotope ratios (of 8 evolved glasses and 10 basaltic glasses) were analyzed in static mode on a

Finnigan MAT 262 mass spectrometer at IFM-GEOMAR, Kiel. The fractionation corrections applied for Sr isotope ratios was 86Sr/88Sr = 0.1194 and 146Nd/144Nd = 0.7219, with repeated measurements of NBS 987 (n = 12) yielding 87Sr/86Sr = 0.710218 (2! = 0.000024). Repeat measurements of the Nd Spex standard (n = 10) and of the La Jolla standard (n = 3) gave an average 143Nd/144Nd of 0.511710 (2! = 0.000015) and 143Nd/144Nd = 0.511827 (2! =

0.000007), respectively. Strontium (supplementary table 1) and Nd analyses (Table 1) were normalized to values of NBS 987 and La Jolla of 0.71025 and 0.511855, respectively.

Glass chips for oxygen isotope analyses were handpicked from crushed sample material. A sample aliquot was checked for purity using X-ray diffraction (XRD). Oxygen isotope ratios were determined on 2.0-2.6 mg sample material weighed before being put into the sample holder. The !18O isotope ratios (of 9 evolved glasses and 5 basaltic glasses) were measured by laser fluorination using a 25 W-Synrad CO2-laser and F2 as reagent, at the

GeoZentrum Nordbayern, Erlangen. The laser was operated in continuous mode and the energy manually adjusted to allow a reaction process as smooth as possible to avoid any sputtering of the sample. The general setup follows that of Sharp (1990). The stainless steel sample chamber is covered with a BaF2-window and the stainless steel sample holder has 16 pits with a diameter and a depth of 3 mm, respectively. Samples were pre-fluorinated at 0.02 bar with F2 overnight. Fluorine pressure during fluorination was 0.1 bar. After reaction in the sample chamber, the gases were slowly passed through an LN2-trap, staying for 3 minutes in a

NaCl-trap heated up to 150-170 °C. The resulting Cl2 gas was trapped in the next segment of the line in another LN2-trap for 3 minutes. During the next step the purified gases frozen onto an LN2 cooled molecular sieve (0.5 nm) in a teflon valved Pyrex tube (3 minutes). This tube then was disconnected from the line and analysed on a Thermo Finnigan Delta Plus masspectrometer at the GeoZentrum Nordbayern. During the period of an entire day, four standard samples (UWG-2, (Valley et al., 1995); NBS-30) where processed and measured together with the samples to ensure accuracy. The long-term reproducibility of the UWG-2 garnet standard is 5.85 ± 0.15 ‰ (1SD, n=67), which is similar to the value of 5.74 ± 0.15 ‰

(1SD, n=1000) given by Valley et al. (1995). The !18O raw values of a run where adjusted by the mean difference of the reference values of the standards (!18O 5.8 ‰ and !18O 5.1 ‰, respectively). Reproducibility of replicates from the samples, which may reflect internal analytical error but also sample heterogeneity and impurity, varies between 0.05 ‰ 1SD and

0.15 ‰ 1SD (mean 0.08 ‰ 1SD), respectively.

Chapter 3: Oxygen isotope evidence for the formation of andesitic-dacitic magmas from the fast-spreading Pacific-Antarctic Rise by assimilation-fractional crystallisation Supplementary Table 3.1 Major elemenets northern andesitic glasses cruise SO100 SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO157 sample 105DS1 6DS1 6DS2 5DS1 5DS5 5DS2 15DS1 15DS2 15DS3 15DS4 4DS5 4DS1 Latitude 37.191 37.562 37.562 37.606 37.606 37.606 37.618 37.618 37.618 37.618 37.635 37.635 Long -110.707 -110.826 -110.826 -110.856 -110.856 -110.856 -110.863 -110.863 -110.863 -110.863 -110.868 -110.868 water depth (m) 2409 2246 2246 2238 2238 2238 2210 2210 2210 2210 2213 2213 material glass glass glass glass glass glass glass glass glass glass glass glass (wt.%)

SiO2 55.5 57.7 57.2 56.4 55.9 55.8 58.9 58.4 58.5 59.0 62.1 60.6

TiO2 2.09 1.55 1.57 1.60 1.58 1.56 1.42 1.46 1.46 1.43 1.03 1.28

Al2O3 12.2 12.7 12.7 13.9 13.9 14.0 13.7 13.5 13.4 13.6 13.3 13.3 FeOT 15.2 12.0 12.0 10.5 10.3 10.3 9.79 10.2 10.3 10.2 8.35 9.06 MnO 0.265 0.207 0.199 0.166 0.139 0.131 0.165 0.151 0.175 0.183 0.157 0.147 MgO 2.03 1.90 2.01 3.12 3.06 3.07 1.99 1.96 1.99 2.00 1.67 1.65 CaO 6.50 5.64 5.71 6.77 6.76 6.67 5.45 5.43 5.61 5.58 4.76 4.76

Na2O 3.35 4.40 4.66 3.98 4.16 4.15 4.59 4.51 4.48 4.41 4.66 4.50

K2O 0.633 0.738 0.736 0.787 0.806 0.809 1.01 1.03 1.02 1.02 1.20 1.11

P2O5 0.813 0.555 0.540 0.384 0.400 0.385 0.484 0.493 0.502 0.467 0.281 0.390

Cr2O3 0.029 0.038 0.017 0.033 0.020 0.035 0.024

SO2 0.164 0.170 0.161 0.151 0.120 0.147 0.086 0.108 0.087 0.097 0.039 0.090 Cl 0.151 1.15 1.14 0.577 0.591 0.590 0.972 0.943 0.950 0.923 0.953 1.01 F

