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
By
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
Copyright by
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 midocean 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.
iv 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.
v 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
vi 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
x 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.
2
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.
4
Table 1. Summary of Estimates of Magma Chamber Depth
Center
Krafla
Depth Range
Theistareykir
Depth Range
Bláfjall
Depth Range
Askja
Depth Range
Grimsvötn
Depth Range
Kistufell
Torfajökull
Range
Hekla
Range
Katla
Range
Eyjafjallajökull
Range
Vestmannaeyjar
Method
Seismology
Geodesy/Gravity
Geodesy
Magnetotelluric
Geodesy
Geodesy
Seismology
Geodesy
Geodesy
Geodesy/Gravity
Geodesy
Geodesy/Gravity
Depth (km)
3-<7
3
3
3,5-7
3, 5-10, >20
3
3
3
3, >5
2.5
2.4,21
2.8, 21
>2.4, 21
30
10-31
10-31
7-10, 14
2-7, 14-26
>2, 26
1.5-3.5
1.5-3.5
1.5-3.5
3,16
3, 20
3,16
>1.5, 20
1.5-4
>1.6
>4
>1.5, 4
>35
10-15, 15-25
>3
8
>7
8, >15
<14
8
>3, 25
8
8
9
6.5
5.0-6.0
>6.5
≥14
11
>5, >14
3
4.7
2
>2, 4.7
3.5
6.3
>3.5, 6.3
15-25
10-15
18-28, 28-35
10-17
>10, 35
8-11
7
7
5.5
3.5
13
2
Petrology
Petrology
Petrology
Petrology
Geodesy
Geodesy
Geodesy
Geodesy/Gravity
Geodesy
Geodesy
Gravity/magnetics
Geodesy
Teleseismic P-wave delays
Petrology
Magnetotelluric
Magnetotelluric, Geotherm
Seismology
Petrology
Seismology
Seismology
Geodesy
Geodesy
Magnetotelluric
Geodesy
Geodesy
Geodesy
Geodesy
Seismic
Geodesy
Seismology
Geodesy
Seismology
Geodesy
Geodesy
Seismology
Seismology
Petrology
Petrology
Range
Hengill-Hrómundartindur Petrology
Geodesy
Geodesy
Petrology
Króksfjördur
Petrology
Thingmuli
Petrology
Snaefell
Petrology
Austerhorn
* Estimated from Fig. 4 in Schiellerup (1995).
5
Reference
Einarsson, 1978
Bjornsson et al., 1979
Tryggvason, 1980
Bjornsson et al.., 1985
Tryggvason, 1986
Ewart and Voight, 1991
Brandsdottir et al., 1997
Sigmundsson et al., 1997
Arnadottir, et al., 1998
Rymer et al., 1998
de Zeeuw-van Dalfsen et al., 2004
de Zeeuw-van Dalfsen et al., 2006
Slater et al., 1997
McClennan et al., 2001
Schiellerup 1995
Schiellerup 1995*
Tryggvason, 1989
Rymer & Tryggvasson, 1993
Sturkell & Sigmundsson, 2000
de Zeeuw-van Dalfsen et al., 2005
Pagli et al., 2006
Sturkell et al., 2006
Gudmundsson & Milsom, 1997
Sturkell et al., 2003
Alfaro et al., 2007
Breddam, 2002
Eysteinsson & Hermance, 1985
Gudmundsson, 1988
Soosalu & Einarsson, 1997
Gunnarsson et al. 1998
Soosalu & Einarsson, 2004
Soosalu et al, 2006a
Sturkell et al., 2006
Kjartansson & Gronvold, 1983
Eysteinsson & Hermance, 1985
Sigmundsson et al., 1992
Linde et al., 1993
Tryggvason, 1994
Jónsson et al. 2003
Soosalu & Einarsson, 2004
Sturkell et al., 2006
Gudmundsson et al..1994
Sturkell et al., 2003
Soosalu et al. 2006b
Sturkell et al., 2003
Pedersen and Sigmundsson, 2004, 2006
Einarsson & Bjornsson, 1979
Gebrande et al., 1980
Furman et al., 1991
Thy, 1991
Hansteen, 1991
Sigmundsson et al., 1997
Feigl et al., 2000
Jónasson et al, 1992
Frost and Lindsley, 1992
Hards et al., 2000
Furman et al., 1992
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 preeruptive 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.
