Toward an Integrated Model of the Crust in the Icelandic Rift Zones

Dissertation

Presented in Partial Fulfillment of the Requirements for the Degree Doctor of Philosophy in the Graduate School of The Ohio State University

Daniel Francis Kelley, B.S., M.S.

Graduate Program in Geological Sciences

The Ohio State University

2009

Dissertation Committee:

Michael Barton, Advisor

Wendy Panero

Hallan Noltimier

Loren Babcock

Daniel Francis Kelley

2009

Abstract

Iceland lies astride the Mid-Atlantic Ridge and was created by seafloor spreading that began about 55 Ma. The crust is anomalously thick (~ 20-40 km) indicating higher melt productivity in the underlying mantle compared with normal ridge segments due to the presence of a mantle plume or upwelling centered beneath the north western edge of the Vatnajökull ice sheet. Seismic and volcanic activity is concentrated in ~ 50 km wide neovolcanic or rift zones, that mark the subaerial Mid-Atlantic Ridge, and in three flank zones. Geodetic and geophysical studies provide evidence for magma chambers located over a range of depths (1.5-21 km) in the crust, with shallow magma chambers beneath some volcanic centers (Katla, Grimsvotn, Eyjafjallajokull), and both shallow and deep chambers beneath others (e.g., Krafla and Askja). I have compiled analyses of basalt glass with geochemical characteristics indicating crystallization of ol-plag-cpx from 29 volcanic centers in the Western, Northern and Eastern rift zones as well as from the Southern Flank Zone. Pressures of crystallization were calculated for these glasses using a method based on phase equilibrium. Comparison with experimental data indicates that calculated pressures are accurate to ±110 MPa (1σ) and are precise to better than 80 MPa (1σ). The results confirm that Icelandic magmas crystallize over a wide range of pressures (1 to ~1000 MPa), equivalent to depths of 0-35 km. This range partly reflects crystallization of melts en route to the surface, probably in dikes and conduits, after they leave intracrustal chambers. There is reasonably good correlation between the depths of deep chambers (>17 km) and geophysical estimates of crustal thickness suggesting that magma ponds at the crust-mantle boundary. Shallow chambers are located in the upper crust (taken here as <7 km), and probably form at a level of neutral buoyancy. There are ii

also discrete chambers at intermediate depths (~11 km beneath the rift zones), and there is good evidence for cooling and crystallizing magma bodies or pockets throughout the middle (7-15 km) and lower crust (>15 km). It has been shown that glasses in magmas erupted at the Kverkfjöll volcanic system in the Northern Volcanic Zone (NVZ) have compositions that are consistent with partial crystallization at average pressures of 445±69 MPa and 794±92 MPa, corresponding to depths of 15.6±2.4 km and 27.9±3.2 km, and I conclude that magma chambers are located at these depths. These results are consistent with interpretation of recent seismic activity beneath Upptyppingar in the Kverkfjöll volcanic system ~50 km north of the Kverkfjöll central volcano. The earthquake hypocenters are concentrated at depths of 15-18 km with a few occurring at greater depths (~25km), and the seismic activity appears to reflect inflow of magma into the base of the crust (Roberts et al., 2007). Custal thickness, temperatures at the base of the crust, and composition of the crust were used to construct geothermal gradients and profiles of density and seismic velocity through the crust to predict these properties in the lowermost crust. Models of mineralogy change and compositional change were considered. The density at the base of the crust is 3120-3134 kg/m3 giving a crust-mantle density difference of 166-188 kg/m3 assuming a mantle density of 3300 kg/m3. The predicted seismic velocity at the base of the crust is 6.8 km/s. The middle and lower crust in the rift zones is relatively hot and porous. It is suggested that crustal accretion occurs over a range of depths as proposed in recent models for crustal accretion at mid- ocean ridges. The presence of multiple stacked chambers and hot, porous crust suggests that magma evolution is complex and involves polybaric crystallization, magma mixing, and assimilation.

iii

Dedication

Dedicated to all of the very important ladies in my life: Liz, Eve, Marie, and Mom.

Acknowledgements

I thank my advisor Mike Barton for helping me to grow tremendously as a

scientist over the past seven years, for being a good friend, and for putting up with all of

the distractions that I have brought on myself in my personal life. I thank Wendy Panero,

Loren Babcock, and Hal Noltimier for contributing to my intellectual and professional

development. I thank all of the other faculty, staff, and students of the School of Earth

Sciences and formerly the Department of Geological Sciences for making my long stay

here productive and enjoyable. This research was supported in part by the Friends of

Orton Hall Fund of The Ohio State University. Much of the material from Chapter 1 was published by Oxford University Press in Journal of Petrology, volume 49, pp. 465-492.

Vita

1997……………………………Findlay High School, Findlay Ohio

2003……………………………B.S. Geological Sciences, The Ohio State University

2005……………………………M.S. Geological Sciences, The Ohio State University

2002-2003……………...………Undergraduate Teaching Assistant, Department of Geological Sciences, The Ohio State University

2003 to present…………..……..Graduate Teaching Assistant, Department of Geological Sciences, The Ohio State University

2008……………………………Instructor, Denison University

Publications

Kelley, D. F., Barton, M., 2008, Depths of magma chambers in the Icelandic crust, Journal of Petrology, 49(3), 465-492.

Fields of Study

Major Field: Geological Sciences

Table of Contents

Abstract……………………………………………………………………..……………ii

Dedication………………………………………………………………….…………….iv

Acknowledgements……………………………………………………….……………...v

Vita……………………………………………………………………….……………...vi

List of Tables…………………………………………………………….………..…….viii

List of Figures………………………………………………………………..…………..ix

Chapter 1: Introduction…………………………………………………………………..1

Chapter 2: Pressures of Crystallization of Icelandic Magmas…………………………….3

Chapter 3: Petrological Imaging of the Magma Chamber beneath Upptyppingar, Kverkfjöll Volcanic System, Iceland…………………………………………….63

Chapter 4: Density and Seismic Velocity of the Crust in the Icelandic Rift Zones...…....81

Chapter 5: Conclusions…………………………………………………………………126

References………………………………………………………………………………128

Appendix A: Calculation of Pressures………………………………………………….148

Appendix B: Accuracy and Precision…………………………………………………..150

Appendix C: Supplemental Data…………………………………………………….…153

vii

List of Tables

Table 1. Summary of Estimates of Magma Chamber Depths…………………………….5

Table 2. Summary of Calculated Pressures and Temperatures………………………..…26

Table 3. Pressures and Depths of Cotectic Crystallization………………………………41

Table 4. Summary of All Basaltic Glass Data From Upptyppingar and Kverkfjöll….….69

Table 5. Major Element Compositions for Neovolcanic Zone crustal Whole Rocks and Glasses, and Calculated Parent Magmas…………………….……………….….90

Table 6. Abundances of Minerals Calculated Using Perple_X Code…………………..107

Table 7. Mineral Assmblages Used for the Shallow Crust...…….………………….….110

Table 8. Example of Method Used to Calculate Pressure………………………….…..149

Table 9. Supplemental Data…………………………………………………………….153

viii

List of Figures

Figure 1. Location of Important Geologic Features of Iceland…………………….……..9

Figure 2. Positions of Cotectic Lines at Different Pressures……………………………12

Figure 3. Comparison of Predicted Melt Compositions at Different Pressures…………14

Figure 4. Comparison of Whole Rock and Glass Compositions………………….…….17

Figure 5. General Chemical Characteristics of Icelandic Glasses………………………20

Figure 6. Plots of MgO vs. Al2O3, Cao, and CaO/Al2O3………………………………..21

Figure 7. Summary of Results Obtained for All Glasses………………………………..27

Figure 8. Histograms Illustrating Results Obtained for Samples From Individual Localities…………………………………………………………….………..31

Figure 9. Possible Interpretations of Calculated Pressures…………………….………..33

Figure 10. Interpretation of Results for Selected Localities…………………………….36

Figure 11. Interpretation of Results for Glasses From the Hengill Complex…………...39

Figure 12. Histograms Showing Averaged Results for the Pressure of Cotectic Crystallization at Individual Volcanic Centers…...………….……………….45

Figure 13. Depths of Shallow and Deep Magma Chambers Along the Rift Zones……..48

Figure 14. Depths of Intermediate Depth Magma Chambers Along the Rift Zones……49

Figure 15. Schematic Representation of Plumbing Systems Beneath Icelandic Volcanoes……………………………………………………………………..51

Figure 16. Petrologic Model for the Icelandic Crust……………………………………58 ix

Figure 17. Map of Central Iceland………………..…………………………………….65

Figure 18. Field Work at Mt. Upptyppingar………………...…………………………..71

Figure 19. Chemical Variations for Kverkfjöll Glasses………………………...…….…72

Figure 20. Magma Chamber Depths Beneath Kverkfjöll System………………………73

Figure 21. Migration of Earthquake Epicenters and Frequency of Earthquakes………..78

Figure 22. Earthquake Hypocenters Beneath Upptyppingar……………………………80

Figure 23. Variation Diagrams of Whole Rock and Glass Analyses…..………………..88

Figure 24. Geothermal Gradients for Two Crustal Models……………………………..96

Figure 25. Calculated Gravity Anomalies for Icelandic Rift Zone ………..……………99

Figure 26. Density change with depth in the crust with constant composition and basalt mineralogy throughout ………………………………………..………...…..102

Figure 27. P-wave change with depth in the crust with constant composition and basalt mineralogy throughout……………………………………………………....104

Figure 28. One Dimensional Cross Section of Predicted Mineral Assemblages…...….109

Figure 29. Density Variation Through the Crust in Model 2…………………………..112

Figure 30. Variation in Predicted P-Wave Velocity From 0-2400 bars Calculated Using Perple_X for Model 1………………………………………………………..113

Figure 31. Variation in Predicted P-wave Velocity Through the Crust Using the Hacker Calculator for Model 1. ……………………………………………………..115

Figure 32. Seismic Velocity Profiles From This Study Compared with Those of Other Workers……………………………………………………………...………116

Figure 33. Variation in density with depth according to Model 3……………….…….120

Figure 34. Variation in P-wave Velocity with Depth According to Model 3….…..….121

Chapter 1: Introduction

The majority of the Earth’s volcanic activity occurs at the mid-ocean ridges.

These underwater volcanic mountain chains are the locality of crustal accretion at

divergent plate boundaries. The island of Iceland is the only portion of any mid-ocean

ridge that is exposed above sea level. It lies astride the Mid-Atlantic Ridge system,

which runs along center of the entire length of the Atlantic Ocean. Iceland’s exposure

above the sea makes it an important place for geologic study as it is the only place on

Earth where direct study of the crustal accretionary processes of divergent plate

boundaries can be directly studied.

A detailed understanding of the crust in Iceland is necessary to add to the

understanding of these processes. The crust in Iceland is generally speaking, fairly

homogenous. It has all been built from basaltic melts produced along the volcanically

active rift zones that run across the island. It is necessary to describe some key

characteristics in order to build a complete model of the crust. There are many types of

data available, which have been collected through the past several decades of geologic

research on Iceland. These include petrologic, geochemical, and geophysical data. The

goal of the research presented here has been to make a step toward integrating different

types of data to further understanding of the processes at work in the active rift zones of

Iceland.

The compositional, thermal, and physical properties of the crust are intimately

interrelated. They must all be investigated in order understand any one. This integrated

1 approach to the study of the Icelandic crust is potentially beneficial to petrologists, geochemists, geodynamicists, and geophysicists. As Iceland is a volcanically active island nation, the understanding of the processes at work on there are important not only to the local population, but to the public at large. There are climatological implications to the potential for large volume eruptions in Iceland. Also, the thermal regime in Iceland is unique globally making the country a working laboratory for geothermal energy research.

Chapter 2: Pressures of Crystallization of Icelandic Magmas

INTRODUCTION

There is considerable interest in the depths of magma chambers beneath active

volcanoes in Iceland (e.g., Soosalu and Einarsson, 2004; Sturkell et al. 2006) for four

main reasons; First, knowledge of the depths of magma chambers is important for

interpreting precursory activity (seismic, deformation, gas emission, etc.) to volcanic

eruptions, and therefore for forecasting eruptions (e.g., Marti and Folch, 2005). Second,

knowledge of the depths of chambers provides constraints on models for magma evolution, because phase relationships and melt compositions vary as a function of pressure (e.g., O’Hara, 1968; Thy, 1991a; Grove et al. 1992; Yang et al. 1996). Third, knowledge of the distribution of magma bodies is important to understanding thermal gradients, which affect variations in density and seismic velocity in the crust (e.g,. Kelley et al. 2005). Fourth, knowledge of the locations and sizes of magma chambers is essential to understand mechanisms of crustal accretion and differentiation (e.g., Pan and Batiza,

2002, 2003).

Different methods have been used to estimate magma chamber depths in Icelandic crust (Table 1). Most recent studies utilize geodetic techniques (see Sturkell et al. 2006), and yield results that in many cases (e.g., Krafla, Grimsvötn, Katla, possibly Torfajökull) agree with those estimated using geophysical methods (see Table 1), although a magma chamber identified at a depth of 7 km beneath Hengill-Hrómundartindur from geodetic data has not been identified from seismic data (Soosalu & Einarsson, 2004). There are

3 shallow chambers beneath Katla, Grimsvötn and Eyjafjallajökull, deep chambers beneath

Vestmannaeyjar, both shallow and deep chambers beneath Krafla and Askja, and intermediate depth chambers beneath Hekla and Torfajökull. These data, along with evidence for lateral transport of magma along fissures and mixing of magmas from different centers (Sigurdsson and Sparks, 1978, 1981; McGarvie, 1984; Mørk, 1984;

Fagents et al., 2001), suggest extremely complex magma dynamics.

Table 1. Summary of Estimates of Magma Chamber Depth

Center Method Depth (km) Reference Krafla Seismology 3-<7 Einarsson, 1978 Geodesy/Gravity 3 Bjornsson et al., 1979 Geodesy 3 Tryggvason, 1980 Magnetotelluric 3,5-7 Bjornsson et al.., 1985 Geodesy 3, 5-10, >20 Tryggvason, 1986 Geodesy 3 Ewart and Voight, 1991 Seismology 3 Brandsdottir et al., 1997 Geodesy 3 Sigmundsson et al., 1997 Geodesy 3, >5 Arnadottir, et al., 1998 Geodesy/Gravity 2.5 Rymer et al., 1998 Geodesy 2.4,21 de Zeeuw-van Dalfsen et al., 2004 Geodesy/Gravity 2.8, 21 de Zeeuw-van Dalfsen et al., 2006 Depth Range >2.4, 21 Theistareykir Petrology 30 Slater et al., 1997 Petrology 10-31 McClennan et al., 2001 Depth Range 10-31 Bláfjall Petrology 7-10, 14 Schiellerup 1995 Petrology 2-7, 14-26 Schiellerup 1995* Depth Range >2, 26 Askja Geodesy 1.5-3.5 Tryggvason, 1989 Geodesy 1.5-3.5 Rymer & Tryggvasson, 1993 Geodesy 1.5-3.5 Sturkell & Sigmundsson, 2000 Geodesy/Gravity 3,16 de Zeeuw-van Dalfsen et al., 2005 Geodesy 3, 20 Pagli et al., 2006 Geodesy 3,16 Sturkell et al., 2006 Depth Range >1.5, 20 Grimsvötn Gravity/magnetics 1.5-4 Gudmundsson & Milsom, 1997 Geodesy >1.6 Sturkell et al., 2003 Teleseismic P-wave delays >4 Alfaro et al., 2007 Depth Range >1.5, 4 Kistufell Petrology >35 Breddam, 2002 Torfajökull Magnetotelluric 10-15, 15-25 Eysteinsson & Hermance, 1985 Magnetotelluric, Geotherm >3 Gudmundsson, 1988 Seismology 8 Soosalu & Einarsson, 1997 Petrology >7 Gunnarsson et al. 1998 Seismology 8, >15 Soosalu & Einarsson, 2004 Seismology <14 Soosalu et al, 2006a Geodesy 8 Sturkell et al., 2006 Range >3, 25 Hekla Geodesy 8 Kjartansson & Gronvold, 1983 Magnetotelluric 8 Eysteinsson & Hermance, 1985 Geodesy 9 Sigmundsson et al., 1992 Geodesy 6.5 Linde et al., 1993 Geodesy 5.0-6.0 Tryggvason, 1994 Geodesy >6.5 Jónsson et al. 2003 Seismic ≥14 Soosalu & Einarsson, 2004 Geodesy 11 Sturkell et al., 2006 Range >5, >14 Katla Seismology 3 Gudmundsson et al..1994 Geodesy 4.7 Sturkell et al., 2003 Seismology 2 Soosalu et al. 2006b Range >2, 4.7 Eyjafjallajökull Geodesy 3.5 Sturkell et al., 2003 Geodesy 6.3 Pedersen and Sigmundsson, 2004, 2006 Range >3.5, 6.3 Vestmannaeyjar Seismology 15-25 Einarsson & Bjornsson, 1979 Seismology 10-15 Gebrande et al., 1980 Petrology 18-28, 28-35 Furman et al., 1991 Petrology 10-17 Thy, 1991 Range >10, 35 Hengill-Hrómundartindur Petrology 8-11 Hansteen, 1991 Geodesy 7 Sigmundsson et al., 1997 Geodesy 7 Feigl et al., 2000 Króksfjördur Petrology 5.5 Jónasson et al, 1992 Thingmuli Petrology 3.5 Frost and Lindsley, 1992 Snaefell Petrology 13 Hards et al., 2000 Austerhorn Petrology 2 Furman et al., 1992 * Estimated from Fig. 4 in Schiellerup (1995). 5

Geodetic and geophysical methods are most useful for locating chambers beneath

active volcanoes, whereas petrologic methods allow the depths of magma chambers

beneath both active and inactive volcanic centers to be determined. Various petrologic techniques are used to determine the pressure, and hence depth, of crystallization (Table

1), but some of the results reported for Icelandic magmas (Table 1) are only qualitative.

Nevertheless, petrologic estimates agree well with those obtained using other methods

when direct comparison is possible (e.g., Hengill, Torfajökull, and Vestmannaeyjar).

In this paper I report new estimates of the pressures of crystallization of Icelandic

magmas and use these to determine the depths of magma chambers. Pressures of

crystallization are determined from experimentally-established phase equilibrium

constraints using the method described by Yang et al. (1996). The results differ from

those obtained in previous studies in two respects. First, quantitative pressure estimates

were obtained using the compositions of glasses, which unambiguously represent pre-

eruptive liquid compositions. Second, pressures were determined using glasses from 29

localities so that the results are applicable to a wide geographic area, and can be

interpreted in terms of crustal thickness. We discuss the implications of the results for the

structure and accretion of the crust, for geothermal gradients, and for magma evolution.

GEOLOGIC BACKGROUND

Iceland lies astride the Mid-Atlantic Ridge (MAR) and is characterized by crust that

is ~ 20-40 km thick (Bjarnason et al. 1993; Darbyshire et al. 1998, 2000a; Menke et al.

1998; Kaban et al. 2003; Foulger et al. 2003; Leftwich et al. 2005) compared with the

7.1±0.9 km thick normal MAR crust (White et al. 1992; Bown and White, 1994). The anomalously thick crust indicates higher melt productivity in the underlying mantle compared with normal ridge segments (White and McKenzie, 1995). The higher melt productivity, together with geochemical differences between Icelandic basalts and normal mid-ocean ridge basalts (N-MORB) (e.g., Schilling, 1973), provides evidence for a mantle plume or upwelling centered beneath the northwestern edge of the Vatnajökull ice sheet (Breddam et al. 2000; Darbyshire et al., 2000a Leftwich et al., 2005)(Fig. 1).

Alternative hypotheses to account for high melt productivity in the sub-Icelandic mantle were discussed by Foulger and Anderson (2005) and Foulger et al. (2005).

The full spreading rate for Iceland is 18-20 mm yr-1 (LaFemina et al. 2005). The subaerial ridge forms ~ 50 km wide neovolcanic or rift zones characterized by abundant seismic and volcanic activity (Fig. 1). The Western Volcanic Zone (WVZ) and the

Northern Volcanic Zone (NVZ) formed at ~7 Ma following eastward relocation of the spreading axis from the Snaefellsnes area (Hardarson et al. 1997). The WVZ can be traced across the Reykjanes Peninsula extensional leaky transform to the Reykjanes

Ridge (RR), whereas the NVZ is offset from the Kolbeinsey Ridge (KR) by the ~120 km long right-lateral Tjörnes Transform Fault. Propagation of the NVZ to the southwest at

~3 Ma (Steinthorsson et al. 1985) formed the Eastern Volcanic Zone (EVZ), so that spreading is partitioned between the WVZ and EVZ in southern Iceland. These rift zones are connected by the complex South Iceland Seismic Zone (SISZ), which accommodates

left-lateral transform motion, and by the Mid-Iceland Belt (MIB), which may be a leaky

transform (Oskarsson et al. 1985; LaFemina et al. 2005) or a non-transform relay zone

(Sinton et al. 2005).

The neovolcanic zones are mostly covered by basalt flows younger than 0.8 Ma.

About 30 en echelon volcanic systems have been recognized that mostly consist of central

shield or composite volcanoes transected by fissure swarms. Eruptions from both

volcanoes and fissure systems occur in the WVZ and NVZ, but central volcanoes are lacking in the EVZ. Many central volcanoes have calderas with typical dimensions of 3 x

4 km, whereas fissure swarms are 5-20 km wide and 40-150 km long, and are often characterized by crater rows formed by scoria and spatter. Other volcanic features include small lava shields, tuff rings and maars, as well as hyaloclastite ridges, hyaloclastite cones and table mountains (tuyas) produced by subglacial and subaqueous eruptions.

Olivine tholeiites and tholeiites dominate the volcanic products, but intermediate and acid volcanics are also erupted, especially at central volcanoes.

Figure 1. Location of important geologic features of Iceland. RR – Reykjanes Ridge; KR – Kolbeinsey Ridge; NVZ – Northern Volcanic Zone; EVZ – Eastern volcanic Zone; WVZ – Western Volcanic Zone; WFZ – Western Flank Zone; SVZ – Southern Flank Zone; EFZ – Eastern Flank Zone; RP – Reykjanes Peninsula; TFZ – Tjornes Fracture Zone, SISZ – South Iceland Seismic Zone, CIB – Central Iceland Belt; SP – Snaefjellsnes Peninsula; V – Vatnajokull Glacier; L – Langjokull Glacier, and sample localities; Th – Theistareykir, Hb – Herdubreid, Bu – Burfell, Bf - Bláfjall Ridge, Ha – Halar, Se – Seljahjalli, As – Askja, Hr – Hrimalda, Gi – Gigoldur, Sp – Sprengisandur, Ki – Kistufell, Kv – Kverkfjoll, Ba – Bardabunga, Gr – Grimsvötn, Vd – Veidivötn, Lk – Laki., La – Langjokull, Hl – Hlodufell, Ra – Raudafell, E – Efstadalsfjall, Kf – Kalfstindar, Tg – Thingvellir, Mi – Midfell, Ma – Maelifell, He – Hengill, Ka – Katla, Hk – Hekla, Ge – Geitafell. Note that samples from Sprengisandur were collected to the NW of the Tungnafellsjokull volcanic system (Meyer et al. 1985), but are thought to originate from this center. In addition, pressures have been calculated for two samples from unspecified localities in the NVZ (NE) and the Reykjanes Peninsula (Rk) (Meyer et al. 1985).