H2O 0.510 Total 99.02 98.77 98.60 98.43 97.73 97.63 98.52 98.23 98.57 98.87 98.53 97.95 cruise SO157 SO157 SO157 SO100 SO100 SO100 SO157 SO157 SO157 SO157 SO100 SO100 SO100 sample 3DS4 3DS6 3DS1 86DS5 86DS1 86DS2 18DS1 18DS2 30GTV1 32DS1 91DS7 91DS5 91DS6 Latitude 37.659 37.659 37.659 37.660 37.660 37.660 37.683 37.683 37.791 38.032 38.092 38.092 38.092 Long -110.872 -110.872 -110.872 -110.874 -110.874 -110.874 -110.883 -110.883 -110.914 -110.980 -111.003 -111.003 -111.003 water depth (m) 2221 2221 2221 2230 2230 2230 2239 2239 2230 2239 2197 2197 2197 material glass glass glass glass glass glass glass glass glass glass glass glass glass

SiO2 62.4 62.0 61.9 62.1 59.9 59.8 59.0 57.9 52.5 57.3 61.8 61.7 61.4

TiO2 1.02 1.06 1.02 1.11 1.30 1.29 1.25 1.46 2.29 1.49 1.14 1.17 1.15

Al2O3 13.5 13.1 13.5 13.5 13.6 13.7 13.6 13.5 13.5 13.5 13.2 13.1 13.1 FeOT 8.10 8.38 7.97 8.51 9.42 9.27 8.86 10.0 13.1 9.75 8.50 8.49 8.51 MnO 0.133 0.134 0.123 0.145 0.158 0.160 0.135 0.157 0.221 0.156 0.157 0.132 0.166 MgO 1.54 1.13 1.57 1.48 1.86 1.77 1.87 2.80 3.96 2.87 1.35 1.35 1.39 CaO 4.44 4.06 4.47 4.52 5.26 5.17 5.09 6.37 8.36 6.46 4.34 4.40 4.42

Na2O 4.47 4.59 4.49 4.01 4.67 4.52 4.63 4.50 3.67 4.24 4.63 4.56 4.64

K2O 1.25 1.36 1.27 1.23 1.10 1.10 1.06 0.850 0.419 0.786 1.21 1.17 1.17

P2O5 0.278 0.341 0.294 0.326 0.403 0.412 0.377 0.387 0.605 0.363 0.405 0.419 0.392

Cr2O3 0.028 0.018 0.028 0.031 0.017

SO2 0.053 0.036 0.060 0.038 0.051 0.046 0.080 0.086 0.218 0.125 0.051 0.056 0.049 Cl 0.918 1.035 0.963 1.01 0.978 0.983 0.868 0.957 0.128 0.552 0.980 0.971 0.960 F

H2O 1.50 0.760 Total 98.13 97.21 97.73 97.97 98.70 98.20 96.86 99.02 99.00 97.68 97.71 97.51 97.39 southern andesitic glasses dacitic glasses cruise SO100 SO100 SO100 SO213 SO213 SO213 SO213 SO157 SO157 SO157 SO157 SO157 SO157 sample 91DS1 92DS1 92DS3 7VSR a 7VSR b 9VSRb 9VSRa 65DS2 63DS3 63DS4 63DS2 63DS1 65DS1 Latitude 38.092 38.156 38.156 38.861 39.030 39.314 39.593 39.505 39.804 39.804 39.804 39.804 39.505 Long -111.003 -111.044 -111.044 -111.191 -111.191 -111.243 -111.243 -111.343 -111.427 -111.427 -111.427 -111.427 -111.343 water depth (m) 2197 2259 2259 2237 2237 2233 2233 2196 2257 2257 2257 2257 2196 material glass glass glass glass glass glass glass glass glass glass glass glass glass

SiO2 60.2 55.7 55.3 60.0 61.0 60.2 59.7 58.3 59.1 57.5 57.3 57.2 68.0

TiO2 1.28 1.51 1.50 1.66 1.60 1.68 1.65 1.75 1.68 1.54 1.71 1.64 0.604

Al2O3 13.5 14.2 14.1 12.9 12.9 12.9 12.9 12.5 12.6 13.3 13.1 13.0 12.4 FeOT 8.94 10.1 10.1 11.8 11.4 11.4 11.2 12.1 11.5 11.3 12.2 11.9 6.14 MnO 0.141 0.170 0.156 0.218 0.183 0.215 0.215 0.218 0.199 0.179 0.223 0.191 0.109 MgO 1.88 3.55 3.50 1.86 1.81 2.03 2.05 2.00 1.79 1.83 1.99 1.99 0.499 CaO 5.12 7.38 7.36 5.45 5.34 5.63 5.65 5.71 5.61 5.52 5.74 5.78 2.67

Na2O 4.35 4.11 4.33 4.27 4.17 4.19 4.20 4.36 4.26 4.49 4.34 4.31 4.70

K2O 1.10 0.715 0.718 0.775 0.805 0.747 0.730 0.711 0.734 0.750 0.693 0.654 1.35

P2O5 0.377 0.411 0.410 0.520 0.486 0.406 0.413 0.565 0.561 0.574 0.778 0.778 0.130

Cr2O3 0.023 0.045 0.027 0.019

SO2 0.052 0.088 0.077 0.096 0.073 0.092 0.095 0.159 0.100 0.103 0.152 0.126 0.025 Cl 0.897 0.811 0.824 0.521 0.549 0.510 0.501 0.530 0.560 0.882 0.643 0.659 0.869 F 0.080 0.065 0.048 0.055