6
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
7
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.
8
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).
9
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).
11
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).
12
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 plagliq 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.
13
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 (18-1 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.
14
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
15
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
16
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 reequilibration 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
18
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).
19
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 subalkaline (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.
20
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
22
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.
23
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
24
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.
25
Table 2. Summary of Calculated Pressures and Temperatures
Zone
NVZ
Locality
n
Ave P
Max
Min
Ave T
Max
Min
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
Bardabunga
17
270.7
841.7
64.9
1175.3 1209.8 1164.0
EVZ
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
Katla
17
426.7
656.3
199.3 1183.8 1212.4 1159.8
SFZ
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
Langjokull
3
253.4
651.5
34.7
1193.4 1216.8 1178.7
WVZ
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
Geitafell
12
399.9
536.0
281.4 1185.1 1197.7 1178.4
RP
1
201.7
1184.6
Reykjanes 1
462
349.6 1032.5
0.4
1186.0 1232.1 1134.4
All Data
1
Samples from Meyer et al. (1985)
Average, maximum, and minimum values based on results obtained for all samples from each locality
26
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).
27
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
29
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.
30
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
31
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.
32
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 olplag-cpx and from j to k by polybaric crystallization of ol-plag accompanied by
assimilation of cpx.
33
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
34
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).
35
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.
36
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
37
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.
38
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.
39
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).
40
41
RP
WVZ
SFZ
EVZ
Zone
NVZ
Locality
Theistareykir
Theistareykir
Herdubreid
Burfell
Blafjall Ridge
Halar
Seljahjalli
Askja
Hrimalda
Gigoldur
Sprengisandur
Kverkfjoll
Kverkfjoll
Kistufell
Northeastern
Bardabunga
Bardabunga
Grimsvotn
Grimsvotn
Veidivotn
Veidivotn
Laki
Katla
Hekla
Hekla
Langjokull
Hlodufell
Raudafell
Efstadalsfjall
Kalfstindar
Thingvellir
Thingvellir
Hengill Gp I
Hengill Gp II
Hegill Gp II
Hengill Gp III
Hengill Gp III
Geitafell
Reykjanes
P MPa
555
555
484
529
511
717
536
593
859
822
518
842
842
512
512
622
622
766
656
1032
956
652
517.0
651.0
627.0
627.0
592.0
556.0
613.0
503.0
507.0
±74
±34
±0.25
±0.86
±26
-
Range
±30
±76
±25
±9
±0.33
-
Max
Preferred
Max
Preferred
Preferred
Upper Limit
Max
Max
Max
Max
Max
Max
Max
Max
Max
Max
Max
Preferred
Max
Max
Max
Preferred
Max
Max
Comment
Max
Max
Table 3. Pressures and Depths of Cotectic Crystallization
±2 .6
±1.26
±0.88
±3.01
±0.91
-
-
-
-
18.0
-
-
18.2
29.6
21.9
±1.1
±2.3
±0.89
±0.32
±1.2
17.0
18.6
18.0
25.3
18.9
20.9
28.9
26.9
23.1
36.3
34.0
22.9
18.2
22.9
22.1
19.6
17.7
17.8
-
Range
-
Z Km
19.5
P MPa
308.0
310.0
325.0
331.0
466.3
425.0
501.7
335.0
313.0
322.0
221.0
257.0
311.0
340.0
340.0
346.0
±33
±32
-
±37
±36
Range
±19
±15
±45
±48
Min
Min
Min
Preferred
Lower Limit
Upper Limit
Preferred
Min
Preferred
Min
Comment
Z Km
10.9
10.9
11.4
11.7
16.4
17.7
11.8
11.0
11.3
9.1
10.9
12.0
12.2
±1.6
±1.7
±1.3
±1.3
±1.2
±1.14
-
±0.66
±0.51
-
Range
-
P MPa
131
149
173
175
185
176
170
54
90
90
143
65
177
97
154
35
141
145
199.3
74
168.0
198.0
202.0
Range
±27
±56
±90
±90
±62
±58
±64
Min
Min
Min
Min
Preferred
Min
Preferred
Min
Preferred
Min
Min
Min
Min
Min
Comment
Min
Preferred
Min
-
-
-
-
7.1
±2.2
±2.0
±2.3
-
Range
±1
±1.99
±2.3
±2.6
5.0
6.2
5.4
5.0
5.1
7.0
2.6
5.9
7.0
-
Z Km
5.3
6.1
6.2
6.5
6.2
6.0
1.9
3.2
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
42
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.