Volcanic activity also occurs in three off-rift flank zones (Fig. 1), the Western

Flank Zone (WFZ), the Southern Flank Zone (SFZ), and the Eastern Flank Zone (EFZ).

Eruptions occur at large shield or composite volcanoes with calderas, and at small lava

shields, tuff rings and scoria cones. However, the extensive fissure swarms characteristic

of the rift zones are absent, and eruptions produce transitional and alkaline basalts along

with intermediate and silicic compositions (Meyer et al. 1985; Oskarsson et al. 1985).

Volcanic products along the NVZ, EVZ and SFZ change from tholeiitic basalts in

northern and central Iceland to Fe-Ti rich transitional basalts in southern Iceland and

alkali basalts in Vestmannaeyjar.

METHODS AND DATA

Method for Determining Pressure of Crystallization

Various petrologic techniques can be used to estimate the pressure (P) and hence

depth (z) of crystallization of magmas (Table 1). The most appropriate method for use

with a large number of samples is based on comparing the compositions of erupted melts

with those of liquids lying along P-dependent phase boundaries. Many basalt magmas

crystallize olivine (ol), plagioclase (plag), and clinopyroxene (cpx), and their

compositions can be compared with those of liquids lying along the ol-plag-cpx cotectic

boundary. The effect of pressure on the latter has been determined experimentally, (e.g.,

O’Hara, 1968; Grove et al. 1992), and can be seen by recasting melt compositions into

normative mineral components and projecting phase relations onto pseudoternary planes

10 in the system CaO-MgO-Al2O3-SiO2. Projection of phase relationships from plag onto the plane ol-cpx-qtz using the recalculation procedure of Walker et al. (1979) clearly shows the shift of the ol-plag-cpx cotectic towards ol with increasing P (Fig. 2).

Crystallization pressure can be estimated by comparing the projected compositions of natural samples with the locations of cotectics on such diagrams, and this method has been used to estimate crystallization pressures for Hengill by Trønnes (1990), for Bláfjall

Table Mountain by Schiellerup (1995), for Kistufell by Breddam (2000), and for

Theistareykir by Maclennan et al. (2001a).

Figure 2. Position of the ol-plag-cpx cotectic at different pressures projected from plagioclase onto the pseudoternary plane Ol-Cpx-Qtz using the method described by Walker et al. (1979). Pressures of cotectic given in GPa. Locations of cotectics based on experimental data from Walker et al. (1979), Yang et al. (1996), Baker and Eggler. (1987), Spulber and Rutherford (1983), Grove and Bryan (1983), Juster et al. (1989), Tormey et al. (1987), Kinzler and Grove (1992), Thy and Lofgren (1994), Sack et al. (1987), Bender et al. (1978), and Shi (1993).

The shift of the ol-plag-cpx cotectic towards ol and plag (see O’Hara, 1968;

Grove et al. 1992) reflects the different pressure dependencies of cpx-liq, ol-liq and plag- liq equilibria, with higher pressure favoring earlier crystallization of cpx. This results in the development of a trend of decreasing CaO with decreasing MgO in liquids at an earlier stage of crystallization. Weaver and Langmuir (1990), Langmuir et al. (1992),

Danyushevsky et al. (1996), Yang et al. (1996), and Herzberg (2004) have proposed models to quantitatively estimate the crystallization pressure based on such relationships.

These models are all calibrated with experimental data, and I have elected to use that of

Yang et al. (1996), who presented equations that describe the composition of liquids along the ol-plag-cpx cotectic as a function of P and T. Hence, the composition of the liquid is used to predict the P (and T) of saturation with ol, plag and cpx, as illustrated in

Fig. 3 using two glass analyses from Bláfjall Table Mountain (Schiellerup, 1995). A series of liquid compositions that lie on the ol-plag-cpx cotectic have been calculated from each glass analysis at increments of 100 MPa. These predicted liquid compositions have been converted to normative mineral components assuming that ΣFe=FeO and projected from plag onto the plane ol-cpx-qtz using the procedure of Tormey et al. (1987) as modified by Grove et al. (1993). Comparison of observed glass compositions and predicted liquid compositions indicates crystallization at ~100 MPa for sample 1.8.5 and

~670 MPa for sample 1.8.1. These pressures agree with those estimated by Schiellerup

(1995) for crystallization of units BII and BI at this locality.

Figure 3. Comparison of predicted melt compositions saturated with ol-plag-cpx at pressures between 0.001 and 1 GPa and observed glass composition for two samples (1- 8-1 and 1-8-5) from Bláfjall Table Mountain (Schiellerup, 1995). Open circles - predicted melt compositions at increments of 0.1 GPa for each sample. Filled circle – analyzed composition of glass in sample 1-8-1. Filled square - analyzed composition of glass in sample 1-8-5. Melt compositions have been converted to normative mineral components and projected from plag onto the pseudoternary plane ol-cpx-qtz using the procedure of Tormey et al. (1987). Comparison of predicted and observed melt compositions indicates crystallization at ~0.67 GPa for sample 1-8-1 and ~0.1 GPa for sample 1.8.5.

Rather than use graphical methods, I have calculated pressures using the

procedure described in Appendix 1. An assessment of the accuracy (~110 MPa, 1σ) and precision (~80 MPa, 1σ) of the calculated pressures is given in Appendix 2.

Herzberg (2004) noted that pressures obtained with his method can differ significantly (up to ~300 MPa.) from those obtained using the method of Yang et al.

(1996). The methods were calibrated with different sets of experimental data, but even so the reason for this discrepancy is not clear. The method used in this paper yields reliable results for the following reasons: 1, Yang et al. (1996) obtained results for MORB and

Hawaiian samples that are consistent with those obtained by other methods; 2, Pressures estimated by Maclennan et al. (2001a) for basalts from Krafla and Theistareykir using the

Yang et al. (op cit.) method agree with those obtained from clinopyroxene geobarometry;

3, Pressures calculated for samples from Midfell in this work agree to better than ±60

MPa with results obtained by Gurenko and Sobolev (2006) using the method of

Danyushevsky et al. (1996), and to better than ±80 MPa with results obtained by these workers using clinopyroxene geobarometry; 4, my results yield estimates of the depths of magma chambers that are consistent with those obtained using geodetic, geophysical, or other petrologic methods. In addition, Herzberg’s (2004) method yields negative and therefore unrealistic pressures for ~ 22% of the samples used to constrain the depth of magma crystallization in this paper. Nevertheless, it is apparent that more experimental data are needed to refine petrologic methods used for the geobarometry of magmas. In particular, there is need for additional experiments to more closely establish the

composition of liquids along the ol-plag-cpx cotectics for a range of basalt compositions

over the pressure range 100-1000 MPa. The implications of using Herzberg’s method

(2004) to calculate crystallization pressures for Icelandic magmas are briefly discussed in

the section “Interpretation of Pressure.”

Samples

Volcanic activity on Iceland produces about ~0.12 km3 yr-1 of fresh lava,

hyaloclastite, scoria, and spatter (see photo in Figure 18). Many recently-erupted samples

contain glass; especially those erupted in subglacial and subaqueous environments (see

photo in Figure 18). Glass analyses are preferable to whole-rock analyses for calculating

the crystallization pressure, because glasses represent samples of quenched melts.

Therefore, glasses formed from liquids in equilibrium with ol, plag, and cpx should have

compositions that lie exactly on the cotectic at the pressure of crystallization. Some whole rock samples represent melts, but others represent mixtures of crystals and melt. It cannot be assumed that the latter formed by closed system crystallization, because the crystals may be of accumulative origin or may represent xenocrysts (eg. Trønnes, 1990;

Hansteen, 1991; Révillon et al. 1999; Hansen and Grönvold, 2000). In such cases, the whole-rock samples do not represent melts and this explains why many whole-rock compositions are displaced from glass compositions in pseudoternary projections (Fig. 4).

The erroneous assumption that whole-rock samples represent melts can lead to large errors in pressure estimates (up to 1000 MPa for the examples shown in Fig. 4).

I have compiled over 800 published glass analyses from localities listed in the

shown in Fig. 1. Most glasses are from localities in the NVZ, WVZ, and EVZ, but 44 are

from Hekla and Katla in the SFZ.

Figure 4. Comparison of whole-rock and glass analyses for samples from Herdubreid, Geitafell, and Hlodufell in projections from plag onto the pseudoternary plane ol-cpx-qtz using the procedure of Tormey et al. (1987). Filled circles – glass analyses. Open Circles – whole-rock analyses. All analyses from Moore and Calk (1991). Fig. 4a. Samples 336, 338, 347 and 350 from Herdubreid. The whole-rock samples have compositions very similar to glasses from the same sample and may represent liquids that crystallized at similar pressure to the liquids represented by the glasses. Fig. 4b. Samples 85T37 and 85T31 from Geitafell. The compositions of the whole-rock samples differ from those of glasses in the same sample and do not represent liquids that crystallized at the same

17 pressure as the liquids represented by the glasses. Fig. 4c. Samples 85T3 and 85T12 from Hlodufell. The whole-rock composition of 85T3 is different from the glass in the same sample and does not represent a liquid that crystallized at the same pressure as the liquid represented by the glass. The whole-rock composition of 85T12 is similar to the glass in the same sample may represent a liquid that crystallized at about the same pressure as the liquid represented by the glass.

I also compiled 201 analyses of glasses in melt inclusions. The compositions of melt inclusions can be modified by post-entrapment crystallization and diffusive re- equilibration with the host mineral (e.g., Danyushevsky et al. 2002). Therefore, I used only those inclusion glasses with compositions that plot along and within arrays defined by groundmass glasses on variation diagrams. Using this criterion, 64 glass inclusion analyses were selected and used to supplement data for groundmass glasses.

Compositional data for samples from each locality are summarized in Table 1 of the Supplemental Data in Appendix C. Detailed discussion of chemical variations shown by the glasses is beyond the scope of this paper, and we describe only the general compositional characteristics are defined along with evidence that these characteristics are consistent with crystallization of ol, plag, and cpx.

Based on their SiO2 to Na2O+K2O ratios 584 out of 588 samples can be classified as basalts (Fig. 5a). The remaining samples are basaltic andesites (2 from Hengill and 2 from Askja). In addition, most glasses (553) are tholeiitic according to the criterion proposed by MacDonald (1968) for Hawaiian lavas (Fig. 5b), and range in composition from olivine tholeiites (normative Ol up to 18 wt. %) to quartz tholeiites (normative Q up

to 8 wt. %). The remaining 35 glasses plot in the alkaline field but only two of these

(from Hekla) contain normative nepheline, and I therefore consider all of these samples to represent transitional basalts. CIPW norms were calculated with Fe2O3 and FeO

contents fixed assuming logfO2=FMQ-1 (McCann and Barton, 2004) (abbreviations: Ol –

Olivine; Q – Quartz). All glasses show strong enrichment in FeO as MgO decreases (Fig.

5c).

Figure 5. General chemical characteristics of Icelandic glasses. Fig. 5a. Plot of total alkalis versus SiO2 with boundaries between different magma types from LeBas et al. (1986). B – Basalt. BA – Basaltic Andesite; TB – Trachybasalt. BTA – Basaltic Trachyandesite. Fig. 5b. Plot of total alkalis versus SiO2 showing boundary between sub- alkaline (SA) or tholeiitic and alkaline (A) compositions proposed by MacDonald (1968) for Hawaiian lavas. Fig. 5c. Plot of MgO versus FeO* (Total Fe as FeO) illustrating the strong iron-enrichment trend developed during differentiation.

Figure 6. Plots of MgO versus Al2O3, CaO, and CaO/ Al2O3 illustrating chemical variations produced by crystallization. Variations of Al2O3, CaO, and CaO/Al2O3 with MgO allow identification of the mineral phases that crystallized during magma evolution (Fig. 7). The decrease in Al2O3 with decreasing MgO (Fig. 7a) is consistent with crystallization of ol+ plag±spinel, and many Icelandic basalts contain phenocrysts or microphenocrysts of these minerals (eg. Meyer et al. 1985). However, the strong decrease in CaO and slight decrease in CaO/Al2O3 with decreasing MgO (Figs. 7b and c) requires crystallization of cpx (see also Michael and Cornell, 1998; Herzberg, 2004).

Chemical variations shown by the glasses can be qualitatively explained by

crystallization of ol-plag-cpx (±spinel) (Fig. 7). Quantitative mass-balance models

21 confirm that removal of these three phases accounts for major-oxide variations at many localities (e.g., Meyer et al. 1985; Schiellerup, 1995; Maclennan et al. 2001a, 2003), but this does not necessarily mean that all glasses represent liquids lying along ol-plag-cpx cotectics. The scatter shown on variation diagrams (Fig. 5c and Fig. 7) indicates that the basalts do not evolve along a single liquid line of descent (LLD). This suggests crystallization along ol-plag-cpx cotectics at different pressures (polybaric crystallization), but other explanations are also possible as described in a later section of this paper (see also Herzberg, 2004). This requires that results for individual localities are evaluated in the context of geochemical and petrographic data. Some samples (24) from the Hengill-Hrómundartindur volcanic complex in the WVZ, as well as from Hrimalda,

Gigoldur, and Sprengisandur in the NVZ have anomalously high CaO contents due to clinopyroxene assimilation (CaO/Al2O3 >1; see Fig. 6) as discussed by Trønnes (1990).

These compositions lie outside the range of those used by Yang et al. (1996) to calibrate the method for pressure calculation. Therefore, I do not report pressures for glasses with anomalously high CaO/Al2O3.

RESULTS

Equilibration Pressures

Pressures were calculated for all glasses, but some results are unrealistic and others are considered unreliable. The average uncertainty in P calculated for nearly all samples is <120 MPa, and is similar to that estimated from experimental data (Appendix

2). In contrast, uncertainties calculated for 24 of the samples are substantially greater (up to 180 MPa). I consider these pressures to be unreliable, and exclude most of them from further consideration. This has no effect on the conclusions, because similar but more reliable estimates of P are obtained for other samples from the same localities. I made an exception for results obtained for three samples from Kverkjoll (Hoskuldsson et al.

2006). Pressures calculated for these samples (790-860±160 MPa) are much higher than those obtained for other samples from this locality (≤490 MPa), and may provide insight into the evolution of magmas at this volcanic center.

Pressures calculated for 55 samples lie between 0 and -100 MPa and can be interpreted as indicating crystallization at ~0.1 MPa if uncertainties are taken into account, and indeed positive pressures close to 0.1 MPa are obtained for other samples from these localities. Nevertheless, I exclude all samples that yield negative pressures from further consideration.

The results demonstrate that Icelandic magmas crystallize over a wide range of P, from ~0.1 MPa to ~1000 MPa (Table 2), indicating that magmas evolve over a range of depths in the crust and, perhaps, upper mantle. Few pressures exceed 600 MPa, and most magmas seem to have crystallized at P~300-400 MPa (Fig. 7a). It is important to note that results obtained using Herzberg’s (2004) method also indicate that Icelandic magmas crystallize over a wide range of P, from ~0.1 MPa to ~750 MPa, although most magmas probably crystallized at ~200 MPa rather than the 300-400 MPa as suggested by my work.

There is no correlation between P and MgO (Fig. 7b), and this is unusual in light

of the positive correlation between P and MgO demonstrated for MORB glasses by

Michael and Cornell (1998). Again, it is important to note that there is also no correlation

between MgO and values of P calculated using Herzberg’s (2004) method. This indicates

that the lack of correlation between P and MgO is not an artifact of my method of

calculation. However, for glasses from most individual localities there is a positive

correlation between P and MgO, whereas for glasses from the Hengill complex there is a

negative correlation between P and MgO. The unusual negative correlation for the

Hengill complex samples is attributed to the effects of crustal assimilation (see later

discussion). The apparent lack of correlation between P and MgO in Fig. 7b is therefore

the consequence of plotting results for a relatively large number of samples (120) from

the Hengill complex together with those for samples from other localities on the same

diagram.

Magma temperatures were calculated using the method of Yang et al. (1996)

(Table 2). Yang et al (1996) noted that their geothermometer reproduces temperatures to

better than 20OC for most samples. The range of temperatures obtained for Iceland

glasses is 1232-1134OC, and is similar to that shown by MORB glasses. A similar range

in T (1254-1099OC) is obtained using the olivine-melt geothermometer of Sugawara

(2000). As expected, there are positive correlations between T and P (Fig 8c) and, with

the exception of Hengill (see later discussion), between T and MgO

Results for individual localities

Representative results obtained for samples from four localities are shown in Fig.

8. Results obtained for a few localities suggest magma evolution at one pressure, for example, 140±80 MPa at Gigoldur and 750±40 MPa at Kalfstindar (Fig. 8a). Results for other localities such as Bláfjall Table Mountain (Fig. 8b) indicate magma evolution at two distinct pressures (590±70 MPa and 100±50 MPa), whereas results for most localities indicate magma evolution over a relatively wide range of pressure rather than at one or two distinct pressures. This is illustrated by results from Laki (380±100 MPa) and

Hengill (260±170 MPa) shown in Fig. 8c and 9d. The results for Hengill are unusual in that many calculated pressures are close to 0.1 MPa (negative pressures were calculated for ~36% of glasses from this locality) and, as described above, there is a negative correlation between P and MgO.

Table 2. Summary of Calculated Pressures and Temperatures

Zone Locality n Ave P Max Min Ave T Max Min NVZ Theistareykir 17 320.1 555.1 131.4 1209.9 1218.2 1200.0 Herdubreid 22 395.7 657.4 173.4 1191.4 1212.9 1173.5 Burfell 3 309.8 324.0 294.9 1174.7 1187.1 1150.1 Bláfjall Ridge 19 329.2 695.0 17.2 1193.8 1215.2 1172.8 Halar 5 511.1 583.4 420.6 1204.9 1212.9 1197.0 Seljahjalli 4 775.6 949.4 701.9 1218.7 1232.1 1212.3 Askja 8 272.8 542.4 124.8 1160.2 1180.7 1134.4 Hrimalda 2 176.0 240.0 112.0 1198.2 1202.2 1194.2 Gigoldur 7 170.4 227.8 81.4 1196.5 1202.7 1190.3 Sprengisandur 7 357.2 592.7 53.8 1180.2 1204.6 1164.7 Kistufell 10 401.4 518.2 143.2 1208.6 1214.5 1196.6 Kverkfjoll 6 507.7 859.4 89.5 1177.8 1200.7 1149.4 Northeastern1 1 331.3 - - 1199.7 - - EVZ Bardabunga 17 270.7 841.7 64.9 1175.3 1209.8 1164.0 Grimsvötn 11 308.1 512.3 97.7 1175.1 1192.2 1149.2 Bardabunga-Grimsvotn 28 284.2 841.7 64.9 1175.1 1209.8 1149.2 Veidivötn 24 209.9 622.5 35.4 1171.5 1194.8 1150.0 Laki 67 379.3 765.5 144.6 1171.4 1192.5 1161.5 SFZ Katla 17 426.7 656.3 199.3 1183.8 1212.4 1159.8 Hekla 17 626.2 1032.5 425.1 1208.6 1227.5 1166.8 Hekla-Katla 34 526.5 1032.5 199.3 1196.2 1227.5 1159.8 WVZ Langjokull 3 253.4 651.5 34.7 1193.4 1216.8 1178.7 Hlodufell 19 356.6 523.4 168.2 1183.5 1198.5 1170.2 Raudafell 12 516.7 721.6 312.9 1199.7 1210.6 1183.9 Efstadalsfjall 30 476.9 813.4 198.1 1198.1 1222.2 1180.5 Kalfstindar 10 745.0 771.1 650.9 1217.1 1218.9 1210.9 Thingvellir 10 374.6 627.4 220.6 1194.0 1205.7 1181.9 Midfell 29 99.8 369.4 0.4 1186.9 1201.4 1176.8 Maelifell 1 14.0 - - 1189.2 - - Hengill 90 314.8 753.9 17.5 1182.4 1217.6 1142.2 Hengill Complex 120 260.3 753.9 0.4 1183.6 1217.6 1142.2 RP Geitafell 12 399.9 536.0 281.4 1185.1 1197.7 1178.4 Reykjanes1 1 201.7 - - 1184.6 - - All Data - 462 349.6 1032.5 0.4 1186.0 1232.1 1134.4 1Samples from Meyer et al. (1985) Average, maximum, and minimum values based on results obtained for all samples from each locality

Figure 7. Summary of results obtained for all glasses excluding those considered unrealistic or unreliable (see text for discussion). Fig. 7a. Histogram indicating range of calculated pressures. Fig. 7b. Plot of calculated pressure in GPa versus MgO. Fig. 7c. Plot of T (OC) calculated using the method of Yang et al. (1996) versus P (GPa).

DISCUSSION

Interpretation of pressure

The calculated pressures reflect the pressure of crystallization only for glasses that

represent liquids lying along ol-plag-cpx cotectics. Glass analyses from some localities

show compositional variations (e.g., increasing CaO and CaO/Al2O3 with decreasing

MgO) that are consistent with crystallization of ol-plag rather than of ol-plag-cpx.

Glasses from other localities have compositions consistent with ol-plag-cpx

crystallization but occur in samples that lack phenocrysts or microphenocrysts of cpx. In

other words, petrographic data provide no evidence for ol-plag-cpx crystallization for these glasses. This is the “pyroxene paradox” recognized for many MORB and other lava suites (e.g., Dungan and Rhodes 1978; Fisk et al. 1982; Grove et al. 1992; Elthon et al.

1995). There are several possible explanations for this paradox, including: cotectic crystallization of ol-plag-cpx followed by crystallization of ol and plag (with dissolution of cpx) during ascent; magma mixing; and crystallization accompanied by assimilation of gabbroic crust (Meyer et al. 1985; Trønnes, 1990; Hansteen, 1991; Schiellerup, 1995;

Hansen and Gronvöld, 2000; Breddam, 2002; Maclennan et al. 2003; Gurenko and

Sobolev, 2006). In principle, the results for each locality can be filtered to identify those melts in equilibrium with ol±plag±spinel. Most of these provide no constraints on crystallization pressure, but some can be used to place limits on the pressure(s) of ol-

28 plag-cpx cotectic crystallization as illustrated in Fig. 9. (see also Michael and Cornell,

1998; Herzberg, 2004).

Melts lying along path a-b in Fig. 9a crystallize ol+plag and have compositions that do not carry a signature of cpx crystallization. Nominal pressures calculated for these melts do not accurately reflect the pressure of magma evolution, but the lowest calculated pressure (for melts with compositions close to b) represents an upper limit for cotectic crystallization.

Melts plotting along the path c-d-e in Fig. 9b isobarically crystallize ol-plag-cpx along the cotectic and then crystallize ol+plag during ascent. Melts saturated in ol+plag with compositions between d and e carry a signature of cpx crystallization, but calculated nominal pressures do not accurately reflect the pressure of magma evolution. In this case, the highest pressure (for melts with compositions close to d) represents a lower limit for cotectic crystallization.

Mixing of ol±plag saturated melts with ol-plag-cpx saturated cotectic melts (Fig.

9c, path f-g-h) will produce hybrids (e.g., path f-h) with the compositional signature of cpx crystallization (Langmuir, 1989). The lowest pressure calculated for the hybrid melts represents an upper limit for cotectic crystallization.