H2O 2.6 1.78 Total 97.82 98.69 98.37 100.01 100.26 99.98 99.20 98.96 98.80 97.99 98.89 98.30 97.47 northern basaltic glasses cruise SO157 SO100 SO100 SO100 SO157 SO157 SO157 SO157 Atalante FH SO100 SO157 SO157 sample 65DS4 106DS2 106DS3 85DS1 2DS1 17DS1 17DS3 17DS2 DR 09-02 102DS1 31GTV5 24DS3 Latitude 39.505 36.907 36.907 37.487 37.666 37.667 37.667 37.667 37.697 37.770 37.776 37.903 Long -111.343 -110.632 -110.632 -110.803 -110.875 -110.877 -110.877 -110.877 -111.150 -110.909 -110.910 -110.947 water depth (m) 2196 2475 2475 2260 2225 2226 2226 2226 2273 2264 2225 2263 material glass glass glass glass glass glass glass glass glass glass glass glass

SiO2 68.0 50.8 50.4 50.1 50.4 50.4 49.9 49.2 50.0 49.4 49.8 50.0

TiO2 0.600 3.55 3.58 1.61 1.74 1.78 1.85 1.89 1.67 1.85 1.58 1.56

Al2O3 12.3 12.3 12.5 14.5 14.1 14.3 13.9 14.2 14.4 14.5 14.4 14.6 FeOT 6.33 16.5 16.5 10.2 11.5 11.2 11.9 11.8 10.8 11.3 10.6 10.5 MnO 0.125 0.255 0.247 0.146 0.214 0.191 0.172 0.200 0.180 0.198 0.211 0.176 MgO 0.504 3.92 3.67 7.43 6.83 6.71 6.36 6.55 7.31 6.93 7.47 7.54 CaO 2.75 8.50 8.10 12.1 11.4 11.4 11.0 11.1 11.8 11.4 12.1 12.1

Na2O 4.57 3.25 3.45 2.71 2.87 2.64 2.87 2.94 2.70 2.81 2.57 2.56

K2O 1.39 0.431 0.453 0.170 0.211 0.194 0.219 0.225 0.150 0.188 0.171 0.159

P2O5 0.137 0.487 0.517 0.250 0.280 0.286 0.279 0.296 0.190 0.277 0.242 0.250

Cr2O3 0.027 0.039 0.048

SO2 0.022 0.227 0.227 0.138 0.228 0.198 0.191 0.241 0.323 0.142 0.183 0.132 Cl 0.844 0.101 0.105 0.006 0.025 0.030 0.033 0.033 0.016 0.025 0.026 0.019 F

H2O 0.450 0.300 Total 97.62 100.30 99.69 99.45 99.90 99.42 98.70 98.66 99.61 99.00 99.42 99.60 southern basaltic glasses cruise SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO157 sample 24DS1 24DS2 29DS6 29DS1 36DS1 36DS2 40DS1 40DS3 40DS2 42DS1 42DS2 42DS3 42DS6 Latitude 37.903 37.903 37.970 37.970 38.215 38.215 38.338 38.338 38.338 38.417 38.417 38.417 38.417 Long -110.947 -110.947 -110.969 -110.969 -111.050 -111.050 -111.073 -111.073 -111.073 -111.082 -111.082 -111.082 -111.082 water depth (m) 2263 2263 2239 2239 2245 2245 2258 2258 2258 2263 2263 2263 2263 material glass glass glass glass glass glass glass glass glass glass glass glass glass

SiO2 49.8 49.6 50.1 49.1 49.4 49.1 50.4 49.9 49.0 50.5 50.4 49.9 49.9

TiO2 1.52 1.56 1.75 1.75 1.64 1.68 1.55 1.57 1.64 1.59 1.64 1.66 1.63

Al2O3 14.7 14.7 14.3 14.3 14.7 14.7 14.7 14.6 14.6 14.4 14.5 14.5 14.5 FeOT 10.2 10.4 11.4 11.0 10.5 10.6 10.3 10.3 10.4 10.6 10.6 10.6 10.5 MnO 0.134 0.200 0.213 0.153 0.168 0.165 0.173 0.185 0.183 0.177 0.181 0.203 0.187 MgO 7.57 7.63 7.08 7.07 7.55 7.50 7.68 7.61 7.68 7.51 7.41 7.41 7.50 CaO 12.0 12.1 11.5 11.3 11.8 11.8 12.2 12.1 11.9 12.1 12.0 12.1 12.1

Na2O 2.56 2.59 2.71 2.72 2.73 2.70 2.67 2.63 2.68 2.69 2.66 2.68 2.69

K2O 0.151 0.168 0.181 0.168 0.151 0.159 0.147 0.148 0.169 0.156 0.151 0.166 0.157

P2O5 0.252 0.249 0.273 0.257 0.266 0.257 0.251 0.253 0.252 0.251 0.261 0.260 0.246

Cr2O3 0.055 0.067 0.053 0.047 0.045 0.072 0.064 0.053

SO2 0.178 0.215 0.202 0.197 0.216 0.224 0.208 0.176 0.218 0.221 0.203 0.138 0.176 Cl 0.017 0.021 0.021 0.020 0.019 0.020 0.013 0.014 0.014 0.017 0.015 0.020 0.017 F

H2O 0.410 0.350 0.320 Total 99.14 99.45 99.75 98.11 99.22 98.98 100.34 99.62 98.77 100.33 100.10 99.57 99.61 cruise SO157 SO157 SO157 SO157 SO157 SO157 SO157 SO213 SO213 SO213 SO213 sample 42DS5 44DS4 44DS3 44DS2 44DS1 45DS1 45DS2 6VSR 1a 6VSR 1 b 8VSRa 8VSRb Latitude 38.417 38.493 38.493 38.493 38.493 38.585 38.585 38.709 38.709 39.004 39.004 Long -111.082 -111.103 -111.103 -111.103 -111.103 -111.131 -111.131 -111.155 -111.155 -111.213 -111.213 water depth (m) 2263 2255 2255 2255 2255 2269 2269 material glass glass glass glass glass glass glass glass glass glass glass