43
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.
44
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.
45
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 (silicarich) 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
46
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.
47
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 yaxis. 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.
48
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 yaxis. 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
49
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.
50
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.
51
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).
52
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
53
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
54
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
55
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.
56
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.
57
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.
58
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
60
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.
62
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
63
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.
64
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.
65
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
66
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).
68
69
2
LC
3
MC
MgO
4.80
7.08
7.14
7.57
7.52
7.50
7.49
7.53
7.36
7.53
4.55
4.83
4.81
4.77
4.81
4.71
4.54
6.33
6.7
CaO
9.85
11.22
11.32
12.12
11.56
11.57
11.59
11.59
11.58
11.86
8.94
9.34
9.24
8.93
9.41
9.95
9.99
10.3
10.58
Na2O
2.73
2.42
2.45
2.00
2.35
2.29
2.35
2.33
2.35
2.29
2.36
2.84
2.76
2.83
2.37
2.75
2.75
1.26
1.26
K2O
0.72
0.42
0.46
0.37
0.39
0.40
0.39
0.38
0.40
0.38
0.63
0.69
0.67
0.67
0.72
0.65
0.69
0.43
0.48
P2O5
0.40
0.23
0.21
0.17
0.19
0.20
0.19
0.21
0.22
0.19
0.46
0.72
0.83
0.83
0.52
0.45
0.47
0.21
0.29
47.91 14.92 2.54 12.79 0.23 6.73 10.97 2.44 0.47 0.36
50.03 13.43 2.83 13.88 0.22 5.72 10.07 2.07 0.57 0.35
MnO
0.24
0.20
0.19
0.21
0.16
0.19
0.19
0.19
0.18
0.20
0.25
0.31
0.31
0.34
0.26
0.27
0.04
0.19
0.19
TOTAL
99.54
96.14
99.67
100.43
99.53
99.44
99.57
99.47
99.34
99.76
98.18
98.87
100.32
99.99
98.01
100.09
97.98
99.81
99.82
0.38
0.38
0.38
0.47
0.24
2.03
2.54
2.31
3.31
1.55
0.73
0.75
0.75
0.83
0.69
21.97
31.53
3.15
1174.36
1223.51
1116.01
3.58
6.15
‐0.70
8.02 0.72 28.21 1.62 1192.08 2.99 4.84
4.40 0.47 15.50 1.15 1161.28 1.94 2.50
6.24
8.96
0.90
Mg# FeO/MgO CaO/Al2O3 P K&B 1σ Z Km 1σ
T
1σ P Herz
0.2418 3.1352
0.7633
4.18 0.70 14.72 0.25 1171.59 4.09 1.40
0.3846 1.6000
0.7536
6.92 0.52 24.36 0.18 1209.03 3.05 4.42
0.3770 1.6525
0.7536
5.21 0.34 18.34 0.12 1195.64 1.97 3.52
0.3880 1.5777
0.7825
6.59 0.32 23.20 0.11 1212.29 1.86 4.41
0.3904 1.5612
0.7260
8.60 0.32 30.27 0.11 1221.45 1.89 6.13
0.3898 1.5654
0.7300
8.53 0.37 30.02 0.13 1221.79 2.18 6.03
0.3880 1.5770
0.7287
8.73 0.38 30.72 0.13 1222.75 2.23 6.15
0.3883 1.5753
0.7304
8.96 0.44 31.53 0.15 1223.51 2.57 6.15
0.3824 1.6149
0.7428
8.05 0.41 28.33 0.14 1219.02 2.40 5.51
0.3918 1.5526
0.7615
7.14 0.38 25.12 0.13 1214.17 2.23 4.87
0.3528 3.2703
0.6990
4.92 0.50 17.30 1.88 1128.71 1.62 2.94
0.3592 3.1801
0.7201
8.10 1.50 28.54 5.16 1121.19 4.50 3.19
0.3557 3.2287
0.7053
7.95 1.60 27.98 5.56 1125.59 4.83 3.12
0.3500 3.3103
0.6880
8.59 1.70 30.24 5.97 1122.10 5.20 3.29
0.3745 2.9771
0.7329
4.40 0.70 15.62 2.55 1131.51 2.19 2.32
0.3543 3.2484
0.8024
3.10 0.80 10.92 2.91 1116.29 2.52 0.37
0.3603 3.1652
0.8311
0.90 0.50 3.15 1.93 1116.01 1.67 ‐0.70
0.4547 2.1374
0.7585
3.41 0.20 12.00 0.63 1170.63 0.53 2.35
0.4661 2.0418
0.7860
4.30 0.40 15.02 1.48 1169.59 1.25 2.45
2
3
Hansen and Grönvold (2000). All iron as FeO; Samples a,e,h,i,51,57; Samples 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.