There is petrographic evidence that some Icelandic magmas, notably those from the Hengill complex, evolve by crystallization of ol+plag combined with assimilation of cpx rather than by cotectic crystallization of ol-plag-cpx (Trønnes, 1990; Hansteen, 1991;

Gurenko and Sobolev, 2006). If assimilation occurs after melts leave the ol-plag-cpx

cotectic and ascend towards the surface (Fig. 9d, path j-k), as appears to be the case at

Hengill, the highest calculated pressure represents a lower limit for cotectic

crystallization. Note, however, that it is extremely difficult to discriminate between

magmas that evolve via ol-plag-cpx cotectic crystallization and those that evolve via

ol+plag crystallization accompanied by assimilation of cpx if petrographic or other (eg.

isotopic) evidence for assimilation is lacking. Indeed, it is possible that the high CaO

contents and high CaO/Al2O3 ratios of some glasses reflect assimilation of cpx.

Consequently, the estimated pressure of cotectic crystallization may be too low if all

samples from a particular locality (e.g., Hrimalda) have these characteristics (see Fig.

9d). However, the glasses from most localities show a range of CaO and CaO/Al2O3, and

it is likely that the highest pressures calculated for these glasses provide a reasonably

accurate estimate of the pressure of cotectic crystallization.

Figure 8. Histograms illustrating raw (unfiltered) results obtained for samples from individual localities. Fig. 8a. Results for Gigoldur (diagonal shading) indicate crystallization at low pressure, whereas those from Kalfstindar (grey stippled pattern) indicate crystallization at high pressure. Fig. 8b. Results for Bláfjall Table Mountain indicate crystallization at two distinct pressures.Fig. 8c. Glasses from Laki indicate crystallization over a wide range of pressure. Fig. 8d. Results for the Hengill complex (Hengill, Maelifell, Midfell) indicate crystallization over a wide range of pressure, though many samples appear to have crystallized at P<0.2 GPa.

Plots of P versus MgO for samples from four localities are shown in Fig. 10 to

illustrate interpretation of the results in light of the relationships described above. For

each locality, plots of CaO and CaO/Al2O3 versus MgO, together with available

petrographic data, were used to discriminate between ol+plag+cpx saturated melts and

ol±plag±spinel saturated melts. For both Herdubreid and Hlodufell (Fig. 10a,b), there is a positive correlation between P and MgO for glasses in equilibrium with ol±plag±spinel whereas there is no correlation between P and MgO for glasses in equilibrium with ol+plag+cpx.

Calculated pressures for ol-plag-cpx saturated melts from Herdubreid (Fig. 10a) define two arrays parallel to the MgO axis that are interpreted to indicate cotectic crystallization at 480±30 MPa and 310±20 MPa. The low pressure calculated for one sample could be an aberrant result, possibly caused by analytical error, or could indicate crystallization along an even lower pressure cotectic (~170 MPa). This interpretation is preferred because results obtained for other localities (Hlodufell, Grimsvötn,

Efstadalsfjall, Askja, Hrimalda, Kverkjoll, Theistareykir, Thingvellir) also provide evidence for low-pressure (90-200 MPa) cotectic crystallization.

Figure 9. Possible interpretations of calculated pressures illustrated using phase relationships projected from plag onto the pseudoternary plane ol-cpx-qtz with the procedure of Tormey et al. (1987). Solid lines mark the positions of the ol-plag-cpx cotectic (P in GPa). POP – Pressure of crystallization of ol-plag assemblages. POPC – Pressure of crystallization of ol-plag-cpx assemblages. PMax – Maximum pressure of crystallization. PMin – Minimum pressure of crystallization. PUL – Upper limit for pressure of crystallization . PLL – Lower limit for pressure of crystallization. Arrows indicate changes in melt composition. Fig. 9a. Melts evolve from a to b by polybaric crystallization of ol-plag. Fig. 9b. Melts evolve from c to d by isobaric crystallization of ol-plag-cpx and from d to e by polybaric crystallization of ol-plag.Fig. 9c. Melts evolve from f to g by polybaric crystallization of ol-plag and from g to h by isobaric crystallization of ol-plag-cpx. The dashed line shows mixing between primitive melt f and evolved melt h. Fig. 9d. Melts evolve from i to j by isobaric crystallization of ol- plag-cpx and from j to k by polybaric crystallization of ol-plag accompanied by assimilation of cpx.

Pressures calculated for ol-plag-cpx saturated melts from Hlodufell (Fig. 10b) define a single broad array parallel to the MgO axis, and could be interpreted to indicate cotectic crystallization between maximum (PMax=380 MPa) and minimum (PMin=290

MPa) pressures. However, the range in pressure is about equal to the expected precision

of the results, and therefore, the results are interpreted as indicating cotectic

crystallization at 340±40 MPa. As with Herdubreid, the low pressure (~170 MPa)

calculated for one glass is interpreted to indicate crystallization along a lower pressure

cotectic.

Calculated pressures for ol-plag-cpx saturated melts from Grimsvötn (Fig. 10c)

show more scatter than those from Hlodufell, and the range is greater than estimated

precision. The results are interpreted to indicate cotectic crystallization between

maximum (PMax= 510 MPa) and minimum (PMin~100 MPa) pressures, but there is clear

evidence for cotectic crystallization at intermediate pressures (see below). A more

realistic estimate of the pressure of low-pressure cotectic crystallization is obtained by

averaging the results for samples that crystallized at similar pressures. Four samples yield

pressures between ~100 and 200 MPa, whereas other samples yield pressures >300 MPa

(Fig 11c). The average pressure for the four samples is 150±60 MPa and is the preferred

value for low pressure cotectic crystallization. Preferred values for low-pressure cotectic

crystallization have been calculated for Theistreykir, Bardabunga, Grimsvötn, Veidivötn,

Katla, Hekla, and Thingvellir. Likewise, results for samples that crystallized at similar

high pressures can be averaged and used to calculate preferred values for high-pressure

cotectic crystallization at some localities (eg.Hengill).

Finally, glasses from Kalfstindar represent melts saturated with ol+plag, and the lowest calculated pressure provides an upper limit (PUL=650 MPa) for cotectic

crystallization (Fig. 10d). Note that pressures calculated for some ol-plag saturated melts

from Herdubreid and Hlodufell are also higher than those calculated for ol-plag-cpx

saturated melts from these localities (Figs. 11a,b).

Figure 10. Interpretation of results for selected localities illustrated by plots of P versus MgO. Open circles – melts in equilibrium with ol±plag±spinel assemblages. Filled circles – melts in equilibrium with ol-plag-cpx assemblages. PC – Pressure of cotectic crystallization. PMax – Maximum pressure of crystallization. PMin – Minimum pressure of crystallization. PUL – Upper limit for pressure of cotectic crystallization. See text for further discussion.

Results from the Hengill complex are anomalous compared with those obtained

from other volcanic centers. Plots of CaO and CaO/Al2O3 versus MgO are consistent with

crystallization of ol+plag+cpx. Nominal pressures calculated for 120 glasses range from

0.4 to 754 MPa (Table 2), and the simplest interpretation is that cotectic crystallization

occurred over a range of pressures between these values. However, detailed studies by

Trønnes (1990), Hansteen (1991), and Gurenko and Sobolev (2006) indicate that this

interpretation is oversimplified. Trønnes (1990) divided Hengill glasses into four groups based primarily on MgO and CaO content. I followed his subdivision for Groups I and II, but have combined glass analyses from his two other groups into a single group (Group

III). Nominal pressures for MgO-rich samples from Group I range from ~0.1 to 311 MPa

(Fig. 11). These samples contain Al- and Cr-rich endiopside as resorbed phenocrysts and as components of partly disaggregated gabbroic nodules, suggesting that the MgO-rich

Group I magmas have assimilated crustal material. The highest calculated pressure for

these samples (311 MPa) represents the lower limit for cotectic crystallization (eg. Fig.

9d). On the other hand, nominal pressures for MgO-poor Group III samples range from

340 to 613 MPa (Fig. 11). Some of these samples contain augite microphenocrysts,

suggesting that the MgO-poor Group III magmas have crystallized ol+plag+cpx along

cotectics between a maximum (613 MPa) and minimum (340 MPa) pressure. The lower limit of cotectic crystallization for Group I glasses and the minimum pressure of cotectic crystallization for Group III glasses are virtually identical (Fig. 11) and provide a tight constraint on low-pressure cotectic crystallization (325 MPa). Pressures calculated for

Group II glasses encompass values obtained for Group I and Group III glasses (Fig. 11).

Crystallization of MgO-rich Group I melts at lower pressure than the MgO-poor Group

III melts produces the unusual negative correlations between P and MgO and P and T observed for Hengill samples. The strong evidence that Group I magmas have interacted with gabbroic crust suggests that the negative correlation between P and MgO for Hengill samples results from assimilation combined with crystallization as magmas ascend beneath this volcanic complex.

Figure 11. Interpretation of results for glasses from the Hengill complex. Fig. 11a. Histogram of results obtained for MgO-rich (Gp I) and MgO-poor (Gp III) samples. The Gp I samples evolve via polybaric crystallization of ol-plag accompanied by assimilation of cpx which allows the lower limit of cotectic crystallization to be determined (PLL). The Gp III samples evolve via polybaric crystallization of ol-plag-cpx which allows the minimum pressure of cotectic crystallization to be determined (PMin). Fig. 11b. Plot of P versus MgO. Open circles - Gp I samples. Filled circles - Gp III samples. . PMax – Maximum pressure of cotectic crystallization. PMin – Minimum pressure of cotectic crystallization. PLL – Lower limit for pressure of cotectic crystallization.

In summary, melts that lie along ol-plag-cpx cotectics can be identified allowing

crystallization pressures to be established. Similarly, melts in equilibrium with ol±plag can be identified and used to place upper limits (PUL) and/or lower limits (PLL) on

crystallization pressures. Finally, melts that provide no constraints on crystallization

pressure can be identified and filtered out of the results.

Pressures of ol-plag-cpx cotectic crystallization for all localities are listed in Table

3. A striking feature of the results (Fig. 12) is the large number of samples (85%) that

have crystallized along either low-P (40-220 MPa) or relatively high P (430-1030 MPa)

cotectics. This near-bimodal distribution of pressures contrasts strongly with the

distribution of nominal pressures obtained using unfiltered results (Fig. 7a), and

illustrates the importance of evaluating geochemical and petrographic data for samples

from individual localities to identify melts that lie along ol-plag-cpx cotectics. It is also

clear that some melts crystallized along ol-plag-cpx cotectics at pressures between 220

and 430 MPa (Table 3).

------7.0 - 6.2 ±2.2 5.3 ±1 6.0 ±2.6 5.4 ±2.0 7.1 - 5.0 ±2.3 6.5 ±1.99 5.0 - 6.2 ±2.3 ------

- -

Min Min 7.0 - Min 3.2 - Min Min 6.2 - Min 5.9 - Min Min 2.6 - Min 1.9 - Min Min 5.1 - Min 6.1 - Min Preferred - - - 62 Preferred 58 Preferred ±64 ± ± ------90 - 177 154 149 ±27 Preferred 141 199.3 - - 170 ±90 - 176 ±90 - 97 - - 202.0 - - 90 - - 131 - - - 175 - - - 35 - - 185 ±56 - 74 - - - 54 - - 145 - - 65 ------9.1 ±1.2 16.4 ±1.6 17.7 ±1.7 10.9 ±0.51 - - 12.2 ±1.14 - - 11.8 ±1.3 168.0 - 10.9 ±0.66 173 - 11.7 - 11.3 ±1.3 198.0 - Limit 11.0 -

Min 12.0 - Min Min 11.4 - 143 - Min Min ------33 Preferred ± ------340.0 - 257.0 466.3 ±45 Preferred 501.7 ±48 Preferred - 310.0 ±15 - 340.0 - - - 311.0 - Limit Lower 10.9 - - - - - 331.0 - - 335.0 ±37 - 313.0 - Upper 2.6 ± ------17.7 ±3.01 19.6 ±0.88 18.9 ±0.32 - 23.1 17.8 ±0.91 346.0 ±32 25.3 ±0.89 - 18.2 ±1.26 322.0 ±36 -

- - Max 36.3 - 425.0 - Max Preferred 34.0 Preferred

74 0.86 Preferred 0.25 Preferred 0.33 Preferred ± ± ± ± ------484 ±30 17.0 ±1.1 308.0 ±19 593 - Max 20.9 - 512 - Max 18.0 - 842 - Max 29.6 - 859 - Max 28.9 ±1.2 - 555 - Max536 ±9 19.5 - 529511 - ±76 Max 18.0 18.6 ±2.3 - - 622 - Max 21.9 - 652 - Max 22.9 - 717 ±25 518 - Max 18.2 - 325.0 - 656 512 - Max 842 - Max 956 622 - Max 555 - Max 822 1032 613.0 - Max 592.0 - Max 507.0 ±26 517.0 ±34 627.0 - Max 22.1 - 221.0 - 651.0 - Limit Upper 22.9 - 503.0 556.0 627.0 - Max Locality P MPa Range Comment Km Z Range P MPa Range Comment Km Z Range P MPa Range Comment Km Z Range Hengill Gp III Hekla Hekla Hengill Gp III Thingvellir Hengill Gp II Hegill GpII Hengill Gp I Kverkfjoll Kverkfjoll Sprengisandur Gigoldur Grimsvotn Hrimalda Bardabunga Grimsvotn Theistareykir Theistareykir Herdubreid Askja Bardabunga Katla Langjokull Geitafell Burfell Blafjall Ridge Halar Seljahjalli Reykjanes Thingvellir Raudafell Efstadalsfjall Kalfstindar Veidivotn Veidivotn Hlodufell Kistufell Laki 766 - Max 26.9 - Northeastern Table 3. Pressures and Depths of Cotectic Crystallization Cotectic of Depths and Pressures 3. Table RP WVZ Zone NVZ EVZ SFZ 41

Effect of H2O

The method used to calculate pressure is based on experiments carried out under nominally anhydrous conditions, and it is important to assess the possible effect of water on the results. Experimental studies show that addition of H2O leads to expansion of the stability field of olivine and contraction of the stability field of plagioclase (Kushiro,

1969; Nicholls and Ringwood, 1973; Baker and Egger, 1987; Sisson and Grove, 1993), so that the location of the ol-plag-cpx cotectic will shift with increasing water content at constant pressure. However, relatively high water contents are required to produce large shifts in the position of the cotectic, and the water contents of Icelandic magmas are probably too low to have much effect. Measured water contents range from 0.1-1 wt%

(Jamtveit et al., 2001; Nichols et al. 2002), and agree with water contents calculated for

88 Icelandic glasses using the method of Danyushevsky et al (1996) and Danyushevsky

(2001) (average – 0.18 wt%, range 0-0.94 wt%). Herzberg (2004) found no correlation between calculated pressures and H2O for MORB with 0.1 to 0.8 wt% H2O. Additional evidence that calculated pressures have not been significantly affected by water is provided by the close agreement (40±40 MPa) between pressures calculated for individual samples using two different projections (see Appendix I), one from plag onto the plane ol-cpx-qtz, the other from ol onto the plane plag-cpx-qtz. Addition of water causes the ol-plag-cpx cotectic to shift towards cpx (away from ol) in the plag projection and away from cpx (towards plag) in the ol projection. In other words, water affects the stability fields of ol, plag, and cpx differently and has a different effect on the positions of

the ol-plag-cpx cotectic in these two projections. Accordingly, pressures calculated for

hydrous melts will be associated with large uncertainties reflecting large differences

between pressures calculated from the two different projections (see Appendix A, Table

A.1). As noted in a preceding section, the uncertainty in calculated pressure for most samples is similar to that estimated from anhydrous experimental data, and it is concluded that the results reported in Tables 4 and 5 closely reflect the pressures of crystallization. Pressures associated with large uncertainties have been filtered out of the results. It is unlikely that these samples crystallized at higher water contents than most

Icelandic magmas, because water contents calculated using the method of Danyushevsky et al (1996) and Danyshevsky (2001) (average 0.28 wt%, range 0-0.85 wt%) are similar

to those obtained for other Icelandic glasses. Use of this method also indicates that

samples that yield nominally anhydrous pressures between 0 and -100 MPa did not

crystallize from magmas containing unusually high water contents, supporting the

interpretation that these samples crystallized at ~0.1 MPa taking uncertainties into

account.

Depths of Magma Chambers and Magma Plumbing Systems

The pressures reported in Table 3 are those at which melts are last saturated with

ol, plag, and cpx. Ascending magmas must pause and crystallize for a sufficiently long

period of time to reach multiple-saturation, and this most likely occurs in magma chambers. The melt is then rapidly erupted from the chamber and quenched to form glass.

The depths of magma chambers may therefore be calculated from pressures of cotectic crystallization using an appropriate value for crustal density. We have used a value of

2900 kg m-3 to calculate the depths listed in Table 3. Preferred values for the pressure of cotectic crystallization were used in the calculations wherever possible.

Figure 12. Histograms showing averaged results for the pressure of cotectic crystallization at individual volcanic centers. Fig. 12a. Cotectic crystallization at relatively high pressure. Fig. 12b. Cotectic crystallization at relatively low pressure. Results indicating cotectic crystallization at intermediate pressures (P>0.22 GPa, <0.43 Gpa) omitted for clarity.

Comparison with results obtained using geodetic and/or geophysical methods (see

Table 1) reveals agreement for the depths of chambers beneath Askja (6.5 and 18.9 km

versus 1.3-3.5 and 16-20 km) and Grimsvötn (5.4 km versus 1.5-4 km), and reasonable

agreement for the depth of a chamber beneath Hengill (11.5 km versus 7 km). Our

estimate of a minimum depth of 17.5 km for a chamber beneath Hekla is greater than the

recent estimate of 11 km based on strain data (see Sturkell et al. 2006), but is consistent

with recent analysis of seismic data (Soosalu and Einarsson, 2004) which provide no

evidence for a significant magma body at depths < 14 km. However, we find no evidence

for the shallow chamber (2-5 km) identified beneath Katla in geophysical and geodetic

studies. This probably indicates that this chamber does not contain basalt, because

Soosalu et al. (2006b) suggest that a cryptodome containing relatively viscous (silica-

rich) magma occurs at shallow (2 km) depths beneath Katla.

Comparison with results obtained in other petrological studies (see Table 1)

reveals excellent agreement in the case of Bláfjall Table Mountain (6.2 and 18.6 km

versus 2-7 and 14-26 km) and the Hengill volcanic complex (11-12 km versus 8-12 km).

However, the depths obtained for crystallization at Kistufell (<18 km) are much lower

than those obtained by Breddam (2000) (>30 km), and the reason for this discrepancy is

unclear.

My results suggest magma chambers that we located at different depths in

Icelandic crust (Figs 14 and 15), and this is corroborated by geodetic and geophysical

studies. There is evidence for only one chamber located in the shallow (Hrimalda and

Gigoldur in the NVZ), middle (Burfell in the NVZ, Raudafell in the WVZ), or deep

(Kalfstindar in the WVZ) crust beneath some centers, but there is clear evidence for two

or more stacked chambers beneath most centers (Fig. 13, 15). Shallow and deep

chambers occur beneath Askja, Bláfjall Table Mountain (including Halar and Seljahalli –

see Schiellerup, 1995) and Langjokull (Fig. 15a), whereas shallow, intermediate, and deep chambers occur beneath Herdubreid and Efstadalsfjall (Fig. 15b). Dikes presumably connect chambers located at different depths, and provide conduits for magmas to reach

the surface.

Figure 13. Depths of shallow and deep magma chambers along the rift zones. Error bars show uncertainty in estimated depth. Abbreviations for volcanic centers are given in Appendix C and Figure 1. Volcanic centers are arranged from south to north along the y- axis. The shaded grey bands show the range of depth of shallow and deep chambers along the NVZ and WVZ. Note the greater depths of chambers beneath Kverkfjoll, Hekla, Grimsvötn-Laki, and Bardabunga-Veidivötn. Also note the lack of evidence for shallow chambers beneath Katla and Hekla.

Figure 14. Depths of intermediate depth magma chambers along the rift zones. Error bars show uncertainty in estimated depth. Abbreviations for volcanic centers are given in Appendix C and Figure 1. Volcanic centers are arranged from south to north along the y- axis. The shaded grey bands show the range of depth of shallow and deep chambers along the NVZ and WVZ.

Results for Theistareykir, Sprengisandur, Kverkfjoll, Bardabunga-Veidivötn,

Grimsvötn-Laki, Langjokull, Thingvellir, and Hengill are consistent with the presence of both shallow and deep chambers, and magmas also appear to have crystallized at intermediate depths even though there is no evidence that this occurred in a discrete chamber (see Fig. 10c). As discussed previously, it is possible that these magmas evolved by high pressure cotectic crystallization of ol+plag+cpx followed by crystallization of ol+plag (with partial or complete resorbtion of cpx) during ascent. However, petrographic data indicate that at least some glasses are from magmas that evolved via ol+plag+cpx crystallization, which suggests the presence of cooling and crystallizing magma bodies or

pockets throughout the middle and lower crust. Multiple, stacked, discrete chambers might exist beneath some volcanoes as shown in Fig. 15c. However, it is also possible that crystallization occurs in extensive crystal mush zones as suggested by Hansen and

Grönvold (2000). In addition, crystallization probably occurs at various depths in feeder

dikes and conduits during lulls in eruptive activity.

Recent work suggests that MORBs partially crystallize over a range of pressures

from 1 to 1000 MPa (Michael and Cornell, 1998; Le Roex et al. 2002; Herzberg, 2004).

The average crustal thickness along mid-ocean ridges is 7.1±0.9 km (White et al. 1992;

Bown and White, 1994), so that crystallization of MORB must begin in the upper mantle.

Some Icelandic magmas may also crystallize in the mantle. The glasses have chemical

characteristics of evolved magmas and must have formed from primary magmas by

crystallization in the deep crust or mantle. Conclusive evidence for crystallization in the

mantle is lacking, but the high pressures (up to 800 MPa) obtained for ol±plag saturated

melts from Herdubreid, Efstadalsfjall, and Kalfstindar (Table 2) are consistent with

crystallization at upper mantle depths.

Figure 15. Schematic representation of plumbing systems beneath Icelandic volcanoes. There is evidence for two or more chambers beneath most volcanic centers (see also Gudmundsson, 2000). Fig. 15a. Shallow (a) and deep chambers (b) linked by a system of conduits and dikes (c). A plumbing system similar to this may exist beneath Askja, Bláfjall Table Mountain and, possibly, Langjokull. Fig. 15b. Shallow (a), intermediate (b) and deep chambers (c) linked by a system of conduits (d) and dikes (e). A plumbing system similar to this may exist beneath Herdubreid and Efstadalsfjall. Fig. 15c. Shallow (a) and deep chambers (b) with a plexus of small chambers or magma bodies in the lower crust (c) linked by a system of conduits and dikes (d). This type of plumbing system probably exists beneath Theistareykir, Sprengisandur, Kverkfjoll, Bardabunga-Veidivötn, Grimsvötn-Laki, Langjokull, Thingvellir, and Hengill. The plumbing systems beneath Katla and Hekla in the SFZ may also be similar to the middle-lower crustal section.