SiO2 49.9 50.2 50.0 49.8 49.7 50.0 49.0 49.6 49.8 50.7 50.8

TiO2 1.72 1.64 1.67 1.77 1.46 1.62 1.64 1.71 1.73 1.78 1.79

Al2O3 14.5 14.4 14.4 14.4 15.2 14.7 14.9 14.2 14.5 14.0 14.2 FeOT 10.7 10.7 10.6 11.0 9.73 10.5 10.5 10.6 10.9 11.5 11.6 MnO 0.205 0.186 0.187 0.185 0.154 0.203 0.195 0.190 0.183 0.188 0.219 MgO 7.28 7.49 7.46 7.19 7.92 7.72 7.71 7.08 7.08 6.91 6.83 CaO 11.8 12.0 12.1 11.4 12.4 12.0 12.0 11.5 11.6 11.4 11.4

Na2O 2.77 2.71 2.69 2.82 2.79 2.71 2.69 2.79 2.76 2.87 2.87

K2O 0.159 0.144 0.150 0.155 0.059 0.138 0.139 0.175 0.173 0.156 0.167

P2O5 0.265 0.256 0.255 0.247 0.213 0.240 0.246 0.174 0.166 0.184 0.176

Cr2O3 0.044 0.056 0.052 0.054

SO2 0.182 0.200 0.133 0.219 0.186 0.210 0.214 0.155 0.162 0.153 0.169 Cl 0.020 0.017 0.018 0.048 0.004 0.014 0.015 0.024 0.025 0.023 0.024 F 0.016 0.023 0.020 0.014

H2O 0.300 0.400 Total 99.42 99.97 99.58 99.19 99.82 100.08 99.36 98.26 99.18 99.98 100.30 Chapter 3: Oxygen isotope evidence for the formation of andesitic-dacitic magmas from the fast-spreading Pacific-Antarctic Rise by assimilation-fractional crystallisation

Supplementary Table 3.1 Trace elements and isotopes northern andesitic glasses cruise SO100 SO157 SO157 SO157 SO157 SO157 SO100 SO157 sample 105DS1 6DS1 6DS2 5DS1 15DS1 3DS1 86DS2 18DS1 (ppm) Li 16.8 14.2 14.2 16.4 18.5 17.3 18.3 Sc 16.2 22.2 17.9 19.8 15.1 10.8 11.8 V 32.7 92.6 202 95.6 68.9 70.4 81.0 Cr 1.07 9.85 29.9 8.76 4.16 6.82 6.36 Co 20.1 19.9 25.7 18.7 14.5 16.4 17.7 Ni 1.48 8.33 18.9 6.62 6.05 6.23 8.16 Cu 16.0 34.0 43.9 35.4 32.9 33.5 36.9 Zn 164 103 86.2 97.7 84.1 86.9 91.7 Ga 26.6 27.0 22.1 25.1 26.5 24.3 25.9 Rb 10.1 10.8 11.5 16.6 22.1 17.4 18.5 Sr 144 134 126 117 107 111 119 Y 97.6 114 74.0 105 109 99 102 Zr 440 569 498 654 714 729 765 Nb 26.8 36.5 23.1 33.7 33.4 33.1 33.7 Cs 0.107 0.109 0.120 0.177 0.231 0.183 0.186 Ba 113 129 111 151 179 152 160 La 25.8 33.4 26.3 37.1 41.3 38.6 39.5 Ce 64.9 83.0 62.2 89.7 96.8 90.3 91.9 Pr 9.98 11.9 8.89 12.3 12.9 12.6 12.9 Nd 48.5 54.7 39.4 54.1 55.2 55.0 56.1 Sm 14.2 15.2 10.7 14.3 14.2 14.4 14.7 Eu 4.11 3.60 2.37 3.00 2.66 2.91 3.02 Gd 17.2 17.3 12.3 15.8 15.4 16.2 16.7 Tb 3.01 3.11 2.15 2.84 2.80 2.85 2.93 Dy 19.5 20.3 14.3 18.4 18.4 18.8 19.3 Ho 4.15 4.31 3.05 3.93 3.95 4.03 4.16 Er 11.8 12.5 8.9 11.5 11.8 11.9 12.3 Tm 1.72 1.85 1.34 1.74 1.80 1.80 1.86 Yb 11.5 12.4 9.0 11.6 12.2 12.1 12.6 Lu 1.68 1.83 1.34 1.70 1.80 1.79 1.87 Hf 10.8 14.8 12.1 16.9 18.9 17.6 18.3 Ta 1.64 2.26 1.42 2.16 2.23 2.02 2.07 Pb 1.88 1.87 1.74 2.25 2.69 2.41 2.39 Th 2.26 3.61 3.32 4.74 6.23 5.34 5.51 U 0.697 1.11 1.04 1.48 1.91 1.66 1.70

87Sr/86Sr 0.702803 0.702878 0.702903 143Nd/144Nd 0.513007 0.513008 0.512997 !18O 5.19 5.08 5.44 5.50 5.43 southern andesitic glasses cruise SO157 SO157 SO100 SO100 SO100 SO100 SO100 SO157 sample 30GTV1 32DS1 91DS7 91DS6 91DS1 92DS1 92DS2 65DS2