1
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 Results
FeO1
15.05
11.33
11.80
11.94
11.74
11.75
11.82
11.86
11.89
11.70
14.88
15.36
15.53
15.79
14.32
15.3
14.37
13.53
13.68
48.79 14.17 2.73 13.35 0.22 6.19 10.58 2.35 0.52 0.37
51.87 15.93 3.72 15.79 0.34 7.57 12.12 2.84 0.72 0.83
46.31 12.02 2.04 11.33 0.04 4.54 8.93 1.26 0.37 0.17
TiO2
3.38
2.04
2.10
2.16
2.10
2.13
2.15
2.12
2.26
2.14
3.41
3.72
3.57
3.51
3.45
3.51
3.48
2.11
2.52
Average
Max
Min
Al2O3
12.91
14.89
15.01
15.49
15.93
15.86
15.90
15.87
15.60
15.57
12.79
12.97
13.1
12.98
12.84
12.4
12.02
13.58
13.46
SiO2
49.46
46.31
48.98
48.38
47.58
47.55
47.50
47.39
47.50
47.91
49.91
48.09
49.5
49.34
49.31
50.1
49.63
51.87
50.66
Sample ID Locality
51
Upp
54
Upp
57
Upp
62
Upp
63
Upp
64
Upp
66
Upp
67
Upp
69
Upp
71
Upp
a
Kverk
b
Kverk
c
Kverk
d
Kverk
e
Kverk
f
Kverk
g
Kverk
h
Kverk
i
Kverk
Table 4. Summary of all basaltic glass data from Upptyppingar and Kverkfjoll.
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
70
68
60 62 63
58 61
64
57
65
54 67 66
50 51
52
71
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
71
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.
72
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.
73
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 341492 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.
74
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
75
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 crustmantle 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
76
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).
77
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.
78
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 crustmantle boundary with geodetic techniques (Sturkell et al., 2006; de Zeeuw-van Dalfsen et
al., 2005).
79
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.
80
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
81
troposphere and, possibly, the lower stratosphere as in the 1783 Laki eruption. The latter
vented 15.1 km3 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.
82
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.
84
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
85
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.
Values for Vp in the upper crust (0-2 km) vary from less than 4.0 km s-1 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.
86
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.
87
Figure 23: Variation diagrams of unfiltered whole rock and glass database.
88
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.
89
90
49
0.61
1.84
1.63
2.13
14.48
14.57
13.3
13.1
12.06
11.78
14.43
12.1
Al2O3
0
0.00
0
0.7
1.19
3.32
1.42
0.4
Fe2O3
8.91
11.53
13.8
11.8
11.47
6.05
10.86
1.4
FeO
12.85
12.35
11.2
9.8
9.30
11.99
11.13
1
CaO
12.07
7.99
5.82
14.4
14.45
11.06
7.04
0.09
MgO
0.17
0.19
0.25
0.17
0.17
0.18
0.20
0.05
MnO
0.05
0.21
0.54
0.18
0.34
0.69
0.41
3.5
K2O
1.5
2.22
2.73
1.79
2.00
2.02
2.39
4.9
Na2O
0.01
0.21
0.41
0
0.20
0.24
0.24
0.04
P2O5
99.19
99.55
100.14
99.92
100.01
98.74
99.10
96.00
Total
Table 5. Major element compositions for neovolcanic zone crustal whole rocks and glasses, and calculated parent magmas.