Thickness and Structure of Icelandic Crust

The factors that determine the particular depth at which magma chambers form are complex and are poorly understood, but include the buoyancy force (reflecting the density contrast between melt and surrounding rocks), any excess pressure over the buoyancy force, and local variations in stress conditions in the crust (Gudmundsson,

2000) . It can be assumed to a first approximation that variations in the buoyancy force, excess pressure, and stress conditions are negligible for magmas entering the base of the crust along the rift zones. Therefore, the main control of the location of deep chambers is likely to be a lithologic boundary associated with a density discontinuity. These chambers occur over a remarkably small range of depths for volcanic centers in the NVZ (19.5±2.6 km, excluding Kverkfjöll) and WVZ (20.9±2.3 km, including samples from

Geitafell)(Fig. 13b). The small range of depths for the rift zones (20.1±2.5 km) suggests that the deep chambers occur at or near the crust-mantle boundary (Kelley et al., 2004;

Kelley and Barton, 2005). There is excellent agreement between magma chamber depth beneath Theistareykir and crustal thickness inferred from seismic data (Brandsdóttir et al,

1997; Staples et al. 1997) beneath nearby Krafla (19.5 versus 19 km). Likewise, there is very good agreement between magma chamber depths and crustal thickness inferred from seismic data (Bjarnason et al. 1993; Weir et al. 2001) for Geitafell (17.8±0.9 versus 16 km), Hengill ( 18.9±1.8 versus 17.5 - 24 km), Hekla (34±2.6 versus 30-35 km), and Katla

(23.1±1.6 versus 20-25 km).

Seismic data do not provide reliable estimates of the depth of the crust-mantle

transition beneath other centers, and crustal thickesses must be estimated from gravity

data (Darbyshire et al. 1998, 2000a; Allen et al. 2002; Kaban et al., 2002; Leftwich et al.

2005; Fedorova et al. 2005) and from the relationship between Moho depth and height

above sea level given by Gudmundsson (2003). These methods do not allow small-scale

variations in crustal thickness to be resolved, and there is considerable uncertainty in the

estimated crustal thickness beneath most volcanoes. In addition, the crust-mantle

boundary may represent a transition zone 5±3 km thick (Foulger et al. 2003). Despite

this, there is reasonable agreement between estimated magma chamber depth and the

base of the crust beneath Kverkfjöll (29 versus 30-35 km) and beneath

Bardabunga/Veidivötn and Grimsvötn/Laki (27-29 km versus 30-40 km) given that

absolute uncertainties in calculated magma chamber depths are ~±3.9 km. These volcanic centers are located on or near the Vatnajökull ice cap (Fig. 1) in a region of thicker crust above the thermal and/or compositional anomaly in the mantle (Darbyshire et al. 2000a;

Kaban et al. 2002; Leftwich et al. 2005; Fedorova et al. 2005). Also, depths estimated for magma chambers and the base of the crust are similar for the Bláfjall complex (~25 km),

Thingvellir (22 km versus 20-25 km), and Kalfstindar (23 km versus 20-25 km).

The agreement between magma chamber depths and Moho depths provides strong evidence that magmas pond at the base of the crust beneath most volcanic centers in

Iceland. However, there is poor correlation between Moho depth and magma chamber depth for some centers, and it is possible that the true, maximum depth of chambers

beneath these centers is greater than that reported in Table 3. As noted previously, some ol-plag saturated melts from Herdubreid and Efstadalsfjall appear to have crystallized at higher pressure than the maximum value calculated for cotectic melts. Also, I am not confident that the maximum depth of chambers has been established for Langjokull,

Sprengisandur and Askja, because of the relatively small number of samples available for study.

The depth of shallow chambers is also relatively constant for volcanic centers along the rift zones (Fig. 13a). The average depths are 5.2±1.6 km for the NVZ, 6.1±2.1 km for the WVZ (including samples from the Reykjanes Peninsula), and 5.4±0.6 km for the EVZ. The average depth for all three rift zones is 5.5±1.6 km, indicating that the chambers are located in the upper crust, which has an average thickness of 5-7 km 2

(Darbyshire et al. 2000b; Foulger et al. 2003). Various authors use different definitions of upper, middle, and lower crust for Iceland (see Foulger et al. 2003). For this paper, I arbitrarily define upper crust as that <7 km, middle crust as that from 7-15 km, and lower crust as that >15 km.The upper crust is a heterogeneous mixture of flows, hyaloclastites, and intrusives that have been metamorphosed under greenschist-amphibolite facies conditions. The observed rapid increase in seismic velocity with depth probably reflects a decrease in the proportion of hyaloclastites and decrease in porosity and permeability

(Foulger et al.2003, and references therein). It seems likely that shallow chambers form at a level of neutral buoyancy at or near the base of the upper crust, although the exact depth

may be controlled by local variations in lithology and/or stress conditions

(Gudmundsson, 2000).

Additional petrologic studies to refine estimates of magma chamber depth are

desirable to provide tighter constraints on the thickness of the whole crust and of the upper crust, as well as to define regional variations in these thicknesses. The occurrence of discrete chambers located at intermediate depths (Fig. 14) also has important implications for models of crustal structure. Those in the NVZ and WVZ occur at a relatively constant depth (11.4±0.4 km). There is no convincing evidence for a seismic discontinuity at this depth, and it seems likely that these chambers are located at a phase transition (amphibolite-granulite facies?), and/or at a rheologic boundary. Fedorova et al.

(2005) suggested that the base of the elastic lithosphere beneath Vatnajökull is 15 km deep, and is underlain by relatively “soft” lower crust. The “intermediate depth” chambers beneath Katla and Hekla are significantly deeper (~16-18 km) than the intermediate depth chambers in the active rift zones, which may indicate a different lithological structure or stress regime in the crust of the SFZ.

The evidence for cooling and crystallizing magma bodies or pockets throughout the middle and lower crust is consistent with results obtained by Tryggvason (1986) for

Krafla and by Maclennan et al. (2001a) for Theistareykir (Table 1). The presence of multiple magma chambers at different depths beneath many volcanic centers implies high geothermal gradients in the crust (see also Meyer et al. 1985; Bjarnason et al. 1993;

Maclennan et al. 2001a; Leftwich et al. 2005; and Kelley et al., 2005a). Petrologic data

therefore strongly suggest that the middle and lower crust is relatively hot and porous, in

contrast to the relatively cool, rigid crust proposed in some geophysical studies (e.g.,

Bjarnasson et al. 1993; Menke and Sparks, 1995). However, receiver function data

(Darbyshire et al. 2000b), combined surface and body wave constraints (Allen et al.

2002), combined results from explosion seismology and receiver functions (Foulger et al.

2003), and combined results from seismic, topography, and gravity data (Fedorova et al.

2005) all provide evidence for low velocity zones (LVZs) in the crust. These may reflect

variations in temperature or lithology, and/or the presence of melt. Most workers favor an

explanation involving both high temperatures and the presence of small amounts of melt.

Darbyshire et al. (2000b) suggest that a prominent LVZ at 10-15 km reflects the presence

of partially molten sills in the lower crust beneath Krafla (cf. Fig. 15c), whereas Fedorova

et al. (2005) proposed that crust deeper than 15 km beneath Vatnajökull contains melt.

Allen et al. (2002) identified LVZs in both the upper (0-15 km) and lower (>15 km) crust.

The former are thought to reflect regions of high temperature and possible melt, whereas

the latter (beneath Vatnajökull) are thought to reflect the thermal halo associated with the fluxing of magma from the mantle to the upper crust. Although detailed correlation of

LVZs with the occurrence of magma chambers is not possible, these results suggest that

seismic and petrologic data can be reconciled to develop an internally consistent model

for Icelandic crust.

Crustal Accretion Models

There is no doubt that some crustal accretion occurs at shallow depths in Iceland

by eruption of lavas and hyaloclastites that are buried beneath the products of later

eruptions, and by crystallization of gabbros in shallow chambers. Shallow accretion is

one of the models proposed for the formation of oceanic crust. Melt is supplied directly

from the mantle to a shallow axial melt lens (Sinton and Detrick, 1992). Some of this

melt is erupted whereas the rest crystallizes to form gabbros that subside to form oceanic

crust. However, it is highly unlikely that the thick (20-40 km) crust in Iceland forms

solely by crystallization in shallow chambers to form gabbroic crust that subsides in

response to plate extension. It is probable that accretion also results from crystallization

in chambers located at Moho depths (underplating), and by crystallization in dikes, melt pockets, or discrete chambers at various depths in the crust (intracrustal accretion). A

model for accretion of Icelandic crust is shown in Fig. 16, and is consistent with recent

seismic reflection and petrologic data for mid-ocean ridges (eg. Singh et al. 2006;

Maclennan et al. 2004; Pan and Batiza, 2002, 2003). These data indicate that lower crust

in axial regions is at least locally porous on an intergranular scale, and suggest that

accretion occurs by crystallization of small pockets of melt over a range of depths

between the shallow melt lens and the Moho.

Figure 16. Petrologic model for Icelandic crust. Shallow chambers are located near the base of the upper crust, and magma is fed to the surface via dike systems. Crustal accretion occurs by eruption of lavas and hyaloclastites, and by crystallization of gabbro in the chamber. The middle and lower crust consists of smaller chambers, pockets of magma, and conduits within a mush zone extending to the Moho. A chamber at the Moho feeds magma to the shallower chambers and directly to the surface. The mush zone may extend to shallower depths than shown (i.e., to the base of the upper crust). Crustal accretion occurs by crystallization of magma within the mush zone and at the base of the crust to form gabbros. Magma chambers or pockets of magma may also occur in the underlying mantle. See text for details.

Petrologic Implications

The results of this study are consistent with complex models of magma evolution that involve polybaric crystallization, magma mixing, and assimilation. Polybaric crystallization is expected in plumbing systems with two or more chambers located at different depths, and supporting evidence is provided by detailed mineral chemical studies of lavas that reveal the presence of different generations of minerals that formed at different pressure (e.g., Hansteen, 1991; Hansen and Grönvold, 2000, Maclennan et al.

2001a). High-pressure crystallization yields residual melts that are compositionally different from those produced by low pressure crystallization (eg. Kinzler et al. 1992), so that polybaric crystallization will produce liquids lying along different liquid lines of descent (see Figs. 6c and 7). It is also likely that these magmas will mix prior to and during sequential eruptive episodes because magmas that pond and partially crystallize in chambers, dikes, and conduits after one eruptive episode will be flushed out and mix with fresh batches of magma rising from deep chambers or from the mantle during subsequent eruptive episodes. Evidence for mixing has been described by Sigurdsson and Sparks

(1981), McGarvie (1984), Mørk (1984), and Fagents et al. (2001). Moreover, crystal aggregates formed by partial crystallization of magma in the plumbing system in an early eruptive phase can be disrupted and incorporated into later batches of magma rising from deep chambers or from the mantle. This yields magmas with complex crystal cargoes derived from different magma batches, such as those described by Hansen and Grönvold

(2000). Interaction between ascending melts and the crystalline products of earlier

59 eruptive episodes can lead to complex variations in melt composition (Kvassnes et al.

2003; Danyushevsky et al. 2004), and interaction between ascending melts and clinopyroxene appears to explain the unusual chemical characteristics (e.g., high MgO,

CaO, CaO/Al2O3) of some magmas erupted in the Hengill complex and elsewhere along the rift zones (e.g., Trønnes, 1990; Hansen and Grönvold, 2000; Gurenko and Sobolev,

2006).

Oxygen isotope studies provide strong evidence for assimilation of hydrothermally altered crust by Icelandic magmas (e.g., Condomines et al. 1983;

Hemond et al. 1988, 1993). Assimilation may be extensive if the middle and lower crust is hot, as suggested above, because less energy is required for a fixed mass of magma to assimilate hot crustal material. Moreover, frequent replenishment with new batches of magma during mixing events leads to thermal buffering in magma chambers, which also facilitates assimilation (Cribb and Barton, 1996). Certainly, the high temperature of basaltic magma emplaced into Icelandic crust (average-1186OC using the method of

Yang et al. (1996) and 1177OC using the method of Sugawara (2000)), together with the relatively small temperature range of magmas erupted at individual volcanic centers

(average-34OC), suggests that abundant heat is available to drive assimilation processes.

High geothermal gradients coupled with high magma temperatures will also facilitate melting of crustal lithologies to form silicic magmas, which are relatively abundant on

Iceland (Gunnarsson et al. 1998). My results suggest that silicic magmas can be

generated over a relatively wide depth range in the middle and lower crust and this prediction can be tested by additional studies of appropriate compositions.

CONCLUSIONS

A method to calculate pressures of crystallization of melts lying along ol-plag-cpx

cotectics based on the procedure described by Yang et al. (1996) yields results that are

accurate to ±110 MPa (1σ) and are precise to 80 MPa (1σ). Pressures calculated for

Icelandic glasses from 29 volcanic centers in the Western, Northern and Eastern rift

zones, as well as from the Southern Flank Zone, indicate crystallization range from 1 to

~1000MPa, equivalent to depths of ~0-35 km. Magma chamber depths estimated from

these results agree well with those estimated using other methods for Askja, Bláfjall

Table Mountain, Grimsvötn, Hengill, and Hekla.

Deep chambers (>17 km) occur beneath most volcanic centers and appear to be

located at the Moho, indicating that magma ponds at the crust-mantle boundary. Shallow

chambers (< 7.1 km) also occur beneath most volcanic centers. These are located in the

upper crust, and probably form at a level of neutral buoyancy. There is good evidence for

cooling and crystallizing bodies and pockets of magma throughout the middle and lower

crust, which probably resembles a crystal mush. This strongly suggests that the middle

and lower crust is relatively hot and porous, and that crustal accretion occurs over a range

of depths as inferred in recent models for crustal accretion at mid-ocean ridges. The

presence of multiple, stacked chambers and hot, porous crust suggests that magma

61 evolution is complex and involves polybaric crystallization, magma mixing, and assimilation.

Chapter 3: Petrologic Imaging of the Magma Chamber beneath Upptyppingar, Kverkfjöll Volcanic System, Iceland

Knowledge of the depths of magma chambers is important to constrain models for

magma evolution because phase relationships and melt compositions vary as a function

of pressure (O’Hara and Herzberg, 2002). In addition, such knowledge is necessary to

understand mechanisms of crustal accretion and differentiation (Pan and Batiza, 2003),

and to interpret precursory activity (seismic, deformation, gas emissions, etc.) to volcanic

eruptions (Pan and Batiza, 2003). Recently developed petrologic methods allow the

pressures and hence depths of magma crystallization to be estimated with reasonable accuracy (Herzberg, 2004; Kelley and Barton, 2008). Here we show that glasses in magmas erupted at the Kverkfjöll volcanic system in the Northern Volcanic Zone (NVZ)

have compositions that are consistent with partial crystallization at average pressures of

445±69 MPa and 794±92 MPa, corresponding to depths of 15.6±2.4 km and 27.9±3.2

km, and conclude that magma chambers are located at these depths. These results are

consistent with interpretation of recent seismic activity beneath Upptyppingar in the

Kverkfjöll volcanic system ~50 km north of the Kverkfjöll central volcano. The

earthquake hypocenters are concentrated at depths of 15-18 km with a few occurring at

greater depths (~25 km), and the seismic activity appears to reflect inflow of magma into

the base of the crust (Roberts et al., 2007; Jakobsdottir et al., 2008). The seismic unrest

at Upptyppingar may herald the onset of a rifting episode along the NVZ. Eruption of

magmas from chambers in the middle to lower crust that reach the surface at near

liquidus will promote rapid degassing leading to explosive activity that could inject

volcanic aerosols sufficiently high in the atmosphere to affect the Earth’s climate.

GEOLOGIC SETTING

Iceland lies on the Mid Atlantic Ridge (MAR), which is expressed by ~50 km wide

axial rift zones that are the loci of active faulting and volcanism (Gudmundsson, 2000;

Roberts et al., 2007). The major rift zones are referred to as the Western (WVZ),

Northern (NVZ) and Eastern (EVZ) Volcanic Zones and contain 30 volcanic systems characterized by extensive fissure swarms as well as central volcanoes and other localized eruptive centers (Fig. 1). The full spreading rate between the North American

Plate and the Eurasian Plate is 18-20 mm/yr (Geirsson et al., 2006), but spreading with associated rifting and magma injection is not continuous in time or space (Thordarson and Larsen, 2007). There are distinct rifting episodes characterized by earthquake swarms and volcanic eruptions from a central volcano and along the associated fissure swarm, which together form a single volcanic system. Some of these episodes (e.g., the 1783–

1784 Laki eruption) were characterized by eruption of large volumes of basalt magma

that produced widespread atmospheric pollution events (Thordarson et al., 1996). The last

major rifting episode in the NVZ occurred at Krafla between 1975 and 1984 and

produced only ~1 km3 of basaltic magma.

Figure 17: Map of Central Iceland. Inset - Map of Iceland with the Northern Volcanic Zone (NVZ), Western Volcanic Zone (WVZ), and the Eastern Volcanic Zone (EVZ) (marked with diagonal lines), and V-Vatnajökull. Shaded square shows the region represented by the larger map. Larger map - Location of Kverkfjöll and Upptyppingar. Light grey shaded area - Vatnajökull ice sheet. Dark grey shaded areas - volcanic centres (note that Bárðarbunga is not exposed above the icecap, so its location is marked ‘x’). Grey lines - fissure zones. Black lines with hash marks-faults. The bold black line through Upptyppingar represents the path of migration of epicentres which are shown in detail in Figure 3.

Volcanic Plumbing Systems on Iceland

Considerable progress has been made in documenting the processes involved in the petrogenesis of Icelandic magmas (Sigmarsson and Steinthorsson, 2007). However, there is limited knowledge of the presence, size, and location of magma chambers, even though this information is key to understanding magma evolution, magma dynamics, and eruption styles. The crust in Iceland is thicker (20-40 km) than normal oceanic crust (~7 km) (Darbyshire et al., 2000; Foulger et al., 2003; Kaban et al., 2003; Bown and White,

2004; Leftwich et al., 2005) due to anomalous melt production in the underlying mantle

(White and McKenzie, 1995) resulting from the juxtaposition of a mantle plume and the

Mid-Atlantic Ridge (MAR). This suggests that magma plumbing systems could be more complex than those beneath normal mid-ocean ridge segments, and the results of geodetic, geophysical, and petrologic studies support this suggestion (Sturkell et al.,

2006; Kelley and Barton, 2008).

The location and size of magma reservoirs that feed eruptions from specific volcanic systems in Iceland are not known with certainty. An anomalous swarm of tectonic earthquakes that began in 2007 near Upptyppingar in the Kverkfjöll volcanic system is thought to reflect flow of magma into the crust at depths of 15 to 18 km

(Roberts et al., 2007), and the seismic activity can be interpreted to indicate the presence of a magma injection at this depth. I used a petrologic method to calculate the pressures of crystallization of magmas erupted at Kverkfjöll and Upptypingar to obtain an

independent estimate of the depth of magma chambers beneath this system. The rationale for this work is that combined results of geophysical and petrologic studies should provide a reliable and accurate estimate of the depth of the main magma chamber beneath the Kverkfjöll volcanic system, and allow the potential effects of eruption of magma from this chamber on the climate system to be assessed.

The 100-150 km long Kverkfjöll volcanic system occurs along the easternmost margin of the NVZ (Fig. 1). The Kverkfjöll central volcano is located at the southern end of the system at the northern edge of the Vatnajökull ice sheet, and smaller eruptive centres occur along the fissure system. Upptyppingar is one of these centers and is located ~50 km north of the Kverkfjöll volcano. Eruption products in this system include lava flows, pillow lavas, hyaloclastites, scoria, and spatter (Hansen and Gronvold, 2000;

Hoskuldson et al., 2006). Many of the hyaloclastites, pillow lavas, and flows formed in eruptions in the last Ice Age. However, the Lindahraun lava field formed <2,800 years ago, and five small eruptions have occurred in historic times (from 1655 to 1968).

Pressures of crystallization were calculated using a method based on the experiments and approach of Yang et al. (1996). The pressure can be found at which a particular liquid was multiply saturated with olivine, plagioclase, and augite prior to eruption (Kelley and Barton, 2008). The liquid compositions are converted to normative mineral components (Grove et al., 1993), and projected from plagioclase onto the plane olivine-augite-quartz and from olivine onto the plane plagioclase-augite-quartz. The pressure dependence of each normative mineral component in the predicted liquids is

67 determined by regression, and the pressure of crystallization is determined from the regression equations using the projected normative mineral components for the original sample. Other workers have developed similar methods to determine the pressures of partial crystallization from melt compositions, and have used these to constrain the evolutionary history of mid-ocean ridge basalts (Herzberg, 2004; Villiger et al., 2007).

The method of Herzburg (2004) has been used here to calculate pressures for the purposes of comparison. The pressures calculated using the method of Herzberg are consistently lower than those using the method of Kelley and Barton (2008). The reason for this is not yet clear and is the focus of ongoing studies. Comparison of the method used here with published experimental data indicates that calculated pressures are accurate to ±116 MPa and precise to ±50 MPa (uncertainties given at the 1σ level)

(Kelley and Barton, 2008). It has been shown that pressures calculated with this method yield estimates of the depths of magma chambers beneath several volcanoes in Iceland

(Askja, Grímsvötn, Hengill, Hekla) that are consistent with those inferred from geodetic and geophysical studies (Kelley and Barton, 2008).

Table 4. Summary of all basaltic glass data from Upptyppingar and Kverkfjoll.