Li 11.4 14.9 17.7 17.6 16.2 13.8 13.0 15.4 Sc 37.1 26.0 9.23 9.42 21.9 16.2 24.1 24.0 V 214 169 50.2 43.7 71.6 155 113 84.1 Cr 93.1 28.9 2.42 2.45 13.9 58.1 48.8 7.20 Co 29.7 23.1 14.6 12.9 17.1 22.9 22.7 20.2 Ni 24.8 14.7 3.57 2.89 7.64 22.3 19.6 5.44 Cu 37.7 31.6 25.7 22.8 29.0 29.1 29.3 22.2 Zn 127 96.7 92.0 85.6 106 78.7 94.2 139 Ga 24.9 25.2 25.9 24.6 25.4 22.7 24.5 27.8 Rb 6.60 12.5 17.2 16.2 17.5 10.2 12.3 8.87 Sr 150 124 112 105 111 128 123 117 Y 79.4 94.0 115 109 109 83.9 96.5 125 Zr 304 530 865 812 748 566 610 630 Nb 19.0 24.1 34.6 32.7 29.0 21.7 23.3 18.1 Cs 0.071 0.124 0.168 0.162 0.211 0.095 0.152 0.118 Ba 82.0 117 147 138 147 90.7 109 85.1 La 17.5 28.0 41.1 38.8 36.2 26.5 29.4 25.9 Ce 46.3 68.6 98.4 92.6 89.4 65.2 74.4 71.1 Pr 6.87 9.70 13.9 13.1 12.3 9.52 10.4 10.8 Nd 33.7 43.6 62.1 58.6 54.1 43.4 47.1 52.2 Sm 10.3 12.2 16.5 15.6 14.3 11.9 12.8 15.5 Eu 3.01 2.67 3.29 3.10 2.89 2.67 2.75 3.70 Gd 12.4 13.9 18.8 17.6 15.9 14.0 14.3 18.0 Tb 2.21 2.52 3.33 3.13 2.84 2.46 2.54 3.25 Dy 14.4 16.6 21.9 20.6 18.6 16.2 16.8 21.5 Ho 3.05 3.53 4.73 4.46 3.99 3.48 3.57 4.56 Er 8.7 10.4 14.0 13.1 11.7 10.2 10.3 13.2 Tm 1.28 1.56 2.11 1.98 1.77 1.52 1.57 1.98 Yb 8.4 10.5 14.4 13.4 12.0 10.2 10.6 13.2 Lu 1.23 1.54 2.13 1.98 1.79 1.51 1.56 1.94 Hf 8.0 14.1 21.1 19.8 18.9 13.7 15.3 16.2 Ta 1.26 1.60 2.13 2.00 1.84 1.33 1.49 1.19 Pb 1.30 1.72 2.60 2.32 2.27 1.65 1.76 1.82 Th 1.43 3.48 5.62 5.25 4.74 3.37 3.73 2.38 U 0.480 1.12 1.88 1.74 1.59 1.14 1.26 0.859

87Sr/86Sr 0.702774 0.702834 0.702739 143Nd/144Nd 0.513024 0.513027 0.513097 !18O 5.59 5.51 5.13 dacitic glassnorthern basaltic glasses cruise SO157 SO157 SO100 SO100 SO100 SO157 SO157 Atalante FH sample 63DS1 65DS1 106DS2 106DS3 85DS1 2DS1 17DS1 DR 09-02

Li 19.8 24.8 10.6 10.6 4.99 6.14 5.26 Sc 22.5 9.91 23.1 23.6 25.8 26.6 34.3 V 60.3 20.8 466 471 339 356 320 Cr 3.68 2.40 10.0 15.1 174 78.4 81.8 Co 16.3 5.72 39.2 39.7 37.1 39.0 42.3 Ni 4.57 1.56 14.5 15.8 52.0 39.0 42.8 Cu 26.2 10.7 39.5 40.4 67.5 66.9 74.2 Zn 158 94.1 127 127 67.5 74.8 96.5 Ga 28.8 29.7 21.9 21.7 16.1 17.3 18.9 Rb 7.17 18.0 6.12 6.03 1.82 2.66 3.04 Sr 118 76.5 142 142 144 142 145 Y 152 162 56.3 56.2 24.6 30.0 32.6 Zr 762 850 249 228 93.0 107 116 Nb 19.8 23.2 18.4 18.3 6.24 7.75 7.70 Cs 0.096 0.171 0.063 0.059 0.010 0.022 0.081 Ba 70.4 133 72.3 72.5 36.9 35.9 39.0 La 26.8 39.6 15.8 15.8 6.41 7.12 7.17 Ce 78.4 105 39.1 39.0 16.0 17.9 18.6 Pr 12.5 15.0 5.91 5.92 2.40 2.77 2.79 Nd 62.6 67.9 28.4 28.5 11.8 13.6 13.7 Sm 19.1 19.4 8.24 8.29 3.63 4.17 4.16 Eu 4.58 3.49 2.55 2.55 1.27 1.44 1.43 Gd 22.3 21.8 9.99 10.0 4.55 5.25 5.01 Tb 4.01 4.11 1.73 1.74 0.792 0.921 0.888 Dy 26.3 27.4 11.4 11.5 5.21 6.11 5.94 Ho 5.56 5.94 2.42 2.41 1.10 1.29 1.24 Er 15.9 17.8 6.82 6.85 3.09 3.67 3.48 Tm 2.36 2.73 0.997 1.00 0.447 0.529 0.508 Yb 15.7 18.3 6.66 6.68 2.94 3.56 3.38 Lu 2.29 2.68 0.973 0.979 0.428 0.519 0.503 Hf 18.7 24.7 6.44 6.42 2.71 3.12 3.16 Ta 1.24 1.60 1.18 1.18 0.407 0.505 0.521 Pb 1.94 3.07 1.21 1.25 0.592 0.960 0.701 Th 1.76 4.70 1.39 1.39 0.461 0.595 0.598 U 0.668 1.65 0.434 0.438 0.150 0.187 0.187