48.54
46.5
Most Mafic Glass
3.09
47.05
Calculated Parent Magma
(Ol Fo88)
Korenga and Kelleman (2000)
(Ol Fo88)
Most Felsic Glass
48.45
1.48
49.78
Most Mafic WR
Average Glass
1.78
48.84
Average WR
0.22
72.3
Most Felsic WR
TiO2
SiO2
Oxide (wt%)
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.
91
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
Ra = [αfgρf2cpfkb(T1-T0)]/μλm
Eq. 1
is used where k is the thermal conductivity of water, and b is the thickness of the section
(3 km). By using the values of Ra = 4π2, density of water (ρf) = 1000 kg/m3 for, thermal
expansivity of water (αf) = 10-3 K-1 for, viscosity (μ) = 1.33 x 10-4 Pa s, heat capacity of
water (cpf) = 4.2 x 103 JKg-1K-1, thermal conductivity of the rock (λm) = 3.3 W m-1 K-1,
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
92
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)
T = To + (qo/k)(z – zo) – (ρH/2k)(z – zo)2
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
93
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
94
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
95
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.
96
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
97
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 crustmantle 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.
98
Figure 25: Calculated gravity Anomalies for Icelandic Rift Zone.
99
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
100
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
KT(T,P) = KT(T){1-f(5-3K΄)}(1+2f)5/2
Eq. 4
K΄ = (dKT/dP)T
Eq. 5
where
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
α(T, P) = α(T)[ρ(P)/ρo] −δ T
Eq. 7
The shear modulus at elevated pressure and temperature G(T,P) is found by
G(T,P) = G(T)(1+2f)5/2{1-f[5-3G΄KT(T)/G(T)]}
Eq. 8
P wave velocity VP is then found by
VP =
( K S + 4 / 3G ) / 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
Mineral
Pl(h)
Amph(DHP)
Amph(DHP)
Amph(DHP)
Cc(AE)
Opx(HP)
Cpx(HP)
spss
mic
ilm
wt %
24.28
27.77
10.37
9.45
1.26
3.79
17.17
0.46
1.58
3.87
vol %
27.71
26.66
10.16
9.91
1.04
3.28
16.4
0.35
1.95
2.55
1800 bars
Mineral
Bio(HP)
Amph(DHP)
Amph(DHP)
Pl(h)
Pl(h)
Cc(AE)
Opx(HP)
Cpx(HP)
spss
q
ilm
wt %
2.61
23.96
19.65
12.7
19.21
1.17
4.16
10.03
0.43
2.21
3.87
vol %
2.66
22.86
19.14
14.22
22.14
1.04
3.48
9.06
0.32
2.59
2.49
3600 bars
Mineral
Bio(HP)
Amph(DHP)
Pl(h)
Opx(HP)
Cpx(HP)
spss
ilm
CO2
wt %
2.62
43.59
34.31
5.25
9.51
0.35
3.87
0.5
vol%
2.62
41.52
38.5
4.39
8.46
0.26
2.47
1.8
5400 bars
Mineral
Bio(HP)
Amph(DHP)
Pl(h)
Opx(HP)
Cpx(HP)
spss
ilm
H2O
CO2
wt % 2.61
5.48
33.45
20.1
32.97
0.25
3.87
0.78
0.5
vol% 2.61
5.16
37.57
17.04
29.35
0.18
2.46
3.89
1.73
600 bars
Mineral
Pl(h)