Sample ID Locality SiO2 Al2O3 TiO2 FeO1 MnO MgO CaO Na2O K2O P2O5 TOTAL Mg# FeO/MgO CaO/Al2O3 P K&B 1σ Z Km 1σ T1σ P Herz 51 Upp 49.46 12.91 3.38 15.05 0.24 4.80 9.85 2.73 0.72 0.40 99.54 0.2418 3.1352 0.7633 4.18 0.70 14.72 0.25 1171.59 4.09 1.40 54 Upp 46.31 14.89 2.04 11.33 0.20 7.08 11.22 2.42 0.42 0.23 96.14 0.3846 1.6000 0.7536 6.92 0.52 24.36 0.18 1209.03 3.05 4.42 57 Upp 48.98 15.01 2.10 11.80 0.19 7.14 11.32 2.45 0.46 0.21 99.67 0.3770 1.6525 0.7536 5.21 0.34 18.34 0.12 1195.64 1.97 3.52 62 Upp 48.38 15.49 2.16 11.94 0.21 7.57 12.12 2.00 0.37 0.17 100.43 0.3880 1.5777 0.7825 6.59 0.32 23.20 0.11 1212.29 1.86 4.41 63 Upp 47.58 15.93 2.10 11.74 0.16 7.52 11.56 2.35 0.39 0.19 99.53 0.3904 1.5612 0.7260 8.60 0.32 30.27 0.11 1221.45 1.89 6.13 64 Upp 47.55 15.86 2.13 11.75 0.19 7.50 11.57 2.29 0.40 0.20 99.44 0.3898 1.5654 0.7300 8.53 0.37 30.02 0.13 1221.79 2.18 6.03 66 Upp 47.50 15.90 2.15 11.82 0.19 7.49 11.59 2.35 0.39 0.19 99.57 0.3880 1.5770 0.7287 8.73 0.38 30.72 0.13 1222.75 2.23 6.15 67 Upp 47.39 15.87 2.12 11.86 0.19 7.53 11.59 2.33 0.38 0.21 99.47 0.3883 1.5753 0.7304 8.96 0.44 31.53 0.15 1223.51 2.57 6.15 69 Upp 47.50 15.60 2.26 11.89 0.18 7.36 11.58 2.35 0.40 0.22 99.34 0.3824 1.6149 0.7428 8.05 0.41 28.33 0.14 1219.02 2.40 5.51 71 Upp 47.91 15.57 2.14 11.70 0.20 7.53 11.86 2.29 0.38 0.19 99.76 0.3918 1.5526 0.7615 7.14 0.38 25.12 0.13 1214.17 2.23 4.87 a Kverk 49.91 12.79 3.41 14.88 0.25 4.55 8.94 2.36 0.63 0.46 98.18 0.3528 3.2703 0.6990 4.92 0.50 17.30 1.88 1128.71 1.62 2.94 b Kverk 48.09 12.97 3.72 15.36 0.31 4.83 9.34 2.84 0.69 0.72 98.87 0.3592 3.1801 0.7201 8.10 1.50 28.54 5.16 1121.19 4.50 3.19 c Kverk 49.5 13.1 3.57 15.53 0.31 4.81 9.24 2.76 0.67 0.83 100.32 0.3557 3.2287 0.7053 7.95 1.60 27.98 5.56 1125.59 4.83 3.12 d Kverk 49.34 12.98 3.51 15.79 0.34 4.77 8.93 2.83 0.67 0.83 99.99 0.3500 3.3103 0.6880 8.59 1.70 30.24 5.97 1122.10 5.20 3.29

69 e Kverk 49.31 12.84 3.45 14.32 0.26 4.81 9.41 2.37 0.72 0.52 98.01 0.3745 2.9771 0.7329 4.40 0.70 15.62 2.55 1131.51 2.19 2.32 f Kverk 50.1 12.4 3.51 15.3 0.27 4.71 9.95 2.75 0.65 0.45 100.09 0.3543 3.2484 0.8024 3.10 0.80 10.92 2.91 1116.29 2.52 0.37 g Kverk 49.63 12.02 3.48 14.37 0.04 4.54 9.99 2.75 0.69 0.47 97.98 0.3603 3.1652 0.8311 0.90 0.50 3.15 1.93 1116.01 1.67 ‐0.70 h Kverk 51.87 13.58 2.11 13.53 0.19 6.33 10.3 1.26 0.43 0.21 99.81 0.4547 2.1374 0.7585 3.41 0.20 12.00 0.63 1170.63 0.53 2.35 i Kverk 50.66 13.46 2.52 13.68 0.19 6.7 10.58 1.26 0.48 0.29 99.82 0.4661 2.0418 0.7860 4.30 0.40 15.02 1.48 1169.59 1.25 2.45

Average 48.79 14.17 2.73 13.35 0.22 6.19 10.58 2.35 0.52 0.37 0.38 2.31 0.75 6.24 21.97 1174.36 3.58 Max 51.87 15.93 3.72 15.79 0.34 7.57 12.12 2.84 0.72 0.83 0.47 3.31 0.83 8.96 31.53 1223.51 6.15 Min 46.31 12.02 2.04 11.33 0.04 4.54 8.93 1.26 0.37 0.17 0.24 1.55 0.69 0.90 3.15 1116.01 ‐0.70

Results MC2 50.03 13.43 2.83 13.88 0.22 5.72 10.07 2.07 0.57 0.35 0.38 2.54 0.75 4.40 0.47 15.50 1.15 1161.28 1.94 2.50 LC3 47.91 14.92 2.54 12.79 0.23 6.73 10.97 2.44 0.47 0.36 0.38 2.03 0.73 8.02 0.72 28.21 1.62 1192.08 2.99 4.84

Samples 51‐71 are from Mt. Upptyppingar and were collected for this study. Samples a‐f are from Kverkfjoll and are from Höskuldsson et al. (2006); samples g‐i from Hansen and Grönvold (2000). 1All iron as FeO; 2Samples a,e,h,i,51,57; 3Samples b,c,d,54,62,63,64,66,67,69,71; a ‐ Karlshryggur; b ‐ Lindarhrygger; c‐ Vegarskard; d ‐ Karlsrani; e ‐ Kreppuhrygger; f ‐ Virkisfell; g ‐ Matrix Glass; Sample NAL455; h,i ‐ Equilibrium compositions of melt inclusions in plagioclase macrocrysts, Sample NAL455; MC – middle crust; LC – lower crust. Columns labeled a‐i show compositional data and pressure depth and temperature of sourcing magma chamber. Columns labeled MC and LC show average compositions, pressures, depths, and temperatures for the mid‐crustal magma chamber and the lower‐crustal magma

chamber. Liquid compositions are preserved as glasses in the matrix of Icelandic lavas, hyaloclastites, and tephra, and also as glass inclusions that represent liquids trapped in minerals during crystallization. Glass analyses are preferable to whole-rock analyses for calculating crystallization pressures, because glasses represent samples of quenched melts. Analyses of glasses from Kverkfjöll have been compiled from published papers

(Hansen and Gronvold, 2000; Hoskuldsson et al., 2006) and are listed in Table 1.

Analyses of melt inclusions with compositions that appear to have been modified by post-entrapment crystallization and diffusive re-equilibration with the host mineral

(Danyushevsky et al., 2002) were excluded from this study. Also included in Table 1 are glass analyses of pillow basalts that were collected from Mt. Upptyppingar in August,

2008. The locations of sample collection are shown in Figure 18. Samples were analyzed by Michael Garcia at the University of Hawaii by electron microprobe analysis.

Analyses listed in Table 1 are average compositions from 5 spots per sample.

70 69 68 60 62 63 58 61 64 71 57 65 54 6766 5051 52

Figure 18: Field work at Mt. Upptyppingar. In left panel is a topographic map of Mt. Upptyppingar with localities of collected samples shown by blue dots. On left side of figure are from upper left and moving clockwise typical exposure of pillow basalts, basaltic glass in hyaloclastites ridge, hyaloclastites ridge outcrops from a distance, and pillow basalt outcrop.

Chemical Variations

The Kverkfjöll magmas are evolved (relatively rich in FeO, low Mg#) tholeiites.

Most of the samples show chemical variations, such as decrease in CaO and CaO/Al2O3 with decreasing MgO (Fig. 3), are consistent with magma evolution in response to crystallization of olivine, plagioclase, and augite. Therefore, most samples have compositions appropriate for calculating the pressures of multiple saturation of liquids with these mineral phases. However, two samples from Kverkfjöll, and one from

Upptyppingar, have anomalously high CaO and CaO/Al2O3 and it is not clear that these

samples represent melts that were saturated with olivine, plagioclase, and augite prior to

eruption. Accordingly, pressures calculated for these samples must be interpreted with

caution. Also, relatively high pressures are found for three samples from Kverkfjöll that

have low MgO contents. The errors on the calculated pressures are high and these

pressures are deemed unreliable (See Figure 19).

Figure 19: Chemical variations for Kverkfjöll glasses, and pressures and depths of partial crystallization of Kverkfjöll magmas. The two diagrams are plots of weight % CaO vs. MgO and CaO/Al2O3 vs. MgO and are used to identify samples formed by partial crystallization of the assemblage olivine-plagioclase-augite. Symbols: ● Reliable Upptyppingar samples, ○ Unreliable Upptyppingar samples, ■ Reliable Kverkfjöll samples, □ Unreliable Kverkfjöll samples. Reliable samples are those that define a trend of decreasing CaO and CaO/Al2O3 with decreasing MgO (arrows) that have compositions consistent with multiple saturation with olivine, plagioclase, and augite, consistent with variation for most other Icelandic volcanic centers (Kelley and Barton, 2008). Pressures calculated for these samples yield reliable estimates of the pressure of cotectic crystallization and can be used to calculate the depths to the magma chambers (assuming a crustal density of 2900 kg m-3). Unreliable samples are those with anomalous CaO and CaO/Al2O3. It is not clear based on either geochemistry or petrology that these samples were multiply saturated with olivine, plagioclase, and augite, and calculated pressures are not considered to provide reliable estimates of the pressure of cotectic crystallization (see text for discussion). The lower two plots show the calculated pressures and corresponding depths of cotectic crystallization.

Figure 20. Samples represented by circles indicate partial crystallization at P=470-521 MPa and P=659-896 MPa. These pressures correspond to depths of 11-19 km, and 23-32 km. Symbols: ● Reliable Upptyppingar samples, ○ Unreliable Upptyppingar samples, ■ Reliable Kverkfjöll samples, □ Unreliable Kverkfjöll samples, ○ Upptyppingar samples calculated using Herzberg method, □ Kverkfjöll samples calculated using Herzberg method. Definitions of reliable and unreliable here are given in Figure 19 caption. The grey shaded area indicates the average pressures or depths of crystallization for reliable samples with uncertainties calculated at the 1σ level. These give the most likely pressure or depth interval of cotectic crystallization, and indicate the most likely depths of magma chamber (16±2.4). The maximum pressure gives an estimate of crustal thickness (31.53 ± 0.15 km) and is shown by a grey dashed line. Grey arrow shows path of magma during ascent and storage.

RESULTS

Calculated pressures are listed in Table 4 and plotted in Fig. 20. The unfiltered results suggest that the Kverkfjöll system magmas crystallized over a relatively wide

range of pressure, from 896 to 90 MPa (average 508±89 MPa). However, the results for

samples that represent olivine-plagioclase-augite saturated melts indicate partial

crystallization over two distinct pressure ranges, one at 470-521 MPa and the other at

659-896 MPa. The uncertainties associated with calculated pressures in the range 341-

492 MPa are uniformly low (average: 46 MPa) indicating that these results are robust.

These results provide strong evidence for partial cotectic crystallization at 445±69 MPa.

The uncertainties associated with calculated pressures in the range 659-896 MPa are also low (average: 43 MPa). The results suggest partial cotectic crystallization at 794±92

MPa. Relatively low pressures (90 and 310 MPa) are calculated for the two compositionally anomalous samples from Kverkfjöll, and there are several possible interpretations of these results (Herzberg, 2004; Kelley and Barton, 2008): the glass analyses are inaccurate or unreliable; the glasses represent melts that crystallized olivine, plagioclase, and augite at higher pressure followed by crystallization of olivine±plagioclase at lower pressures, or; the calculated pressures represent the actual pressures of olivine-plagioclase-augite cotectic crystallization. The available data do not allow discrimination between these possibilities, and the results for the compositionally

anomalous samples are not considered further.

When only the reliable pressure values are considered (the filled symbols in Figure

20), points in MgO vs. P space show a trend that can be interpreted in terms of the likely ascent history of the magma in this plumbing system. The gray arrow in Figure 20 shows that the less evolved magmas ascend from the base of the crsut to the shallower chamber.

After the ascending to ~15 km, the magmas are stored in a mid-crustal magma chamber and evolve to compositions with lower MgO abundances.

Pressures were also calculated for all samples using the method of Herzberg

(2004). These results are plotted in red symbols in Figure 20. As stated above, the pressures found using this calculation method are systematically lower than those using the methods described by Kelley and Barton (2008). The reason for this is a topic of ongoing investigation.

DISCUSSION

It is necessary to consider a limitation inherent in the method – the assumption that the magmas crystallized under anhydrous conditions. Addition of water to anhydrous melts depresses liquidus temperatures and affects the pressure of cotectic crystallization

(Almeev et al., 2008; Asimov et al., 2004; Danyushevsky, 2001; Danyushevsky et al.,

1996). Measured H2O contents of Kverkfjöll glasses (Hoskuldsson et al., 2006) average

0.9±0.07 wt.%. I calculated H2O contents of 0.3±0.23 wt.% using the method of

Danyushevsky (2001) and Danyushevsky et al. (1996). Although low amounts of water

will shift the calculated pressures of cotectic crystallization, the effect is small compared

with the uncertainties in the pressure calculations (Herzberg, 2004; Kelley and Barton,

2008; Almeev et al., 2008). I conclude that the calculated pressures closely match the

actual pressures of partial crystallization of the Kverkfjöll magmas.

The pressures can be used to calculate the depths of crystallization assuming a crustal density of 2900 kg m-3, which is appropriate for crust mostly composed of basalt.

The depths are 12-18 km and 23-31 km, and represent the depth at which the melts were

last saturated with olivine, plagioclase, and augite. Ascending magmas must pause and

crystallize for a sufficiently long period of time to reach multiple-saturation, and this

most likely occurs in magma chambers. Accordingly, the results indicate a magma

chamber at depth of 15.6±2.4 km and crystallization in the lower crust to depths as high

as 31.53 ± 0.15 km (Table 4 and Figure 20). The crustal thickness at Kverkfjöll has been

estimated from seismic and gravity data to be 30 to 35 km (Darbyshire et al. 1998, 2000a;

Allen et al. 2002; Kaban et al., 2002; Leftwich et al. 2005; Fedorova et al. 2005), which

is consistent with the estimate of 30 km for crustal thickness at the nearby Askja volcano

from surface deformation studies (de Zeeuw-van Dalfsen et al., 2005). Therefore, the

results provide evidence for a magma chamber in the middle crust and magma rising

through the lower crust beneath the Kverkfjöll volcanic system. The evidence for a

chamber in the lower crust is not as strong as that for the chamber in the middle crust, because of the MgO – P systematics. However, the presence of a chamber at the crust- mantle boundary (~30km) is consistent with evidence that some magmas erupted in other

Icelandic volcanic systems originated from chambers at the base of the crust (Kelley and

Barton, 2008). There is a caldera complex at Kverkfjöll volcano, and shallow magma chambers occur beneath many volcanic centres with calderas in Iceland (Kelley and

Barton, 2008) (e.g. Askja, Grímsvötn, Katla). However, there is no convincing evidence for a chamber in the shallow crust.

Swarms of micro-seismicity that occurred near Upptyppingar (16.2° W, 65° N) from February to August 2007 provide evidence that magma is being injected into the crust at ~15 km depth. The focal depths of the earthquakes are concentrated at depths between 15 and 18 km, and this has been interpreted to reflect influx of magma into the deep crust (Roberts et al., 2007). However, the seismicity could be caused by inflation of the mid-crustal chamber in response to input of magma from deeper levels in the crust or mantle. Data from the South Iceland Lowland (SIL) seismological data acquisition system seismic network (Icelandic Meteorological Office, Department of Geophysics,

SIL database) indicates that activity is continuing presently. Earthquake hypocenters from 16.0° to 16.32° W, 64.98° to 65.1° N, and February 1, 2007 to May 8, 2008 at magnitudes >0.8 are plotted in Figs. 3 and 4. These are preliminary results acquired from the SIL on-line database and therefore should be considered cautiously. However, some general trends can be seen. The swarms have migrated ~10 km to the northeast since July

2007 (Jakobsdottir et al., 2008) (Fig. 21.), possibly indicating dike intrusion, or propagation of the fault system. Recent analysis of these data by the IMO has led to the interpretation that the seismic events represent the intrusion of a planar magma body dipping southward at 41° at depths of ~4-20 km (Jadobsdottir et al., 2008).

Figure 21: Migration of earthquake epicenters and frequency of total earthquakes earthquakes. a.) Earthquake epicenters beneath Upptyppingar from February 1, 2007 through May 8, 2008 (SIL database). From February-August 2007, there was not much horizontal variation in the location of the earthquake clusters. Since August 2007, however, the clusters have been migrating along a N68E orientation. b.) Frequency of seismic events since 2/2007. Activity is shown to occur in clusters, and there appears to be an increase in activity.

A few swarms originated at depths >18 km (up to 25 km), and might be associated with the inferred chamber at the base of the crust (Fig.4). However, the seismic data do not provide conclusive support for the existence of this chamber. The deep crust in Iceland is relatively hot (Kelley and Barton, 2008), so that influx of magma might be accommodated by ductile deformation of the crust (Sigmundsson, 2007) and might not be accompanied by seismic activity. It should be possible to detect the chamber at the crust- mantle boundary with geodetic techniques (Sturkell et al., 2006; de Zeeuw-van Dalfsen et al., 2005).

Figure 22: Earthquake hypocenters beneath Upptyppingar. Latitude, longitude and depth for all earthquake foci beneath Upptyppingar from February 1, 2007 through May 8, 2008. These data are unfiltered, and include 7560 foci that range in magnitude from 0.17 to 3.35 (SIL database). The majority of hypocenters are clustered at 15 km, although some occur at depths up to 25 km and others occur at shallower depths. The cluster of hypocenters at 15 km is interpreted to mark the location of a magma chamber, and agrees with the depth calculated from glass compositions.

The seismic unrest at Upptyppingar may herald the onset of a discrete rifting

episode along the eastern margin of the NVZ. Our results support a model for spreading accompanied by intrusion of magma into relatively deep crustal reservoirs (≥15 km depth) with subsequent injection of dikes into the upper crust along the plate margin

(Gudmundsson, 2000). In addition, the evidence for magma chambers in the middle and

lower crust beneath the Kverkfjöll system, and beneath other volcanic systems in Iceland

(Kelley and Barton, 2008), strongly suggests that the crust forms from multiple magma

bodies. This model for accretion of thick crust in Iceland is similar to that recently

proposed for formation of thinner crust along the submerged mid-ocean ridge system

(Pan and Batiza, 2003; Nedimović et al., 2005).

Seismic activity has preceded recent eruptions in Iceland (e.g. the Krafla eruptions

between 1975 and 1984, the Gjalp eruption in 1996, the Hekla eruption in 2000, and the

Grímsvötn eruption in 2004). The seismic activity usually begins a few hours to a few months before the onset of eruptions. In contrast, the activity at Upptyppingar was ongoing for more than 14 months. Consequently, there is no certainty that an eruption is

imminent, although this must be considered a distinct possibility.

Our results indicate that if an eruption does occur, the magma will be erupted from

depths of ~15km, and that magma ascent, given the Kverkfjöll system glass data, will be

sufficiently rapid to inhibit significant cooling and crystallization. Accordingly, glass

compositions will preserve high-pressure phase relationships. Rapid ascent will promote

rapid near-surface degassing leading to explosive activity that could inject SO2 into the

troposphere and, possibly, the lower stratosphere as in the 1783 Laki eruption. The latter

3 vented 15.1 km of magma and released ~120 Mt SO2, which formed aerosols that

affected the climate in the northern hemisphere and produced extreme air pollution

(Thordarson et al., 1996; Thordarson et al., 2003). Calculated S concentrations

(1791±173 ppm) from the Kverkfjöll samples are similar to those for Laki magmas

(1600-1700 ppm), indicating that a large volume (≥15 km3) eruption from the Kverkfjöll

system could produce similar quantities of SO2 aerosols and pose a considerable hazard both on a local and a global scale. Measured S concentrations (727±187 ppm) from

Upptyppingar glasses are lower and could indicate some degassing. Careful analysis of the seismic data coupled with detailed studies of surface deformation should provide an estimate of the size of the mid-crustal chamber and hence of the volume of magma in the chamber. This will allow a more realistic evaluation of the potential size and impact of an eruption.

Chapter 4: Density and Seismic Velocity of the Crust in Icelandic Rift Zones

INTRODUCTION

Typical oceanic crust has been described through drilling (Teagle et al., 2005), seismic studies (White et al., 1992), and through study of ophiolites (Korenaga and

Kelleman, 1997). The crust along the spreading axis of mid-ocean ridges is typically ~7

km thick (White et al., 1992). The crust in Iceland is anomalously thick (up to 40 km)

(e.g., Leftwich et al., 2005). This thick crust is the result of magma generation at a mantle

plume that is located on the MAR. The amount of magma produced, and thickness of

crust generated depends on factors such as the radius, volume flux, buoyancy, and

viscosity of the plume, as well as on the spreading rate of the ridge. Various numeric

models have been proposed for the interaction of a plume and ridge (Ito et al., 1996, Ribe

et al., 1995, Jones, 2003) and show that it is not easy to find the balance between these

variables that accurately reproduces observed crustal thicknesses, seismic P-wave

velocities, and geochemical anomaly distributions.

The crust in Iceland was thought early on to consist of three layers similar to

oceanic crust. These layers consist of hydrothermally altered, porous basalt in the upper

crust, sheeted dike complexes in the middle crust and gabbros and cumulates in the

lower crust. Palmason (1964) introduced a model with a Layer 4 at the base of the crust

which had high seismic velocity and density values. Modeling efforts have since tried to

reproduce crustal accretion through different mechanisms of magma injection and

83 crystallization. Problems that exist with these efforts include relating thermal regimes with compositional models in order to reconcile the amount of melt that may be present in the lower crust and consequently affect the seismic observations, and creating seismic and gravity models that agree with one another in terms of the thickness of the crust and the temperature at the base of the crust.

More recently, seismic and gravity models have begun to converge to provide a consistent picture of crustal thickness (Darbshire et al., 2000; Leftwich et al., 2005), but there are still uncertainties primarily associated with the thermal state of the lower crust.

Some interpretations involve a model in which the lower crust (15 to 25 km) is relatively cool and solid (Menke and Levin, 1994; Menke and Sparks, 1995, White et al., 1996), whereas other interpretations yield a model in which the lower crust is relatively warm and possibly partially molten (Palmason, 1971; Beblo and Bjornson, 1980; Flovenz and

Saemundsson, 1993). Few models integrate geophysical models with petrologic and geochemical constraints with the exceptions of Maclennan et al,. (2001), and Farentini et al., (1996).

The objectives of this study are to create a simple integrated (petrological, geophysical, geochemical) model for Icelandic crust in the active rift zones. Models will be created that will serve as a guide to future studies, and will compliment the work of

Maclennan et al. (2001), and Farentani et al. (1996). While this modeling is not comprehensive, it provides a basis for developing more sophisticated models.

GEOLOGIC SETTING AND SUMMARY OF PREVIOUS WORK

The MAR runs approximately NE-SW through Iceland creating a rift system with

abundant volcanic activity. There are three distinct rift zones that are the focus of most

of the volcanic activity on the island: The Western Volcanic Zone (WVZ), the Eastern

Volcanic Zone (EVZ), and the Northern Volcanic Zone (NVZ) (Figure 1). The three

zones collectively are referred to as the neovolcanic zone. The WVZ is the extension of

the Reykjanes Ridge, and marks the trace of the MAR as it runs into Iceland from the

south. The NVZ extends offshore to the north into the Kolbeinsey Ridge.

The NVZ and WVZ are similar in terms of structure and crustal thickness (20

km). In the eastern central part of the island roughly where the three zones meet, the crust is much thicker (40 km) (Staples et al, 1997; Darbyshire et al., 1998; Leftwich et al,

2005). This area, near the northwestern edge of the large Vatnajökull ice sheet, is where the mantle plume is thought to be centered (Leftwich et al., 2005). The volcanic centers of the neovolcanic zone can be grouped together to represent the rift-related volcanism on

Iceland. There are 30 defined volcanic systems in the rift zones. These systems are shown in Figure 1. The numerous eruptions provide abundant samples of fresh lava.

Estimates of crustal thickness in Iceland are based on seismic and gravity studies.

The crust is generally considered to be 20 km thick in the vicinity of the rift system, and up to 40 km in the off-axis and plume-related portions of the island (Staples et al, 1997;

Darbyshire et al., 1998, Maclennan et al., 2001; Leftwich et al, 2005; Kelley and Barton,

2008). The thickest crust in Iceland has been determined to be 40 km by wide-angle

Moho reflections, receiver functions, and surface-wave dispersion (Darbyshire et al.,

1998, Darbyshire et al., 2000, Gudmundsson, 2003).

The temperature of the crust increases rapidly with depth, and is quite variable laterally across the island. Kaban et al. (2002) proposed that the 1200˚C isotherm lies at

30-50 km depth in some parts of the island, but at 20 km or less in the neovolcanic zone.

Palmason (1971), and Flovenz (1992) concluded that the temperature at the base of the crust has to be ~1200˚C. However, Menke and Levin (1994), and Menke (1995) claim that the upper limit for temperature in the lowermost crust is only 700-775˚C.

Seismic velocity models are consistent in different studies of the crust in Iceland.