87Sr/86Sr 0.702582 0.702684 0.702844 0.702804 0.702805 0.702646 143Nd/144Nd 0.513101 0.513094 0.512987 0.513011 0.513015 0.513059 !18O 5.43 5.56 5.62 southern basaltic glasses cruise SO100 SO157 SO157 SO157 SO157 SO157 SO157 sample 102DS1 24DS1 29DS1 36DS1 40DS1 42DS1 44DS1

Li 6.31 5.37 5.06 5.93 4.70 4.93 5.941 Sc 28.1 29.1 34.549 28.957 36.007 38.702 48.263 V 375 355 324 358 292 306 328 Cr 119 219 125 213 274 227 239 Co 39.9 39.5 42.3 39.5 39.9 40.9 39.4 Ni 52.6 60.5 54.6 64.5 68.7 64.4 57.2 Cu 60.9 65.2 62.6 60.1 68.4 67.5 63.6 Zn 78.8 70.4 93.7 73.2 388.5 87.5 73.9 Ga 18.1 16.9 18.5 17.6 17.5 18.1 17.6 Rb 2.48 2.00 2.36 1.88 1.77 2.02 1.68 Sr 151 144 140 147 135 142 141 Y 31.1 26.6 30.7 29.4 28.0 29.2 28.1 Zr 109 86.0 97.2 96.6 89.4 96.1 96.7 Nb 7.53 5.91 6.02 5.52 4.30 4.95 4.89 Cs 0.021 0.015 0.052 0.014 0.055 0.062 0.018 Ba 35.1 29.3 33.1 26.3 22.9 25.8 24.5 La 6.97 5.52 5.70 5.49 4.57 5.03 4.89 Ce 17.8 14.2 15.1 14.5 12.6 13.8 13.6 Pr 2.80 2.22 2.33 2.37 2.00 2.16 2.14 Nd 13.8 11.3 11.6 12.1 10.4 11.0 10.9 Sm 4.32 3.57 3.73 3.92 3.39 3.53 3.59 Eu 1.50 1.26 1.32 1.39 1.22 1.26 1.29 Gd 5.40 4.54 4.60 5.02 4.23 4.47 4.50 Tb 0.950 0.807 0.811 0.883 0.768 0.790 0.812 Dy 6.31 5.31 5.56 5.92 5.13 5.26 5.37 Ho 1.33 1.15 1.16 1.25 1.08 1.11 1.15 Er 3.80 3.24 3.27 3.59 3.03 3.12 3.24 Tm 0.549 0.466 0.478 0.518 0.437 0.459 0.472 Yb 3.62 3.08 3.16 3.42 2.92 2.99 3.07 Lu 0.540 0.464 0.467 0.508 0.432 0.442 0.452 Hf 3.16 2.61 2.74 2.79 2.53 2.62 2.62 Ta 0.484 0.383 0.410 0.358 0.297 0.333 0.325 Pb 0.594 0.450 0.537 0.467 2.32 0.474 0.435 Th 0.560 0.428 0.445 0.396 0.314 0.358 0.328 U 0.178 0.145 0.147 0.136 0.107 0.123 0.114

87Sr/86Sr 0.702754 0.702678 0.702617 0.702637 0.702606 143Nd/144Nd 0.513021 0.513033 0.513064 0.513101 0.513067 0.513072 !18O 5.51 5.67 5.57 standards standards (for measurement of: (for measurement of: cruise SO157 SO213 SO213 SO157, SO100, Atalante FH) SO213) sample 45DS1 6VSR 1a 8VSRa BHVO (n=2) BIR (n=3) BHVO (n=1) BIR (n=8)

Li 4.971 6.1 6.21 4.82 3.25 4.27 3.22 Sc 35.3 48.8 47.1 34.6 47.5 31.8 46.9 V 293 372 370.9 341 341 333 364 Cr 267 182 98 259 372 299 439 Co 41.1 47.6 48.5 41.7 47.8 46.0 56.6 Ni 73.0 56.2 50.5 107 152 121 182 Cu 64.9 80.8 77.3 134 108 142 132 Zn 87.5 98.8 101 92.1 60.3 111 74.0 Ga 17.8 22.5 15.6 Rb 1.79 2.11 1.79 9.13 0.222 9.21 0.200 Sr 140 161 145 405 107 411 117 Y 32.5 34.1 34.1 23.9 13.9 24.8 15.1 Zr 109 116 112 170 13.5 177 14.9 Nb 4.43 6.09 5.28 17.2 0.516 18.1 0.525 Cs 0.048 0.095 0.006 Ba 22.7 27.6 24.8 138 6.89 137 6.72 La 5.03 5.87 5.30 15.3 0.636 15.8 0.620 Ce 14.4 15.9 14.7 38.2 1.89 39.2 1.93 Pr 2.33 2.44 2.30 5.35 0.369 5.36 0.376 Nd 12.0 13.1 12.4 24.3 2.29 25.7 2.48 Sm 3.92 4.10 3.96 6.15 1.09 6.21 1.12 Eu 1.36 1.49 1.46 2.03 0.500 2.14 0.542 Gd 4.97 5.75 5.64 6.02 1.7 7.00 2.01 Tb 0.883 0.955 0.944 0.942 0.350 0.991 0.373 Dy 6.00 6.39 6.32 5.33 2.51 5.60 2.72 Ho 1.25 1.31 1.31 0.980 0.556 1.00 0.589 Er 3.55 3.89 3.90 2.48 1.63 2.71 1.80 Tm 0.517 0.540 0.543 0.326 0.242 0.335 0.257 Yb 3.45 3.57 3.59 1.99 1.62 2.08 1.72 Lu 0.512 0.535 0.534 0.274 0.240 0.289 0.258 Hf 3.06 3.08 2.97 4.51 0.576 4.68 0.618 Ta 0.308 0.384 0.330 1.14 0.038 1.14 0.042 Pb 0.566 0.475 0.442 2.02 3.04 1.84 2.84 Th 0.320 0.380 0.325 1.26 0.034 1.23 0.029 U 0.115 0.130 0.115 0.430 0.012 0.421 0.010