Amph(DHP)
Cc(AE)
Opx(HP)
Cpx(HP)
spss
ab
mic
q
ilm
wt %
12.15
49.26
1.23
3.85
13.07
0.46
5.37
1.58
9.17
3.87
vol % 13.77
47.56
1.03
3.28
12.19
0.34
6.42
1.94
10.93
2.53
1200 bars
Mineral
Pl(h)
Amph(DHP)
Cc(AE)
Opx(HP)
Cpx(HP)
spss
ab
mic
q
ilm
wt %
12.75
49.14
1.19
3.92
13.18
0.46
5.05
1.58
8.86
3.87
vol %
14.46
47.54
1.04
3.34
12.21
0.34
6.04
1.93
10.56
2.53
2400 bars
Mineral
Bio(HP)
Amph(DHP)
Pl(h)
Opx(HP)
Cpx(HP)
spss
ilm
CO2
wt %
2.62
43.57
34.58
5.25
9.22
0.41
3.87
0.5
vol %
2.62
41.41
38.8
4.34
8.19
0.3
2.46
1.9
3000 bars
Mineral
Bio(HP)
Amph(DHP)
Pl(h)
Opx(HP)
Cpx(HP)
spss
ilm
CO2
wt %
2.62
43.58
34.45
5.25
9.36
0.39
3.87
0.5
vol %
2.62
41.47
38.65
4.36
8.32
0.28
2.46
1.84
4200 bars
Mineral
Bio(HP)
Amph(DHP)
Pl(h)
Opx(HP)
Cpx(HP)
spss
ilm
H2O
CO2
wt %
2.61
27.43
34.93
11.56
18.47
0.32
3.87
0.33
0.5
vol %
2.61
26.03
39.14
9.76
16.42
0.23
2.46
1.58
1.76
4800 bars
Mineral
Bio(HP)
Amph(DHP)
Pl(h)
Opx(HP)
Cpx(HP)
spss
ilm
H2O
CO2
wt %
2.6
9
34.81
18.78
29.46
0.29
3.87
0.71
0.5
vol %
2.61
8.48
39.01
15.86
26.2
0.21
2.46
3.45
1.74
6000 bars
Mineral
Bio(HP)
Pl(h)
Opx(HP)
Cpx(HP)
spss
ilm
usp
H2O
CO2
wt %
2.6
33.47
20.79
36.88
0.2
2.13
2.55
0.89
0.5
vol %
2.6
37.52
17.69
32.79
0.14
1.36
1.65
4.54
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 912 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 Pwave 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)
aqz
hAb
lAb
an
or
hb
daph
ep
mt
cc
0
7
14.6
0
0
2.4
5
40
0
1
30
1
7
14.6
0
0
2.4
5
40
15
1
15
2
8
17.6
0
0
2.4
20
30
15
1
6
3
14
16
0
7
2.4
34
4.6
15
4
3
4
12
14.6
0
8
2.4
46
0
10
4
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.
111
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.
112
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.
113
Vp
6
6.2
6.4
6.6
6.8
7
0
500
P (bars)
1000
1500
2000
2500
Figure 30: Variation in predicted P-wave velocity from 0-2400 bars calculated using
Perple_X for Model 2.
114
Vp (km /s)
6.2
6.4
6.6
6.8
7
7.2
0
2
4
6
Depth (km)
8
10
12
14
16
18
20
Figure 31: Variation in predicted P-wave velocity through the crust using the Hacker
calculator for Model 2.
115
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
116
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.
117
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.
119
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
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
122
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
123
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.
124
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-plagcpx 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 crustmantle 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.