-1 Values for Vp in the upper crust (0-2 km) vary from less than 4.0 km s in fresh lava

flows to 5.0 km s-1 in more altered and dense basalts (Palmason, 1971; Flovenz, 1980;

Flovenz and Gunnarsson, 1991). These increase to ~6.5 km s-1 at 3-10 km depth and then to 7.0-7.4 km s-1 at 10-30 km depth (Palmason, 1971; Bjarnason et al., 1993; White et

al., 1996; Staples et al., 1997; Brandsdottir et al., 1997; Darbyshire et al., 1998). The lowest layer of the crust is defined seismically and is referred to as Layer 4 (Palmason,

1971). Kaban et al. (2002) estimated that the density in layer 4 must be on the order of

3050-3150 kg m-3. Menke (1999) contended that the mantle density beneath Iceland is as low as 3150±60 kg m-3, and that the density contrast across the crust-mantle transition is

90±10 kg m-3.

AVERAGE COMPOSITION OF THE CRUST

Volcanic Glass analyses

A database of over 800 analyses of Icelandic volcanic glasses was compiled from

published papers. These data are discussed in detail by Kelley and Barton (2008). These

glasses were taken from volcanic centers located in the active rift zone at the locations

shown in Figure 1. No picritic glasses (MgO contents higher than 12%) are found in

published analyses. This means that no melts with MgO>12 wt % have erupted on

Iceland.

Whole Rock analysis

A total of 1788 basalt samples from the Georoc database, managed by the Max

Plank Institute for Chemistry, were used to compile a database of whole rock

compositions for crust in the rift zone. Samples were limited to those containing less

than 12% MgO because no glasses from Iceland have MgO higher than 12%. It was

assumed, therefore, that rocks with MgO>12% have been affected by olivine

accumulation and do not represent liquid compositions. Samples with anomalous

compositions due to alteration were eliminated. These were samples with high K2O, low

K2O, high H2O, Fe2O3, high Rb, high Ba, or high Sr. The remaining compositions were then cut down to only those with less than 12% MgO. The data were plotted on variation diagrams (Figure 23) with glass data and those samples that plot off of the main data array were removed.

Figure 23: Variation diagrams of unfiltered whole rock and glass database.

Composition of the Crust

The whole rock and glass analyses indicate that the crust is predominantly basaltic

(Table 5). Silicic magmas have erupted in Iceland, most notably at Torfajokull volcano,

and are thought to have formed by differentiation and remelting (Gunnarsson et al., 1998;

Marsh et al., 1991). The whole rock samples range in SiO2 from 35 to 73 wt.%, and

MgO from 7 to 12 wt.%. This range in whole rock compositions agrees well with that of the glass analyses (Figure 23). The glass analyses show a smaller compositional range than the whole-rock analyses.

Oxide (wt%) SiO2 TiO2 Al2O3 Fe2O3 FeO CaO MgO MnO K2O Na2O P2O5 Total

Most Felsic WR 72.3 0.22 12.1 0.4 1.4 1 0.09 0.05 3.5 4.9 0.04 96.00

Average WR 48.84 2.13 14.43 1.42 10.86 11.13 7.04 0.20 0.41 2.39 0.24 99.10

Most Mafic WR 49.78 1.63 11.78 3.32 6.05 11.99 11.06 0.18 0.69 2.02 0.24 98.74

Calculated Parent Magma 47.05 1.78 12.06 1.19 11.47 9.30 14.45 0.17 0.34 2.00 0.20 100.01 (Ol Fo88) 90 Korenga and Kelleman (2000) 46.5 1.48 13.1 0.7 11.8 9.8 14.4 0.17 0.18 1.79 0 99.92 (Ol Fo88) Most Felsic Glass 49 3.09 13.3 0 13.8 11.2 5.82 0.25 0.54 2.73 0.41 100.14

Average Glass 48.45 1.84 14.57 0.00 11.53 12.35 7.99 0.19 0.21 2.22 0.21 99.55

Most Mafic Glass 48.54 0.61 14.48 0 8.91 12.85 12.07 0.17 0.05 1.5 0.01 99.19

Table 5. Major element compositions for neovolcanic zone crustal whole rocks and glasses, and calculated parent magmas.

CALCULATION OF A 1-D GEOTHERM

The geothermal regime in Iceland is complicated by many factors, including the

presence of shallow partially molten magma lenses in the crust beneath some localities,

cooling intrusive bodies, mechanical energy released by crustal movement, radioactive

heat generation, latent heat of phase changes, the thermal effects of alteration processes,

and the redistribution of heat by groundwater circulation. To create a model of the heat

flow regime for all of Iceland, the effects of each of these processes needs to be evaluated. Furthermore, a great deal of lateral and temporal variation occurs in most if not all of these factors across Iceland. Creation of a model for near surface distribution of heat is beyond the scope of this study. My goal is to build a simple crustal model that is firmly based on compositional data. To simplify matters, all of the rift zones in Iceland were assumed to be fairly uniform lithologically, the crustal composition laterally homogenous, the structure along the spreading center uniform, and the crustal thickness constant. The magma plumbing structure and temperatures reported in Chapter 1 along with the compositions discussed above show that the rift zones are sufficiently uniform to

calculate a representative geothermal gradient.

Here I consider the crust in the rift zones to be 20 km thick (Darbyshire et al.,

1998; Allen et al., 2002; Kelley and Barton, 2008), and that hydrothermal circulation

occurs in the upper 3 km of the crust facilitates (Flovenz, 1993). The temperature at the

base of the crust is 1200° C (Kelley and Barton, 2008). The temperature at the surface is

taken to be 25° C.

The geothermal gradient in the upper 3 km must be considered differently from

that in the lower 17 km of the crust because of the presence of hydrothermal circulation.

In order for hydrothermal circulation of any fluid to take place, a critical Raleigh number

of 4π2 must be exceeded. The the upper few kilometers of the rift zones are very porous

rock (Flovenz and Saemundsson, 1993) and relatively high temperatures throughout the crust, so there is a sufficiently high Raleigh number to allow hydrothermal circulation. In

order then to determine how temperature varies with depth in this zone of hydrothermal

circulation, the equation

2 Ra = [αfgρf cpfkb(T1-T0)]/μλm Eq. 1

is used where k is the thermal conductivity of water, and b is the thickness of the section

2 3 (3 km). By using the values of Ra = 4π , density of water (ρf) = 1000 kg/m for, thermal

-3 -1 -4 expansivity of water (αf) = 10 K for, viscosity (μ) = 1.33 x 10 Pa s, heat capacity of

3 -1 -1 -1 -1 water (cpf) = 4.2 x 10 JKg K , thermal conductivity of the rock (λm) = 3.3 W m K , and gravity (g) = 9.8 m s-1, Equation 1 can be rewritten as

dT/dy = (4.2 x 10-10)/kb2 Eq. 2

Flovenz and Saemundsson (1993) gave the value kb = 2x1012 for the rift zones. Inserting

this value into Eq. 3 gives a value of dt/dy=71.2 °C/km for the upper 3 k m of the crust in

the rift zones.

Below 3 km depth, the porosity of the rocks no longer allows for hydrothermal

circulation (Flovenz and Saemundsson, 1993) and therefore thermal conduction controls

the flow of heat as long as the crust is completely solid. This conduction of heat is

controlled by the equation (Turcotte and Schubert, 2002)

2 T = To + (qo/k)(z – zo) – (ρH/2k)(z – zo) Eq. 3

where z is crustal thickness, ρ is density, q is heat flow, k is thermal conductivity, and H

is internal heat generation.

Thermal conductivity, k, is more or less constant for rocks in oceanic crust

because they consist of relatively homogenous basalt. There are other rock types in

Icelandic crust, but they form a volumetrically minor crustal component so they will be

removed from consideration here. An average value of 1.9 W/m/K for basaltic (Drury,

1985; Flovenz and Saemundsson, 1993) is appropriate. Variation in this value due to variation in rock composition and mineralogy will be considered below. Porosity has the

effect of decreasing the thermal conductivity in basaltic rock. For rocks with a porosity

of 0 to 10%, which is appropriate for the Icelandic rift zones (Stefansson, 1997) the

thermal conductivity ranges from 1.6 to 1.9 W/m/K (Stefansson, 1997; Clauser and

Huenges, 1995). Here, I follow the suggestion of Flovenz and Saemundsson (1993) that

there is no porosity below 3 km and therefore, a constant value of 1.9 W/m/K is

appropriate.

Internal heat generation, ρH is considered constant with depth at 0.3 Wm-3

(Rybach, 1988). Small changes may occur with depth but they have no significant effect on the geothermal gradient.

The surface heat flow varies laterally because of magma intrusion and hydrothermal activity, but similar rock types, and heat source, along the rift zones permits the use of an average value to find a representative geothermal gradient. Kononov and

Polyak (1975) reported the heat flow to be 8 μcal/cm2/sec (328 mW/m2) in the EVZ, 7

μcal/cm2/sec (287 mW/m2) in the NVZ, and 5 to 7 μcal/cm2/sec (205 to 287 mW/m2) in

the WVZ. Friedman et al. (1976) suggested a value of 780 μcal/m2/sec (~320 mW/m2)

for thermal modeling. Values of 1.2 to 1.6 HFU (~50 to 66 mW/m2) are reported for

background heat flow in the volcanic zones by Kononov and Polyak (1978). Values

range from ~90 mWm-2 outside of the rift zones, approaching 200 to 300 mWm-2 in the rift zone with anomalous areas reaching values of more than 400 mWm-2 according to

Flovenz and Saemundsson (1993). A heat flow value of 150 mWm-2 is used here to calculate the geotherm in the rift zones.

Geotherms for Models

Two different models for the crust have been considered and are discussed in detail in the following section. The possibility of mineralogic change due to

metamorphism, and the possibility of compositional change due to magma fractionation

are considered. Because these two models for the crust are different, the calculated

geothermal gradient for each model will be different because of differences in physical

properties with depth (namely the thermal conductivity). I have calculated a geotherm to

be used in Model 2 (mineralogic changes on a constant composition) and in Model 3

(compositional changes). These are shown in Figure 24 and are very similar.

For Model 2, the values of thermal conductivity changed due to the changing

mineralogy represented by basalt (1.95 Wm-1K-1) in the upper 5 km, greenschist (3.0

Wm-1K-1) from 5-9 km, eninsula es (2.46 Wm-1K-1) from 9-12 km, and granulite

(2.475 Wm-1K-1) from 12-20 km (Zoth and Haenel, 1988; Ray et al., 2006).

In the calculation of the geotherm for Model 3, the thermal conductivity changes

because it is assumed that basalt occurs in the upper 10 km of the crust, whereas gabbro

with increasing amounts of pyroxinite and ultramafic cumulates occurs in most of the lower crust with the lowermost 2 km of the crust consisting entirely of ultramafic cumulates. Because magma chambers tend to exist at ~10 km depth beneath volcanic centers in the rift zones (Kelley and Barton, 2008), I have limited all cumulate to be beneath this depth. The values of thermal conductivity used for this geothermal gradient calculation are: basalt: 1.95 Wm-1K-1, gabbro: 2.63 Wm-1K-1, cululates: 4.0 Wm-1K-1

(Zoth and Haenel; 1988).

The geothermal gradients for Model 2 and Model 3 are shown in Figure 24. In

both models, the upper 3 km are treated as discussed above to include the effects of

hydrothermal circulation. Below 3 km, the two gradients are similar, although there are minor differences. These models lead to temperatures of 1185 °C (Model 2) and 1194 °C

(Model 3) at the base of the crust (20 km).

Figure 24: Geothermal gradients for the crustal models.

APPROACH AND RATIONALE

Predicted Variation in Density with Depth

The composition of a rock can be cast into a mineral assemblage by calculating

the normative mineralogy. The CIPW norm is a method for finding the mineralogy of a

rock based on its major element composition. This method was devised by Cross,

Iddings, Pirsson, and Washington (Johannsen, 1931). Once the normative mineralogy is

calculated, the physical properties of the rock as a whole can be determined. The

densities of individual minerals are well known through direct measurement and theory.

Therefore, the density of a rock is the addition of proportional amounts of the densities of

the minerals that it is made of. Changes in the major element composition of a rock lead

to changes in the calculated mineral assemblage (the norm) and therefore changes in the

density.

The density of the crust is generally assumed to increase with depth. Another

arrangement would be unstable due to the tendency of higher density materials wanting

to drop through lower density materials due to gravity. Likewise, the density of the mantle is greater than that of the crust. The density of the mantle is approximated due to the mineral assemblages that have been observed to exist in mantle xenoliths and determined to exist using petrologic theory. For the purposes of this study the density of

the mantle will be considered to be 3300 kg/m3 because that is what is commonly used in

geodynamic modeling efforts and is a generally accepted value (e.g., Philpotts and Ague,

2009). Often the density contrast between the lower crust and the mantle is more

important than actual density values because it is the difference in densities that gives rise to gravity anomalies.

This density contrast is important in the inverse modeling of gravity anomaly data. Figure 25 shows a generalized model of the Icelandic crust in cross section perpendicular to the rift system. This model shows the mantle rising to ~20 km depth at the spreading center with the crust thickening to about 40 km in the off-rift sections of the

island. The density of the mantle was kept constant at 3300 kg m-3, while the density of

the crust was varied to explore the effects of density contrast on calculated gravity

anomalies. The density of the crust is first taken as 3080 kgm-3 which gives a crust-

mantle difference of 220 kg/m3. This produces a positive gravity anomaly of 22 mGals

(Figure 25) which is consistent with values obtained for the neovolcanic zone in many

gravity anomaly studies of Iceland (Darbyshire et al., 2000; Leftwich et al., 2005;

Fedorova et al., 2005).

Lowering the density of the crust to 2670 kg m-3, which approximates the density

of granite and is sometimes erroneously used in modeling of the crust regardless of the

geology, dramatically increases the anomaly to 80 mGals (Figure 25). Clearly, the

gravity data of the rift zone do not support this lower crustal density and this is

inconsistent with the crustal composition in Iceland. Increasing the crustal density to

2900 kg/m-3 (roughly the density of basaltic surface rocks in Iceland) reduces the calculated gravity anomaly to 42 mGals, which also is higher than gravity observations over the rift zone.

Figure 25: Calculated gravity Anomalies for Icelandic Rift Zone.

The gravity modeling efforts illustrate the importance of effectively density

variations. In the case of the Icelandic neovolcanic zone, values for crustal density can

be estimated from petrologic and geochemical studies for effective gravity modeling of

the crust. It is clear from these models that a crustal density of 3080 kg m-3 provides a density contrast with the mantle that yields a reasonable gravity anomaly. The values of

3120 and 3134 kg/m3 found for Model 2 and Model 3 here are even higher values for the

lower crust and would therefore give an even lower gravity anomaly. The values given

here produce a crust-mantle density difference that is close to 200 kg/m3, which is the

value that produces gravity anomalies similar to those that have been measured in the

region. This is in contrast to the suggestions of Menke (1999) that the crust-mantle

density difference must be ~90 kg/m3. It is important when studying the lower crust that

a model that is consistent among all available data is utilized. That is what has been done

in this study.

If the normative mineralogy of a rock is calculated and the density is determined,

this density only applies to surface temperature and pressure conditions. The density of

each mineral is affected by pressure and temperature changes. An increase in

temperature has the effect of decreasing the density of a mineral due to an increase in

volume because of thermal expansion. An increase in pressure has the effect of

increasing the density of a mineral because the volume is reduced.

A problem arises in modeling the crust for the purposes of geophysical and

geodynamic studies. The crust in Iceland commonly taken to be basaltic in composition

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and therefore a normative mineralogy is calculated for the average composition of a

basalt. From here on I will refer to this situation (constant basaltic composition and mineralogy throughout the depth of the crust) as Model 1. When increases in pressure

and temperature are applied according to the geotherm discussed above, the minerals

included in the norm are affected. The resulting density variation with depth for Model 1

is shown in Figure 26. The density of the rock in Model 1 decreases with depth. The

high geothermal gradient in Iceland means that increasing temperature has more affect on

the minerals than the increasing pressure in the depth range of the crust. Therefore,

thermal expansion affects the minerals and the density decreases. As mentioned above

this is an inherently instable situation.

101

Figure 26: Density change with depth in the crust assuming constant composition and basalt mineralogy throughout.

102

PREDICTED VARIATION IN SEISMIC VELOCITY WITH DEPTH

Seismic velocities can be calculated as a function of depth along the geothermal

gradient using calculated density variations. Seismic velocity was calculated using the

methods described in Hacker and Abers (2002). In this method the isothermal bulk

modulus KT(T,P) at elevated pressure and temperature is given as

5/2 KT(T,P) = KT(T){1-f(5-3K΄)}(1+2f) Eq. 4

where

K΄ = (dKT/dP)T Eq. 5

and f is the finite Eulerian strain. The isentropic bulk modulus KS is

KS = KT(T,P)[1+Tγthα(T,P)] Eq. 6

where γth is the first Grüneisen parameter, and α(T,P) is the expansivity at elevated

pressure and temperature which is given by

−δ α(T, P) = α(T)[ρ(P)/ρo] T Eq. 7

The shear modulus at elevated pressure and temperature G(T,P) is found by

5/2 G(T,P) = G(T)(1+2f) {1-f[5-3G΄KT(T)/G(T)]} Eq. 8

P wave velocity VP is then found by

VP = S + GK /)3/4( ρ Eq. 9

and the shear wave velocity VS is found by

VS = G / ρ Eq. 10

Again, I first consider Model 1, the case where the entire crust consists of basalt

103 of average composition and the normative mineralogy at surface conditions. Calculating the expected Vp along the geothermal gradient found above produces values that decrease with depth. The results are shown in Figure 27. This situation is not correct. Seismic studies have always shown that P-wave velocities increase with depth in the crust with the highest values at the base of the crust (e.g., Farnetani et al., 1996).

Figure 27: P-wave change with depth in the crust assuming constant composition and basalt mineralogy throughout.

104

Summary of Results for Model 1

This model fails for both prediction of density and seismic velocity. In order to calculate densities and seismic velocities that increase with depth, either the mineralogy must change with depth to cause these changes in physical properties, or the composition of the crust must vary with depth, or some combination of these two possibilities. The possibilities are not mutually exclusive. A change in composition gives rise to a change in mineralogy. However, it is useful to separate these effects and independently study the magnitude of each one on density and seismic velocity. Changing mineralogy in the crust would be due to the effects of metamorphism. As temperature increases with depth in the Iceland crust, the mineral assemblages that are stable change. Changing composition in the crust would be the result of intracrustal differentiation. This would likely follow the crustal accretion models discussed in Chapter 1. There is evidence from geophysical studies of seismic discontinuities within the crust (e.g., Palmason 1971).

These could result from either mineral or composition changes. The effects of these changes are explored in the following sections.

The geotherms are high in the models presented here, so the application of this problem to MORB is significant. In the mid-ocean ridges, the crust is thinner and therefore the geothermal gradient will be even higher. The geotherm estimates for the

Iceland rift zones are robust. It is hard to justify changing any of the parameters. The crustal thickness is well established by gravity, seismic, and petrologic work. The geotherm is also consistent with heat flow studies and petrologic models

105

METAMORPHISM OF ICELAND CRUST

In this model (Model 2), all changes in lithology encountered with depth are

assumed to be changes in mineralogy resulting from metamorphism at constant

composition. The composition of the melts coming into the crust is that of the lavas that

have been erupted from the base of the crust. Those melts were identified in Chapter 1.

Melts that have been erupted from shallower depths in the crust are similar in

composition to the latter (see Table 5). Therefore, it is practical to model changes in the

crust considering a constant chemical composition.

Mineral modeling

The crux of Model 2 is that the mineral assemblages change with depth.

Knowledge of the bulk composition of the crust (Table 5) and the geothermal gradient

(Figure 24) allow the stable mineralogy to be determined. The Perple_X software

developed by James Connolly was used to calculate the mineral assemblages in Figure 28

and Table 6. These are assemblages common for mafic compositions metamorphosed to

the greenschist, amphibolite, and granulite facies (e.g. Spear, 1993). However, the

assemblages produced are dependent on H2O and CO2 content of the initial composition.

We varied the H2O content from 0% to 5%. CO2 in amounts of 0.5-1% also produced carbonates in the stable assemblages in shallow depths. The calculations were run with

varying amounts of H2O present in the rocks to investigate the effect that this had on the

106

stability of mineral phases. This had little effect on the resulting assemblages at depths below ~4 km. With low amounts of H2O and CO2 (0.5% of each) present in the rock, the

calculated mineral assemblage is such that the density profile is reliable. Physical

properties in the uppermost crust are difficult to reproduce because there is porosity and

water present in the rocks. The purpose of this study is to determine physical property

values in the middle and lower crust for use in geophysical analysis of the crust, so these

problems were not investigated in detail. It is pointed out here because the density and

seismic velocities that I have calculated do not match those of other workers in the upper

5 km of the crust.

Predicted variations in density and seismic velocity with depth

In Model 2, the geothermal gradient was closely approximated using the third

order polynomial

T = 3E-09P3 – 3E-05P2 + 0.2675P + 300.95 Eq. 11

where P is pressure in bars, and T is temperature in K. This polynomial expression was

applied to the calculated phase diagram using in the Perple_X software package and the

mineralogy was found at all depths in the crust (Figure 28). The mineral phases that were

included for consideration were olivine, clinopyroxene, orthopyroxene, plagioclase,

chlorite, cummingtonite, amphibole, ilmenite, epidote, spinel, and biotite.