87Sr/86Sr 0.702578 143Nd/144Nd 0.513085 !18O 5.63 Chapter 3: Oxygen isotope evidence for the formation of andesitic-dacitic magmas from the fast-spreading Pacific-Antarctic Rise by assimilation-fractional crystallisation supplementary Table 2 Partition coefficiants / = not required 0 = not available

Kd (assimilant) AFCKd (parental magma), AFCKd (parental magma) FC cpx plag Amph Mag Ol Opx Cpx Pl Ilmenit Ol Cpx Opx Plag Ilmenit ///// U 20 22 4 29 ///// 7 12 7 7 22 Nb 9 9 18 23 7 7 7 7 9 7 11 7 7 9 La 28 22 18 29 ///// 7 11 7 7 24 Ce 28 22 18 29 ///// 7 11 7 7 24 Pb 9 9 9 9 ///// 2 12 0 0 9 Nd 28 6 18 29 7 7 11 7 24 7 11 7 7 24 Sr 9 6 9 9 7 16 11 1 0 7 11 16 7 0 Sm 14 22 18 29 ///// 7 11 7 7 24 Eu 14 6 18 29 ///// 7 11 7 7 24 Gd 27 6 18 0 ///// 7 10 0 7 0 Yb 14 6 18 29 ///// 7 11 7 7 24 Tb 14 6 19 29 ///// 7 11 7 7 24 Hf 20 6 18 29 ///// 7 11 7 7 24 Cl 13 3 5 0 0 0 13 0 0 0 13 0 3 0 K 14 6 21 25 26 17 11 0 0 26 26 0 26 0

REF 1 Aigner-Torres et al. (2007) 2 Beattie, P. (1994) 3 Bindeman et al. (1998) 4 Brenan et al. (1995) 5 Coogan et al. (2003). 6 Dudas et al. (1971) 7 Dunn, T. & Sen, C. (1994) 8 Elkins et al. (2008) 9 Ewart & Griffin (1994) 10 Hack et al. (1994) 11 Hart & Dunn (1993) 12 Hauri et al. (1994) 13 Hill et al. (2000) 14 Huang et al. (2006) 15 Jenner et al. (1994) 16 Green et al. (2000) 17 Kelemen and Dunn (1992) 18 Klein et al. (1997) 19 Luhr, & Carmichael(1980) 20 Mahood & Hildreth (1983) 21 Nagasawa & Schnetzler (1971) 22 Nash & Crecraft (1985) 23 Nielsen & Beard (2000) 24 Nielsen et al. (1992) 25 Okamoto, K. (1979) 26 Philpotts & Schnetzler (1970) 27 Schnetzler & Philpotts (1970) 28 Sisson (1991) 29 Streck & Grunder (1997) Chapter 4: Constraints on the formation of geochemically variable plagiogranite intrusions in the Troodos Ophiolite, Cyprus

Supplementary Table 4.1

Mineral assemblage and composition of Troodos plagiogranites - thin-section observations (microscopy & EMP) & X-ray diffraction

quartz feldspar amphibole Cpx Tit Cze Ep Chl Mag Tmag Il T Zr Pre U-sp measurements (n): An average: mm textures content textures Zoopigi group some grains show oprical n=19: magnesiohbl, Tro Zoo38A 0.1-0.3 granular n=27: An19-62 zoning actinolitic hbl / <3% /// x / x ? x // Tro Zoo 38B 0.3-1.5 granophyric x, altered x, highly solved /// x / ??? / x // altered, two grain sizes: 0.01 Tro Zoo 39 0.1 & 1-2 granular & 1 mm /// x x / ??? //// Tro Zoo 40 ? ? / / ferro-act.///// <1% ////// optical zoned, some grains highly solved, n=2: Tro Zoo 41 0.5-2.5 granular n=10: An7-31, Or2 highly solved Ferrohbl, hastingitichbl /// x x ??? / x // Tro Zoo 42 1.0-4.0 granular, recristall.selv. altered x, highly solved /// 1.80% / 1.50% /// x // Tro Zoo 43 0.001-3 granular altered solved? ///// ??? //// Tro Zoo 45 0.2-1.5 granophyric zoned x /// 1.50% / 0.50% // x // main group /