107

1bar 600 bars 1200 bars Mineral wt %vol %Mineral wt %vol % Mineral wt %vol % Pl(h) 24.28 27.71 Pl(h) 12.15 13.77 Pl(h) 12.75 14.46 Amph(DHP) 27.77 26.66 Amph(DHP) 49.26 47.56 Amph(DHP) 49.14 47.54 Amph(DHP) 10.37 10.16 Cc(AE) 1.23 1.03 Cc(AE) 1.19 1.04 Amph(DHP) 9.45 9.91 Opx(HP) 3.85 3.28 Opx(HP) 3.92 3.34 Cc(AE) 1.26 1.04 Cpx(HP) 13.07 12.19 Cpx(HP) 13.18 12.21 Opx(HP) 3.79 3.28 spss 0.46 0.34 spss 0.46 0.34 Cpx(HP) 17.17 16.4 ab 5.37 6.42 ab 5.05 6.04 spss 0.46 0.35 mic 1.58 1.94 mic 1.58 1.93 mic 1.58 1.95 q 9.17 10.93 q 8.86 10.56 ilm 3.87 2.55 ilm 3.87 2.53 ilm 3.87 2.53

1800 bars 2400 bars 3000 bars Mineral wt %vol %Mineral wt %vol %Mineral wt %vol % Bio(HP) 2.61 2.66 Bio(HP) 2.62 2.62 Bio(HP) 2.62 2.62 Amph(DHP) 23.96 22.86 Amph(DHP) 43.57 41.41 Amph(DHP) 43.58 41.47 Amph(DHP) 19.65 19.14 Pl(h) 34.58 38.8 Pl(h) 34.45 38.65 Pl(h) 12.7 14.22 Opx(HP) 5.25 4.34 Opx(HP) 5.25 4.36 Pl(h) 19.21 22.14 Cpx(HP) 9.22 8.19 Cpx(HP) 9.36 8.32 Cc(AE) 1.17 1.04 spss 0.41 0.3 spss 0.39 0.28 Opx(HP) 4.16 3.48 ilm 3.87 2.46 ilm 3.87 2.46 Cpx(HP) 10.03 9.06 CO2 0.5 1.9 CO2 0.5 1.84 spss 0.43 0.32 q 2.21 2.59 ilm 3.87 2.49

3600 bars 4200 bars 4800 bars Mineral wt %vol%Mineralwt %vol %Mineral wt %vol % Bio(HP) 2.62 2.62 Bio(HP) 2.61 2.61 Bio(HP) 2.6 2.61 Amph(DHP) 43.59 41.52 Amph(DHP) 27.43 26.03 Amph(DHP) 9 8.48 Pl(h) 34.31 38.5 Pl(h) 34.93 39.14 Pl(h) 34.81 39.01 Opx(HP) 5.25 4.39 Opx(HP) 11.56 9.76 Opx(HP) 18.78 15.86 Cpx(HP) 9.51 8.46 Cpx(HP) 18.47 16.42 Cpx(HP) 29.46 26.2 spss 0.35 0.26 spss 0.32 0.23 spss 0.29 0.21 ilm 3.87 2.47 ilm 3.87 2.46 ilm 3.87 2.46 CO2 0.5 1.8 H2O 0.33 1.58 H2O 0.71 3.45 CO2 0.5 1.76 CO2 0.5 1.74

5400 bars 6000 bars Mineral wt % vol% Mineral wt %vol % Bio(HP) 2.61 2.61 Bio(HP) 2.6 2.6 Amph(DHP) 5.48 5.16 Pl(h) 33.47 37.52 Pl(h) 33.45 37.57 Opx(HP) 20.79 17.69 Opx(HP) 20.1 17.04 Cpx(HP) 36.88 32.79 Cpx(HP) 32.97 29.35 spss 0.2 0.14 spss 0.25 0.18 ilm 2.13 1.36 ilm 3.87 2.46 usp 2.55 1.65 H2O 0.78 3.89 H2O 0.89 4.54 CO2 0.5 1.73 CO2 0.5 1.72

Table 6. Abundances of minerals calculated using Perple_X code. Mineral abbreviations are Pl ‐ plagioclase, Amph ‐ amphibole, Cc ‐ calcite, Opx ‐ orthopyroxene, Cpx ‐ clinopyroxene, spss ‐ spessartine garnet, mic ‐ mica, ilm ‐ ilmneite, q ‐ quartz, Bio ‐ biotite, CO2 ‐ CO2 fluid phase, H2O ‐ H2O fluid phase, usp ‐ ulvospinel. References for solid solution mineral data are: (h) ‐ Newton et al., 1980; (DHP) ‐ Dale et al., 2000; (AE) ‐ Anovitz and Essene, 1987; (HP) ‐ Holland and Powell, 1995. 108

Figure 28: One-dimensional cross section of predicted mineral assemblages. For mineral abbreviations and references, see Table 6 caption.

109

The changing mineralogy produced by the Perple_X calculations provides

mineral assemblages that are consistent with basalt in the upper crust (0-3 km), a

greenschist facies assemblage from 3-9 km depth, amphibolite facies assemblage from 9-

12 km and granulite facies assemblage from 12-20 km. The temperature ranges that are

encountered by the geotherm at these depths are coincident with the temperature ranges where the respective facies assemblages are stable (Spear, 1993). Also, Palmason (1986) suggested that the greenschist facies is encountered in the Icelandic rift zones at ~3 km or

~250° C, and the greenschist to amphibolite transition is ~5.6 km or ~600° C.

There is a density increase associated with the amphibolite-granulite facies transition in all models due to the dehydration reaction that occurs there as amphibole becomes unstable. Seismic velocities were not able to be extracted from Perple_X below depths of ~12 km because the code calculated liquid water as a stable phase and so the P- wave velocities went to 0.

The density throughout the crust is 3050 to 3150 kg/m3. A significant density

increase was encountered around 3600 bars. This jump in density represents the

amphibolite-granulite facies transition facilitated by the dehydration of amphiboles.

The densities and seismic velocities calculated for 0 to 4 km depth were not

consistent with observations and other estimates. The difficulty in using Perple_X to

recreate the physical properties in the shallow crust is most likely due to the high amount

of porosity and water that is present in Iceland the shallowest rocks in Iceland’s rift

zones.

110

To attempt to better model the density and seismic velocity in the uppermost 4

km, I changed the mineralogy from that which was the calculated stable mineralogy

according to Perple_X. I used the calculator of Hacker et al. (2005) to find the density

and seismic velocities of the constructed mineralogy. In order to represent the alteration products in the basalts and greenschist facies rocks in the upper 4 km of the crust I added daphnite, epidote, and calcite to the basalt mineralogy. The mineral abundances are given in Table 6. The combination of densities and seismic velocities calculated using the Hacker calculator in the depths 0 to 4 km and Perple_X for depths from 4 to 20 km gives the profiles shown in Figures 29 to 31. This combination gives profiles that closely represent those of other workers including the shallow crust.

Depth (km)01234 aqz 7 7 8 14 12 hAb 14.6 14.6 17.6 16 14.6 lAb00000 an00078 or 2.4 2.4 2.4 2.4 2.4 hb 5 5 20 34 46 daph 40 40 30 4.6 0 ep 015151510 mt11144 cc 30 15 6 3 3 Table 7. Mineral assemblages used in the shallow crust. Appreviations for minerals are: aqz – alpha quartz; hAb – high albite; lAb – low albite; an – anorthite; or – orthoclase; hb – hornblende; daph – daphnite; ep – epidote; mt – magnetite; cc – calcite.

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Figure 29: Density variation through the crust in Model 2.

The density variation in the crust is due to the mineralogy changes. Figure 29 shows the density changes with depth due to facies changes. The crust is made of basalt in the uppermost crust and the density increases with depth due to compaction. Then the greenshist facies minerals produce a density ~3100 kg/m3. This is not a realistic value.

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This high density mineral assemblage cannot be stable at these depths. This problem is

something that I will need to explore more in the future. Next the amphibolite facies

minerals control the density down to 4000 bars where the mineralogy changes to a

granulite facies assemblage. The density at the base of the crust is found to be 3134

kg/m3. This gives a crust-mantle density difference of 166 kg/m3 assuming a mantle

density of 3300 kg/m3. I find this density to be a bit low in comparison to the estimates of some other workers (e.g., Menke, 1999).

The calculated P-wave velocities are shown in Figure 31. The predicted P-wave velocity is 6.95 km/s at the base of the crust and the profile agrees well with that of other workers (Figure 32). Perple_X did not calculate P-wave velocities for assemblages beneath ~2400 bars (~8 km) because it found that CO2 and H2O should be stable in the

fluid phase at these depths and so the value for Vp goes to 0 (Figure 30). Water is

calculated as a stable phase below 4000 bars using the Perple_X code because of the

dehydration reactions associated with the amphibloite-granulite facies transition. Again, the upper 4 km of the profile shown in Figure 25 are calculated using the input

mineralogy listed in Table. The calculator of Hacker et al. (2005) was also used to

calculate P-wave velocities with depth through the entire crust (Figure 31), and so the

problem of stable liquid water was not encountered because the solid phases only were input to the calculator. This P-wave profile is shown in Figure 31.

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Vp 6 6.2 6.4 6.6 6.8 7 0

500

1000 P (bars)P

1500

2000

2500

Figure 30: Variation in predicted P-wave velocity from 0-2400 bars calculated using Perple_X for Model 2.

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Vp (km/s) 6.2 6.4 6.6 6.8 7 7.2 0

10 Depth (km) 12

Figure 31: Variation in predicted P-wave velocity through the crust using the Hacker calculator for Model 2.

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Figure 32: Seismic velocity profiles from this study compared with those of other workers (Figure after Lippitsch et al., 2005).

Whereas the density values in the lower crust are appropriate with the mineral

assemblages considered here, the P-wave profiles are problematic. The results of the calculations done using Perple_X and the Hacker et al. (2003) calculator produce profiles that are roughly similar to observed seismic profiles (Figure 32). That is, the seismic velocities are increasing with the change from greenschist to amphibolite and from amphibolite to granulite depths in the crust. However, within the granulite facies mineral assemblage (the lower crust) the calculated P-wave velocities are decreasing with depth

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due to the temperature increase (see Figure 31). Therefore, the P-wave velocity value

that is expected in the lowermost crust is somewhat unreliable, but is 6.95-7.0 km/s. The

possibility of compositional change within the crust has been modeled to explore the potential effects.

COMPOSITIONAL VARIATION IN ICELANDIC CRUST

In this model (Model 3) changes in the physical properties of the crust with depth are due to changes in chemical composition of the crust. This occurs by fractionation within magma chambers in the crust. It was shown by Kelley and Barton (2008) that magma chambers exist in the crust beneath most volcanic centers in the rift zones at intermediate depths (~10 km on average). Crustal accretion occurs through differentiation of magmas at various depths. As differentiation occurs, cumulates are formed. In general (assuming at least some melt-solid separation) cumulates are more mafic than erupted liquids (see MacLennan et al. 2001).

No primary (mantle derived) magmas have been identified amongst the eruptive materials in Iceland. Even those magmas that have erupted from basal crustal chambers are evolved. An explanation for this is that primary magmas are trapped in chambers at

the base of the crust where they differentiate prior to ascending. Therefore, the lower crust has a more mafic bulk composition than any melts that are erupted and likely

contains olivine rich cumulates. This characterization for the lower crust is also

suggested by Farnetani et al. (1996) and is the model adopted here.

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I consider in Model 3 that the crust above 10 km is basaltic in composition. From

10 to 15 km the crust is made up of more and more mafic compositions. The lowermost

5 km of the crust is composed of the most primitive composition in the crust. In fact, this

composition is more mafic than any that have been erupted. This composition was

calculated by Korenaga and Kelleman (2000) and is listed in table 5. I have also

calculated the composition of the most primitive melt to enter the base of the crust before

any effects of fractionation. This composition is also listed in Table 5. To do this, I took

the most mafic composition in the glass database and incrementally added the

components of olivine to it in order to remove the effect of the olivine crystallization that

took place in the magma chamber prior to eruption. The most primitive composition that was started with was that of a glass from the Reykjanes Peninsula. Olivine was added to

this composition until it was at a composition that is in chemical equilibrium with the

olivine composition of Fo91.8. This is the composition of the most magnesian olivine

crystal to be reported from Icelandic rocks. It comes from Theistareykir (MacLennan,

2003). The calculated parent composition is listed in Table 5. Also listed is a parent

composition that was calculated by Korenaga and Kelleman (2000). This composition

agrees with mine and was used in the modeling.

Predicted variations in density and seismic velocity

The calculator of Hacker et al. (2003) was used to find the density and P-wave

velocities through the crust while considering compositional change. In the Hacker

118 calculator, it is necessary to input mineral assemblages at varying depths. The mineral assemblages that were input for the varying depths and corresponding compositions were found using the IgPet software to calculate normative mineralogies. The results of the density and seismic velocity calculations are shown in Figures 33 and 34 respectively.

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Figure 33: Variation in density with depth according to Model 3.

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Figure 34: Variation in P-wave velocity with depth according to Model 3.

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This P-wave velocity is plotted on Figure 32 with those of previous studies and

that of Model 2. The values are steady but slightly increasing with depth below the

highly altered rocks of the upper few kilometers. The value in the lowermost crust is found to be 6.78 km/s. These depths would correspond to Palmason’s (1971) Layer 4.

The density that is found in Model 3 at the base of the crust is 3120 kg/m3 – quite similar to the value produced by Model 2. This is a robust estimate. The density contrast at the crust-mantle boundary according to Model 3 is 180 kg/m3 assuming a mantle density of 3300 kg/m3.

DISCUSSION AND CONCLUSIONS

The Icelandic rift zones are of course a dynamic system. Local temperature and

compositional variations exist. In particular, near active volcanic centers, geothermal

gradients will change temporarily due to magma within the crust. However, the models

that have been presented here are applicable to the Icelandic rift zones as a whole. They

are based on compositional, thermal, and geophysical data from across the island. The

models are intended to be used as a starting point for any regional or local studies. All

studies of local variation in crustal properties should be considered with respect to the

modeling discussed here.

My preferred model is that of compositional variation with depth in the crust. In

reality, some combination of Model 2 and Model 3 must exist. However, when evaluated

on their own, Model 3 better reproduces the physical properties of the crust. The effects

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from metamorphism in the middle to deep crust will be minimal in the absence of water.

The half-spreading rate of the divergent plate boundary in Iceland is 10mm/yr which leads to 1 km of spreading in 1 million years. Subsidence in the rifts is 1 km/my. The age of the rifts is 7 my. Therefore, the maximum depth of hydrated lavas in the rift zones is 7 km (and this ignores lateral motion). So then, there cannot be water present in the middle to lower crust.

It is useful to compare the crust of Iceland to oceanic crust at mid-ocean ridges.

There are seismic discontinuities at the base of lavas and base of the sheeted dikes in the lower parts of the oceanic crust. There is a relatively uniform and small increase in seismic velocity with depth.

I have shown that it is possible to develop a model for the crust that is consistent with petrological, geochemical, and geophysical information, but the only when the crust changes toward a more mafic composition with depth.

It is likely that the structure of the middle to lower crust is a mixture of sills, massive gabbro (i.e., solidified magma chambers), and layered gabbro (cumulates). The proportion of olivine rich gabbro and dunite increases with depth.

A potential problem of the models discussed here is that the high temperatures in the lower crust should lead to melting. The presence of large amounts of melt in the lower crust is inconsistent with seismic data. Therefore, maybe the lower crust is dominated by relatively refractory cumulates. Alternatively, the lower to middle crust could be mushy. That is, a mixture of melts and crystals. The melts may not exist in

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extensive bodies and therefore would not be detectable by seismology. Detailed structure

of the lower crust in Iceland has been difficult to determine using seismic data (Allen et

al., 2002) possibly because of the mushy crust.

It has been shown that magma chambers exist at the base of the crust throughout

the Icelandic rift zones. Furthermore, the geothermal gradients that have been shown

here confirm that the lower crust in Iceland must have high temperatures, and most likely

in many places partial melting occurs. This is in contrast to the claims that the lower

crust in Iceland must be cool (Menke and Levin, 1994; Menke 1995). These studies were

based on seismic profiles that poorly sampled the active rift zones and so are only appropriate for interpretation of the off-rift portions of the Icelandic crust. Much of the

ambiguity in the results of these more recent seismic studies and earlier ones that reported

higher lower crustal temperatures and partial melting in the crust (Palmason, 1971; Beblo

and Bjornsson, 1980; Gebrande et al., 1980; Eysteinsson and Hermance, 1985) can be

resolved by the results of this study, which predict physical properties based on a well

constrained composition and geothermal gradient.

The density and seismic velocities were calculated for Model 2 and Model 3 with

respect to depth through the changing crustal mineralogy or composition. The results are

shown in Figures 28 to 31. The results for Model 2 and Model 3 are quite similar to each

other. The density at the base of the crust is found to be 3120 kg/m3 and 3134 kg/m3 for

Model 2 and Model 3 respectively. The P-wave velocity is roughly 6.8 km/s and 6.78

km/s in the lower crust in Model 2 and Model 3 respectively.

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It is interesting to point out that at an approximate depth of 10 km in the crust in

the rift zones, there are a number of key changes. First, 10 km is the lowermost depth for

the upper to mid-crustal magma chambers to exist beneath volcanic systems that have

erupted lavas from a shallow magma chamber (Figure 13). Furthermore, even in systems

where the mid-crustal magma chamber is deeper than this limit, the chamber does not

exist much deeper than 10 km (perhaps 13 km; see Figure 14). This suggests a density

change in the crustal rocks that results in a loss of buoyancy of magma at this depth. This

could also be due to the crust becoming more brittle and therefore more capable of

allowing intrusion of large volumes of magma. Model 2 is based on a change in

composition at 10 km because of the tendency of magmas to pond around this depth.

Therefore, the density modeling is consistent with a change at this depth, which would

cause magma ponding.

The temperature at this depth according to the geotherm that has been presented

here is ~700 C. This temperature marks the approximate depth of the brittle ductile

transition. With some exceptions, earthquake foci have not been observed to occur

beneath 10 km (Bjornsson, 2008) further supporting the idea of a brittle-ductile transition

there. Furthermore, Bjornsson (2008) reports that there must be partial melting in the

crust beneath 10 km on the basis of electrical conductivity studies.

Model 2 produced yet another change at ~10 km in the crust of the rift zones. The thermodynamically predicted mineral assemblages that occur along the geotherm in this region show that the change from amphibolites facies rocks to granulite facies rocks

125 occurs at or around this depth. The density profile for Model 2 also show a contrast at

~10 km which might contribute to magma ponding. Also, the Perple_X software calculated that melt should be present in the crust below 12 km considering the composition and temperature that are present.

The broad scope of this study allowed for the observation that there is a tendency for magma to pond in the crust throughout the rift zones. The results are consistent for most of the volcanic systems in Iceland. Therefore, the rift zones have a common crustal structure with a density change in the middle crust caused by a compositional and/or mineralogic change. Furthermore, the temperature gradient combined with the composition and/or mineralogy change at this depth leads to a brittle-ductile transition in the rocks. The lower crust throughout the Icelandic rift zones is warm with probable partial melting throughout and magma ponding at the base of the crust beneath all volcanic systems, if not beneath the entire length of the spreading center.

126

Chapter 5: Conclusions

Pressures of crystallization of melts in chemical equilibrium with an ol-plag- cpx mineral assemblage were calculated for samples from all volcanic systems in Iceland.

The method based on the work of Yang et al. (1996) produces results that are accurate to

±110 MPa (1σ) and are precise to 80 MPa (1σ). These pressures correspond to depths of magma ponding in magma chambers within the crust. Evidence has been found for magma ponding at depths of ~0-35 km in the Icelandic crust. Magma chamber depths calculated beneath Askja, Bláfjall Table Mountain, Grimsvötn, Hengill, and Hekla all agree with estimates of other workers.

Most localities have erupted melts from deep chambers that appear to be at the base of the crust. Magma rising from the mantle has a tendency to pond at the crust- mantle boundary. There are also shallow chambers beneath most volcanic systems where magma ponds at a point of neutral buoyancy. The middle to lower crust is warm and probably partially molten. Crustal accretion occurs over a range of depths. The presence of multiple, stacked chambers and hot, porous crust suggests that magma evolution is complex and involves polybaric crystallization, magma mixing, and assimilation.

Magmas have erupted from a magma chamber located at 15.6±2.4 km depth beneath the Kverkfjöll volcanic system during its history. The greatest depth of eruption source in this system gives an estimate for the crustal thickness of ~31.5 km, which is in agreement with previous estimates. Recent and ongoing microseismic activity at Mt.

Upptyppingar in the Kverkfjoll system can be interpreted as the influx of magma into a

127

dike at 15-18 km depth. This depth is similar to previous eruption depths in this system

suggesting that the possibility exists for eruption at Mt. Upptyppingar as a result of the

current magma influx.

The geothermal gradients in the crust in the Iceland rift zones are high. The

mid to lower crust must be warm and is possibly partially molten. There must be either a

compositional change or a change in mineral assemblage in the crust with depth. Models

for both of these possibilities can be constructed based on geochemical data and

petrologic theory that are in agreement with observed density and seismic velocity

profiles through the crust. Temperatures at the base of the crust are 1185-1.194 °C. The density at the base of the crust is 3120-3134 kg/m3. The P-wave velocity in the lower

crust is predicted to be 6.78-7.0 km/s. The range in these values comes from the two

different models considered (Model 2 and Model 3). In reality, some combination of

these models must exist.

Iceland crustal studies are significant because while crustal accretion is

extremely important along mid-ocean ridges, they are very hard to study. The wealth of

geochemical, petrologic, and geophysical data that is available from Iceland provides the

unique opportunity to develop sophisticated, integrated models for crustal accretion

processes there. This work is a crucial first step toward these efforts. The integration of

all types of data is the only way to develop a truly complete model.

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148

Appendix A

Calculation of Pressure

Yang et al. (1996, Table 3) present three equations that allow direct calculation of pressure. These equations are used to calculate a series of liquid compositions (LP) lying along the ol-plag-cpx cotectic for the sample of interest at increments of 100 MPa (see

Table 8). The liquid compositions are converted to normative mineral components using the procedure described by Grove et al. (1993) assuming that ΣFe=FeO, and projected from plag onto the plane ol-cpx-qtz and from ol onto the plane plag-cpx-qtz. The pressure dependence of each normative mineral component in the predicted liquids (LP) is found by regression, and the pressure of crystallization is found from the regression equations using the projected normative mineral components for the original sample (LS). Thus for the projection from plag, values of P are calculated from predicted and observed ol, cpx and qtz, whereas for the projection from ol, values of P are calculated from predicted and observed plag, cpx and qtz. We have used these two projections because most basalt melts are saturated with plag and ol, and obtain six values of P for each sample. The average value is taken as the pressure of crystallization, and all values are used to calculate the uncertainty (1σ) associated with the calculated pressure. An Excel spreadsheet to perform these calculations is available from the authors upon request.

The approach described above uses all three of the equations given by Yang et al.

(1996) to calculate pressure. In contrast, Michael and Cornell (1998) used one of the

149 equations of Yang et al. (1996, Eq. 2) to calculate the pressures of crystallization of

MORB. Results obtained using the two methods are compared in Appendix B.

Table 8. Example of Method Used to Calculate Pressure

Sample 1.7.41 P Kb Plag Projection Ol Projection Results P Kb SiO2 47.6 Ol Cpx Qtz Plag Cpx Qtz Ave 6 6.70 Al2O3 15.9 Sample2 42.78 51.89 5.33 68.66 28.43 2.92 1σ6 0.42 TiO2 1.23 Ol Cpx Qtz Plag Cpx Qtz Range 6 1.17 Fe2O3 0 0.001 Predicted3 28.93 71.45 -0.38 61.07 39.14 -0.21 ΔPlag-Ol7 0.30 FeO 10.08 1 30.80 68.83 0.37 62.26 37.54 0.20 MnO 0.17 2 32.71 66.15 1.14 63.47 35.91 0.62 MgO 9.18 3 34.67 63.41 1.92 64.69 34.27 1.04 CaO 12.9 4 36.67 60.61 2.72 65.93 32.61 1.46 Na2O 1.8 5 38.71 57.75 3.54 67.18 30.92 1.90 K2O 0.1 6 40.80 54.82 4.38 68.46 29.21 2.33 P2O5 0.11 7 42.94 51.82 5.24 69.74 27.48 2.78 TOTAL 99.070 8 45.13 48.75 6.12 71.05 25.72 3.23 9 47.38 45.61 7.01 72.37 23.94 3.68 10 49.67 42.39 7.93 73.71 22.14 4.14 11 52.03 39.10 8.88 75.07 20.31 4.61 12 54.44 35.72 9.84 76.45 18.46 5.09 Slope4 0.47 -0.34 1.17 0.78 -0.58 2.27 Intercept4 -13.36 24.25 0.70 -47.49 22.86 0.62 R2 0.9985 0.9985 0.9985 0.9995 0.9995 0.9995 P Calc5 6.77 6.82 6.95 6.06 6.36 7.23 1Glass analysis from Schiellerup (1995) 2Projected mineral components for sample 1.7.4 calculated using procedure of Tormey et al. (1987) 3Projected mineral components calculated at the pressures listed in Column 3 using procedure of Tormey et al. (1987) 4Slope and intercept from regression of projected mineral components (Columns 5-10) versus P (Column 3) 5Pressures calculated from regressions and projected mineral components for sample 1.7.4. 6Average, 1σ, and range calculated from six pressures obtained by regression 7Difference between average P calculated for Plag projection and average P calculated for Ol projection.

150

Appendix B

Accuracy and Precision

In principle, accuracy can be determined by comparing pressures calculated for glasses

produced in experiments with the reported run pressure. A problem with this approach is that experimental glass compositions show considerable variability, much greater than that expected from analytical uncertainties, which reflects problems inherent in experimental studies (eg. Ford et al. 1983; Yang et al.1996; Faloon et al. 2004). These include uncertainties in measurement of pressure and temperature, unknown amounts of volatiles in nominally anhydrous experiments, loss of Fe2+ and alkalis from the charge

during the experiment, modification of melt compositions by quench crystallization, and

run durations too short for the attainment of equilibrium. It appears that some

experimental glasses reported to be in equilibrium with ol, plag, and cpx do not have

compositions that represent liquids lying on the ol-plag-cpx cotectic. Accordingly, large

differences are expected between calculated and experimental pressures leading to poor

estimates of accuracy.

The experimental glass compositions used to construct Fig. 2 are taken from the

sources listed in the caption. Glass compositions from experiments on natural samples

were combined with glass analyses from experiments by Shi (1993) in the system CaO-

MgO-Al2O3-SiO2-FeO-Na2O and used to map the likely positions of the cotectic at

different pressures (samples with anomalous compositions, that plot off the main data array at a given pressure, were discarded). The actual positions of the cotectic were

151

located using a subset of these glasses that have compositions that form tightly defined

arrays in projections such as those shown in Fig. 2, and therefore appear to anchor the

positions of the cotectic at different pressures. The glass compositions in both data sets

(excluding alkaline compositions and glasses from Shi’s experiments) were used to test

the accuracy of pressures calculated for sub-alkaline and transitional compositions. For

the extended data set (91 glasses), calculated pressures agree with experimental pressures

to ±120 MPa (1σ), which is about the same as uncertainties for pressures calculated by

similar methods (Herzberg, 2004). For the smaller data set (59 glasses), calculated

pressures agree with experimental pressures to ±90 MPa (1σ). These results suggest that

reported pressures are accurate to ~110 MPa.

Pressures were also calculated using the method described by Michael and

Cornell (1988). For the extended data set (91 glasses), pressures calculated using Eq. 2 of

Yang et al. 1996) agree with experimental pressures to ±160 MPa (1σ), whereas for the

smaller data set (59 glasses), calculated pressures agree with experimental pressures to

±120 MPa (1σ). These results indicate that pressures calculated using the method used here are more accurate than those estimated using the method described by Michael and

Cornell (op cit).

Yang et al. (1996) estimated the uncertainty in pressure estimates (precision) from the standard deviation of the mean of replicate electron microprobe analyses, and obtained a value of about ±50 MPa (2σ). The uncertainties in calculated pressures for the

152 experimental data sets determined as described in Appendix 1 are somewhat larger than this value, ±60-80 MPa (1σ).

153

Appendix C

Supplemental Data ual ngill, , 2006), I I , 2006), d these with with d these 9 1 NVZ Blafjall Ridge (Bf) 3 NVZ Burfell(Bu) 2 2 Moore and Calk, 1991 Hansenand 2000 Gronvold, Schiellerup, 1995 NVZ Herdubreid (Hb) les. 7 p 0.59 0.79 0.440.05 2.13 0.100.69 2.96 0.02 0.71 1.31 0.29 0.61 2.06 0.46 0.52 3.31 0.13 0.59 1.41 0.47 0.45 2.10 0.75 0.48 2.26 0.31 0.57 0.21 1.89 0.31 0.28 0.54 0.16 0.57 0.51 ill sam g Table 9. Supplemental Data. Sample Localities and Summary of Glass Compositions Glass of Summary and Localities Sample Data. Supplemental 9. Table other Hen other Maelifell and Midfell). It seems likely that some samples from Thingvellir also belong to thissystem, though Ihave not groupe Zone NVZ I retainthe locality names given thein original papers.However, following the suggestionof P. Meyer (personal communication Locality (Th) Theistareykir have used reported values of latitude and longitudevolcanicto assign centers samples (Tablefrom Bardabunga/Grimsvötn 3). Also, I have andgrouped from Katla/Heklatogether to samples individ belonging to the Hengill volcanic system (eg. samples from He reference Slater et al. 2001 sample_idn1 Average Max Min Average Max Min Average Max Min Average Max Min SiO2Al2O3TiO2 FeOT 49.04MnO 16.29MgO 49.98CaO 16.66 8.02Na2O 48.43 0.15K2O 15.61 9.82 9.16P2O5 48.78 13.76 0.21 14.34TOTAL 1.59 10.56Mg# 7.32 49.40 15.03 16.00 0.11FeO/MgO 0.15 1.95 99.40 8.17CaO/Al2O3 48.30 11.97 12.78 13.00 0.82Na+K 0.20 0.84 100.09 0.20 1.32 50.18 13.40 7.30 12.03 14.48 1.12 98.23 0.22 0.91 50.69 0.05 10.60 12.80 2.35 1.64 15.62 8.42 99.64 0.72 0.17 0.78 49.82 12.12 11.10 0.25 12.30 2.79 100.66 2.05 6.17 1.67 0.11 0.84 48.10 15.59 10.88 14.23 99.12 0.36 1.34 1.38 6.26 2.17 49.57 0.15 10.36 0.89 12.15 15.00 99.18 0.14 2.41 2.64 7.55 46.98 1.26 11.39 0.04 9.00 0.79 13.64 99.86 0.22 2.68 11.92 3.25 2.24 3.98 0.17 97.85 12.67 0.75 10.48 0.38 2.22 3.92 1.77 98.93 13.72 7.53 0.18 0.78 0.12 99.95 11.92 1.37 2.24 2.88 0.13 8.28 0.73 98.19 0.28 1.52 2.48 3.43 0.89 6.84 1.74 0.32 1.96 1.00 2.57 1.33 0.21 0.80 2.45 2.71 2.14 Continued

154

Table 9 continued. 4 NVZ (Hr) Hrimalda 8 Askja (As) Askja NVZ Sigurdsson&Sparks, 1981 Hansen and Gronvold, 2000 4 NVZ Schiellerup,1995 Seljahjalli (Se) Seljahjalli (Se) 0.14 0.16 0.11 0.16 0.18 0.14 0.48 0.64 0.34 0.10 0.13 0.05 1.71 1.77 1.68 2.430.57 2.45 0.58 2.41 0.56 2.06 0.53 2.81 0.53 1.59 0.52 1.08 0.44 1.23 0.56 0.93 0.36 0.63 0.69 0.59 Zone NVZ sample_id Averagen5 Max Min Average Max Min Average Max Min Average Max Min LocationReference (Ha) Halar Schiellerup, 1995 SiO2Al2O3TiO2 FeOT 48.76MnO 15.29MgO 49.23CaO 15.71 10.90Na2O 48.42 0.15 14.91K2O 11.20 8.14P2O5 47.15 12.25 14.62 0.18TOTAL 10.60 2.13 8.40Mg# 47.77 12.72 15.12 12.22 0.12FeO/MgO 0.15 2.23 99.61 7.86 46.68CaO/Al2O3 12.03 14.30 12.36 1.34Na+K 0.18 0.80 100.30 0.17 51.21 2.00 7.58 11.52 13.63 12.00 1.42 98.69 0.19 0.81 52.68 0.14 11.78 7.77 2.45 14.72 2.26 13.40 98.60 1.28 0.17 0.79 49.88 10.92 12.80 0.29 7.44 15.54 2.96 2.38 99.01 1.61 49.24 0.20 0.79 10.20 14.52 11.02 5.93 0.33 2.26 98.21 2.14 49.80 1.66 0.30 12.40 0.82 15.92 9.22 7.73 99.63 0.27 2.39 48.58 2.61 1.59 13.57 8.90 0.14 0.74 100.29 9.98 4.84 0.15 3.20 3.13 2.34 14.32 0.07 99.06 0.75 8.44 8.88 0.27 0.77 14.92 3.17 2.43 99.56 0.12 0.84 10.50 13.46 0.00 2.03 99.95 1.43 0.01 2.87 0.64 8.00 99.06 0.11 1.06 2.19 3.64 0.99 1.24 0.15 1.81 1.10 1.34 0.82 0.10 0.89 2.13 2.32 1.86 Continued

155

Table 9 continued 9 kverkfjoll (Kv) kverkfjoll (Kv) NVZ Hansenand Gronvold, 2000; Hoskuldsson et al. 2006 0 1 Kistufell (Ki) (Ki) Kistufell NVZ Breddam, 2002 1 1 Sprengisandur (Sp) NVZ 2 0.08 0.11 0.05 0.14 0.32 0.05 0.07 0.09 0.05 0.63 0.72 0.43 0.65 0.70 0.57 0.55 0.62 0.39 0.65 0.68 0.64 0.38 0.47 0.35 LocationReference Gigoldur (Gi) 2000 Hansen Gronvold, and sample_id Meyer et al.,n1 Average 1985 Max Min Average Max Min Average Max Min Average Max Min Zone NVZ SiO2Al2O3TiO2FeOT 49.45MnO 15.77MgO 50.35 0.80CaO 16.58 8.63Na2O 48.90 0.14 14.28K2O 1.19 10.25 8.84P2O5 49.92 13.93 14.32 0.16TOTAL 0.37 2.10 7.58 9.98Mg# 51.00 14.71 16.00 0.11FeO/MgO 0.07 1.36 2.35 99.80 11.40 48.90CaO/Al2O3 7.73 13.55 12.45 0.99Na+K 0.20 0.89 101.15 15.87 0.14 48.34 2.39 1.84 12.34 15.87 7.70 1.33 98.81 0.26 0.96 9.38 48.98 0.01 0.77 2.24 13.90 16.28 2.18 8.92 99.75 0.77 0.15 0.82 47.70 9.18 10.06 15.10 0.13 0.99 2.76 100.81 2.42 5.65 1.54 49.82 0.16 0.86 12.90 13.80 9.49 99.14 0.27 1.11 1.90 1.93 51.87 9.71 2.81 0.18 0.96 13.58 14.09 99.95 8.69 0.00 0.87 1.78 48.09 2.38 10.56 1.12 12.02 13.49 0.15 100.70 0.68 14.75 0.06 3.25 1.87 3.00 9.27 99.22 0.95 9.63 0.24 0.87 15.79 0.18 3.72 1.71 99.23 1.99 5.12 1.02 10.58 0.34 13.53 0.93 100.32 0.00 2.11 2.35 1.85 0.85 6.70 8.93 0.04 0.84 97.98 0.53 2.84 2.95 1.96 4.54 0.75 0.83 3.31 1.26 1.77 0.83 0.21 2.04 2.98 0.69 3.53 1.69 Continued

156

Table 9 continued Z V Hansen and Gronvold, 200 Hansen and Gronvold, 9 9 ZE V Bardabunga-Grimsvotn Veidivotn (Vd) Compiled 12 ZE V Grimsvotn (Gr) Grimsvotn Meyer et al., 1985 81 1 ZE V E e n o SiO2Al2O3TiO2FeOT 49.39MnO0.200.250.170.210.270.160.200.270.160.110.180.01 13.54MgO6.808.595.036.668.674.516.758.674.517.899.266.10 50.60 2.21CaO 14.67 12.83Na2O 47.31 12.27K2O0.270.630.080.290.560.080.280.630.080.150.320.05 4.16 15.12P2O5 49.78 11.50 13.52TOTAL 9.89 0.93 2.52Mg#0.490.610.370.470.610.330.480.610.330.590.670.46 50.86 13.60 14.60FeO/MgO1.943.011.152.143.561.162.013.561.151.312.070.87 13.00 0.19 2.33 3.38 99.46CaO/Al2O30.850.940.740.820.950.710.840.950.710.850.900.80 49.06 9.74 12.08Na+K 16.09 100.88 0.44 49.54 3.76 1.84 13.54 11.17 98.46 10.09 50.86 0.00 0.98 2.44 14.67 2.79 13.80 99.64 12.88 47.31 12.08 0.22 2.24 3.11 100.23 8.74 4.01 16.09 49.58 14.95 99.07 0.40 11.39 4.16 1.87 9.89 1.92 50.18 15.88 99.50 13.80 0.00 0.93 2.48 9.79 2.73 48.03 100.88 12.93 8.74 0.21 12.66 1.32 3.38 98.46 3.67 12.79 0.44 7.95 98.96 2.24 1.84 1.95 14.08 99.90 0.00 0.82 2.26 2.76 10.96 98.00 0.12 2.70 4.01 0.22 1.82 1.92 0.06 2.41 3.02 1.90 sample_idn Average Max Min Average Max Min Average Max Min Average Max Min LocationReference Bardabunga (Ba) Meyer et al., 1985 Z Continued

157

Table 9 continued 2 7 Laki (Lk) EVZ 4 2 Veidivotn (Vd) EVZ 4 Veidivotn (Vd) EVZ Thordason etThordason al., 2003 Compiled et Thordason al., 1996 1 0.42 0.610.45 0.26 0.47 0.21 0.43 0.26 0.51 0.15 0.61 0.28 0.46 0.61 0.51 0.05 0.67 0.51 0.43 0.90 0.39 0.32 0.50 0.29 Reference Mork, 1984 sample_idn1 Average Max Min Average Max Min Average Max Min Average Max Min Location Veidivotn (Vd) Zone EVZ SiO2Al2O3TiO2FeOT 50.30MnO 13.76MgO 51.75 2.01CaO 14.38 12.86Na2O 49.05 0.22 13.27K2O 2.68 13.51 5.80P2O5 49.63 10.67 0.26 13.64TOTAL 11.95 1.79 2.68 6.43Mg# 50.46 11.71 13.94 12.73 0.19FeO/MgO 0.19 1.53 3.12 98.92 5.20CaO/Al2O3 48.00 9.97 13.30 14.11 2.23Na+K 0.22 0.78 99.71 0.22 49.92 1.97 2.12 7.36 14.19 11.32 10.56 2.38 0.26 0.87 97.79 51.75 0.16 0.78 2.21 9.38 15.88 3.10 11.73 11.69 2.00 98.99 0.19 0.70 48.00 12.93 0.15 6.42 10.86 14.11 1.67 2.27 3.73 99.66 1.79 49.75 0.18 0.83 12.82 11.58 7.95 0.20 6.84 2.68 2.11 2.48 97.40 50.51 2.07 0.26 0.84 13.70 14.08 14.72 9.38 0.05 98.95 2.44 0.78 48.81 2.42 1.13 0.01 11.45 0.82 9.97 17.69 5.20 99.90 0.16 3.12 3.23 2.52 1.81 0.23 0.81 10.01 12.17 97.40 5.26 0.22 4.37 1.82 2.31 2.38 0.29 11.70 0.90 99.56 6.96 0.05 2.17 2.69 0.87 2.73 100.19 8.81 0.17 0.70 4.10 0.34 98.07 2.84 3.01 3.73 0.78 0.73 4.31 2.21 1.90 0.86 1.75 0.19 0.73 3.20 3.46 2.83 Continued

158

Table 9 continued WVZ Langjokull (La) Langjokull Meyer et al., 1985 4 3 SFZ Hekla-Katla Compiled 54 SFZ (Hk) Hekla Meyer et al., 1985; Moune et al, 2007 92 1 0.58 0.98 0.21 0.56 1.29 0.34 0.57 1.29 0.21 0.02 0.04 0.01 0.40 0.52 0.30 0.44 0.57 0.31 0.42 0.57 0.30 0.63 0.65 0.61 Meyer et al., 1985; Lacasse et al 2007 Reference Zone SFZ Location (Kt) Katla sample_idn Average Max Min Average Max Min Average Max Min Average Max Min SiO2Al2O3TiO2FeOT 48.74MnO 13.07MgO 50.54 3.47CaO 13.94 14.65Na2O 46.60 0.23 11.60K2O 5.32 16.99 5.43P2O5 47.57 10.43 0.29 13.75TOTAL 11.81 1.75 2.74 7.23Mg# 50.00 12.36 15.90 0.19 13.55FeO/MgO 0.24 3.44 3.22 99.56 4.03CaO/Al2O3 46.40 8.77 11.60 17.09 2.80Na+K 0.22 0.80 101.43 0.70 48.08 5.27 1.50 5.93 13.45 10.74 9.23 4.12 98.56 0.36 0.89 50.54 0.00 7.73 2.36 2.88 15.90 3.32 13.60 99.00 14.02 1.63 0.15 0.69 46.40 11.60 0.35 3.83 3.45 3.88 100.35 8.25 4.02 17.09 2.41 48.80 0.22 0.78 15.38 97.11 5.72 0.95 10.60 5.32 2.31 9.23 1.88 49.43 4.04 0.36 0.89 16.18 99.24 13.60 7.73 0.00 1.75 2.82 9.64 47.69 3.44 1.35 101.43 0.15 14.80 0.66 8.25 3.83 0.30 0.92 3.88 9.88 97.11 4.73 2.58 0.13 0.79 13.70 9.15 0.95 99.83 1.02 1.50 9.46 2.65 4.12 0.14 0.88 14.18 100.48 9.73 0.00 0.82 2.07 1.35 3.39 0.11 0.66 12.95 99.00 8.68 0.00 1.06 2.18 4.73 0.89 0.01 1.14 2.00 1.88 0.94 0.97 0.00 0.80 2.09 2.22 2.02 Continued

159

Table 9 continued 0 WVZ Kalfstindar (Ka) Kalfstindar 01 WVZ Efstadalsfjall (Ef) Efstadalsfjall 23 WVZ Raudafell (Ra) 91 1 0.19 0.27 0.12 0.28 0.31 0.26 0.18 0.32 0.14 0.13 0.13 0.12 0.50 0.55 0.44 0.55 0.56 0.52 0.53 0.58 0.47 0.59 0.60 0.58 Location Hlodufell (Hl) P2O5TOTALMg# 0.16 99.97 101.34 0.21 99.07 0.13 99.49 0.22 100.13 98.95 0.25 99.47 0.18 100.95 0.24 98.90 0.29 99.41 0.20 99.66 99.24 0.13 0.14 0.12 Reference Moore Calk, and 1991sample_idn AverageSiO2 Max Moore and Calk, 1991Al2O3TiO2 MinFeOT 49.22MnO 14.30 Moore and Calk, 1991 AverageMgO 50.00 1.87 MaxCaO 15.30 12.46Na2O 48.50 0.22 13.10K2O 2.68 14.40 Min 7.03 Moore and Calk, 1991 48.46 12.17 0.25 15.00 11.20 1.55 2.36 7.85 Average 48.80 12.70 15.60 Max 0.20 11.44FeO/MgO 1.76 2.68 6.26CaO/Al2O3 48.00 11.40 13.90 12.10 1.79Na+K 0.20 Min 0.85 48.08 1.94 2.15 7.74 12.18 14.40 11.20 2.23 0.22 0.91 Average 48.80 12.40 1.66 7.99 2.22 15.60 2.54 12.03 Max 1.48 0.19 0.81 47.40 11.60 13.50 7.44 13.70 1.97 2.27 2.95 Min 1.48 0.20 0.81 12.52 47.88 11.20 7.65 2.80 2.17 15.95 2.30 1.63 0.22 13.00 0.88 48.10 8.62 16.10 10.68 1.68 2.21 2.49 1.40 11.60 0.18 0.76 47.70 15.70 6.68 10.80 2.45 2.56 1.58 0.87 12.30 1.38 10.60 0.17 2.08 2.44 2.05 8.60 12.40 0.95 1.42 0.17 1.30 12.20 8.79 2.39 0.78 1.34 2.19 0.16 8.39 2.77 1.24 2.24 0.77 2.25 1.29 2.17 0.79 1.21 0.76 2.32 2.37 2.29 Zone WVZ Continued

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Table 9 continued 3 0 Z 1 V Hengill (Hg) Hengill Tronnes, 1990 4 ZW V Maelifell (Ma) 9 8 ZW V Midfell (Mi) Midfell Gurenko Sobolev, 2006 and Hansteen, 1991 0 ZW 0.13 0.240.58 0.04 0.63 0.04 0.53 0.18 0.63 0.02 0.67 0.03 0.50 0.03 0.66 0.02 0.68 0.32 0.64 0.87 0.50 0.01 0.65 0.30 V W e n o Location (Tg) Thingvellir SiO2Al2O3TiO2FeOT 48.79MnO 15.09MgO 49.39 1.49CaO 15.84 10.87Na2O 48.04 0.14 14.00K2O 2.07 12.17 8.33P2O5 49.03 12.82 0.16 14.80TOTAL 9.38 1.02 2.35 9.08Mg# 49.63 13.89 15.72 0.10FeO/MgO 0.09 100.09 9.55 0.85 2.60 7.63CaO/Al2O3 47.97 11.86 14.15 1.32 100.81Na+K 0.17 0.85 11.78 0.20 48.80 2.14 2.07 9.32 15.04 14.60 99.20 1.60 0.21 0.89 8.53 49.06 0.00 15.54 9.99 0.71 1.62 100.46 14.90 2.48 1.03 0.15 0.78 48.63 101.45 8.68 13.33 14.28 0.04 6.62 0.92 2.15 2.80 1.03 99.73 48.77 0.15 1.02 14.36 14.04 8.95 9.37 0.28 1.00 1.47 2.11 98.82 53.40 1.78 0.15 14.80 1.07 15.70 8.41 9.83 0.01 0.85 1.87 99.22 1.66 47.10 0.88 13.77 0.14 12.90 0.87 12.12 8.75 0.05 98.65 2.11 1.96 12.16 2.30 0.93 0.20 0.98 15.30 6.78 99.27 0.06 15.50 3.43 1.81 1.49 1.02 0.29 1.03 8.75 101.03 9.74 7.83 0.05 0.76 2.49 0.86 1.90 0.14 97.55 0.94 3.35 0.27 1.88 3.29 1.99 0.87 1.03 4.24 1.54 1.84 1.05 0.94 0.03 0.60 2.81 4.07 1.55 Reference Meyer et al., 1985 sample_idn1 Average Max Min Average Max Min Average Max Min Average Max Min Z Continued

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Table 9 continued 12 RP Moore and Calk, 1991 1.51 3.43 0.710.18 2.56 0.87 3.09 0.01 2.13 0.43 0.54 0.36 0.188.01 0.292.10 9.99 0.100.16 3.29 3.35 0.23 6.440.56 1.03 1.47 0.25 6.98 2.65 0.68 0.00 0.21 2.28 5.82 0.33 2.84 0.30 4.07 0.47 0.41 2.50 1.49 0.50 0.26 3.09 0.43 3.28 2.86 48.8914.43 53.40 15.8410.88 47.10 12.90 15.30 48.7813.48 13.86 49.00 8.41 15.54 15.00 12.97 48.50 13.00 7.83 13.90 11.49 12.10 11.80 11.00 206 Location (Hg) complex Reference Hengill Compiled sample_idn Average (Ge) Geitafell SiO2 MaxAl2O3 TiO2 MinFeOT MnO MgO AverageCaO MaxNa2O K2O Min P2O5 TOTALMg# FeO/MgOCaO/Al2O3 99.81Na+K 1.47 101.45 0.93 97.55 4.24 1.07 99.74 0.86 0.60 100.55 2.03 0.83 98.89 2.37 0.85 1.76 0.79 Zone W VZ

162