n=8: ferrohbl, magnesiohbl, Tro Pal 1 0.5-1.5 granular n=8: An45-60 actinolitic-hbl / x x x x x ?? / x // granophyric/granu n=5: magnesiohbl, Tro Pal 3B 0.5-1 lar n=4: An24-60 actinolitic-hbl, actinolite ///// ??? / x // Tro Pal 4 0.3-0.5 granular altered x, combined with oxides ///// ??? / X // Tro Pal 5A 0.1-0.5 granular altered x ///// ??? / x // Tro Pal 6 0.5 & 0.5-1 granular n=5: An36-56 n=2: magnesiohbl ///// ??? / X // n=15: ferrohbl, magnesiohbl, Tro Pal 7A 0.5 & 0.5-1 granophyric x altered ferroactinolite //// x ??? / x / Tro Pal 9 0.5-1 granophyric altered x ///// ??? / x // n=2: magnesiohbl, Tro Pal 10A 0.2-1 granophyric n=5: An45-60 actinolite ///// ? / x / X / x Tro Pal 10B 0.2-2 granophyric optically zoned, altered x ///// ??? / x // n=11: magnesiohbl, ferro- Tro Pal 11 0.2-1 recystall. Selv. n=1: An21 altered actinolite ///// ??? / x / Tro Pal 13 0.5 granophyric very samll grains x ///// ??? / x // Tro Pal 14B 0.5-1.5 granophyric n=3: An50-55 n=7: ferro-hbl, ferro-act // x x / x ?? ?? x / x Tro Pal 15 0.1-1 recystall. Selv. altered x, combined with oxides // x // ??? / x x / some optical zoned, some n=6: ferrohbl, ferroact, act- Tro Pal 16 0.1-1.5 granophyric n=36: An20-75 partly solved hbl, magnesium-hbl // x x / ??? / x // Tro Pal 17 0.3-1.5 granular altered x, combined with oxides // x // ??? / x // Tro Pal 18 0.1-1 granular optical zoned, partly altered x ///// 1.50% /// x // x,solved, combined with Tro Pal 19A 0.1-1 granular altered oxides // x 1.60% / <0.5% /// x x / Tro Pal 20 0.1-0.5 granular alterd, partly solved x ///// ??? //// x n=11: actinolitic hbl, ferrohbl, magnesiohbl, Cy10.05-05 0.5 - 2.0 granular n=13: An38-60 zoned actinolite, ferroact ///// x x // x // Cy11.05-01 0.1-1 granular altered x /// x / ??? //// Cy27.05-02 <0.1 granular altered x /// x x ??? / x / n=14: actinolite, actinolitic Cy11.05-03 <0.05 granular n=14: An51-65 some zoned, some altered hbl /// x / x / x / x // n=3: An45-76(aplitic huge graines, hornfelsic vein) & n=3: An75- texture with plag, n=3: Cy27.05-03 <0.25 granular 83(Hornfels) partly solved magnesiohbl //// x / x x //// n=22: actinolitic hbl, actinolite, ferroact, ferrohbl, Cy05.05-03 0.5-2.0 granuphyric n=18: An34-66 zoned magnesiohbl ///// x x x / x //

granular, n=18: ferroact, actinolite, Cy10.05-04 1.0-3.0 granopyric n=19: An41-58 altered actinolitic hbl, magnseiohbl ///// x x x //// granular, n=7: actinolitic hbl, Cy12.05-01 0.25-1.0 granopyric n=6: An12-40 altered magnesiohbl, ferrohbl / x / x x x x x //// n03: actinolitic hbl, ferrohbl, Cy14.05-01 0.1-1.5 granular n=4: An24-51 zoned magnesiohbl / x / x x x x ///// granular, n=10: actinolitic hbl, Cy05.05-01 0.5-1.0 granopyric n=6: An41-50 some zoned ferrohbl / x / x / x x x / x // n=22: actinolitic hbl, Cy12.05-02 0.25-1.0 granular n=4: An86-89 altered magnesiohbl, ferrohbl / x / x x x x x / x //

n=4: Woll54-56, granophyric, some optical zoned, partly En25-33, Tro Am 23B 1.0-4.0 recystall. Selv. n=10, An46-71 solved, x,highly solved Fer19-27 //// ??? / x / Tro Fte 46 0.1 fine granular alterd, partly solved x / x /// ??? //// n=10: ferro-hbl, act, ferro- Tro Fte 47 0.1 - 2 granophyric n=18, An14-53 act, act-hbl / x / x / x / x / x / x

Tro Fte 48 0.1-1 granular x x /// x / ??? //// Tro Lem 49 0.3-1 granular x x /// x / ??? //// Tro Ped 50 0.2-1 granular x x /// x x ??? //// granophyric & Tro Am 54 0.3 recrystall.selv. x x ///// ??? / x // granophyric & Tro Am 55 > 3 recystall. Selv. x x ///// ??? //// Tro Am 56 0.5- >3 granophyric optically zoned x ///// ??? //// some optically zoned, partly Tro Pla 58B 0.1-0.6 granular n=13, An49-63 solved n=3: act-hbl magnesio-hbl ///// ???? x // Spilia group /

Tro Spi 24 0.5-2.5 granular up to 2 mm, zoned X / x /// ??? ////

Tro Spi 25 0.1 - 1 granophyric altered /// < 2% < 1% / 1.60% //// x /

Tro Spi 30 < 2 granular altered x, solved? // x x / ??? //// granophyric & Tro Spi 31 < 3 recystall. Selv. altered x, solved? // x x / ??? ////

Tro Spi 32 0.1-0.5 granophyric altered x // x X ////////

Tro Spi 33 ? ? / albit magnesio-hbl / < 1% /// 1.30% ////// n=11: Ferrohbl, act-hbl, granular & n=6: Core An73, rim magnesian-hbl, subcalcic- Tro Spi 35 0.5-2 granophyric An16-42 optical zoned, very fine laths act / x /// ? / x ? /// Tro Spi 36A 0.3-2 granular / zoned, altered n=17: ferro-hbl / <1.5% x <1.5% ///// x //

Tro Spi 36B 0.6-1 granular altered x / x /// ??? / x // abbreviations recrystall. Sel. = recrystallisation selvages Clinopyroxen Cpx x = occured (found by microscopy) titanite Tit ? = opaques found via microscopy clinozoisite Cze % =analysis via x-ray deffraction epidote Ep /= not found Fe-chlorite Chl magnetite Mag titanomagnetite Tmag ilmenite Il zircone T apatite Zr prehnite Pre ulvo-spinell U-sp