128
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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
SiO2
47.6
Al2O3
15.9
TiO2
1.23
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K2O
P2O5
TOTAL
0
10.08
0.17
9.18
12.9
1.8
0.1
0.11
99.070
P Kb
Plag Projection
Qtz
Plag
Cpx
Qtz
Results
Ave 6
P Kb
Cpx
42.78
51.89
5.33
68.66
28.43
2.92
1σ6
0.42
Ol
Cpx
Qtz
Plag
Cpx
Qtz
Range 6
1.17
Predicted3
28.93
30.80
32.71
34.67
36.67
38.71
40.80
42.94
45.13
47.38
49.67
52.03
54.44
71.45
68.83
66.15
63.41
60.61
57.75
54.82
51.82
48.75
45.61
42.39
39.10
35.72
-0.38
0.37
1.14
1.92
2.72
3.54
4.38
5.24
6.12
7.01
7.93
8.88
9.84
61.07
62.26
63.47
64.69
65.93
67.18
68.46
69.74
71.05
72.37
73.71
75.07
76.45
39.14
37.54
35.91
34.27
32.61
30.92
29.21
27.48
25.72
23.94
22.14
20.31
18.46
-0.21
0.20
0.62
1.04
1.46
1.90
2.33
2.78
3.23
3.68
4.14
4.61
5.09
ΔPlag-Ol 7
0.30
Slope4
Intercept4
R2
P Calc 5
0.47
-13.36
0.9985
6.77
-0.34
24.25
0.9985
6.82
1.17
0.70
0.9985
6.95
0.78
-47.49
0.9995
6.06
-0.58
22.86
0.9995
6.36
2.27
0.62
0.9995
7.23
Sample2
0.001
1
2
3
4
5
6
7
8
9
10
11
12
Ol Projection
Ol
6.70
1
Glass analysis from Schiellerup (1995)
2
Projected mineral components for sample 1.7.4 calculated using procedure of Tormey et al. (1987)
Projected mineral components calculated at the pressures listed in Column 3 using procedure of
Tormey et al. (1987)
3
4
Slope and intercept from regression of projected mineral components (Columns 5-10) versus P (Column 3)
Pressures calculated from regressions and projected mineral components for sample 1.7.4.
6
Average, 1σ, and range calculated from six pressures obtained by regression
7
Difference between average P calculated for Plag projection and average P calculated for Ol projection.
5
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 CaOMgO-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
Continued
154
48.43
15.61
0.44
7.32
0.11
8.17
12.78
1.32
0.02
0.05
98.23
0.61
0.72
0.78
1.38
Min
Average
22
48.78
14.34
2.13
11.97
0.20
7.30
12.03
2.35
0.29
0.25
99.64
0.52
1.67
0.84
2.64
49.40
16.00
2.96
13.40
0.22
8.42
12.80
2.79
0.46
0.36
100.66
0.59
2.17
0.89
3.25
Max
NVZ
Herdubreid (Hb)
Moore and Calk, 1991
48.30
13.00
1.31
10.60
0.17
6.17
11.10
1.34
0.13
0.14
99.12
0.45
1.26
0.79
1.77
Min
Average
3
50.18
14.48
2.06
12.12
0.11
6.26
10.88
2.41
0.47
0.22
99.18
0.48
2.24
0.75
2.88
50.69
15.62
3.31
15.59
0.15
7.55
12.15
2.68
0.75
0.38
99.86
0.57
3.92
0.78
3.43
Max
49.82
12.30
1.41
10.36
0.04
3.98
9.00
2.22
0.31
0.12
97.85
0.31
1.37
0.73
2.57
Min
NVZ
Burfell (Bu)
Hansen and Gronvold, 2000
Average
19
48.10
14.23
2.10
11.39
0.17
7.53
12.67
2.24
0.21
0.28
98.93
0.54
1.52
0.89
2.45
49.57
15.00
2.26
11.92
0.18
8.28
13.72
2.48
0.28
0.32
99.95
0.57
1.74
1.00
2.71
Max
NVZ
Blafjall Ridge (Bf)
Schiellerup, 1995
46.98
13.64
1.89
10.48
0.13
6.84
11.92
1.96
0.16
0.21
98.19
0.51
1.33
0.80
2.14
Min
I retain the locality names given in the original papers. However, following the suggestion of P. Meyer (personal communication, 2006), I
have used reported values of latitude and longitude to assign samples from Bardabunga/Grimsvötn and from Katla/Hekla to individual
volcanic centers (Table 3). Also, I have grouped together samples belonging to the Hengill volcanic system (eg. samples from Hengill,
Maelifell and Midfell). It seems likely that some samples from Thingvellir also belong to this system, though I have not grouped these with
other Hengill samples.
49.98
16.66
0.79
9.16
0.21
10.56
15.03
1.95
0.10
0.20
100.09
0.71
1.12
0.91
2.05
Max
NVZ
Theistareykir (Th)
Slater et al. 2001
sample_id Average
n
17
SiO2
49.04
Al2O3
16.29
TiO2
0.59
FeOT
8.02
MnO
0.15
MgO
9.82
CaO
13.76
Na2O
1.59
K2O
0.05
P2O5
0.15
TOTAL
99.40
Mg#
0.69
FeO/MgO
0.82
CaO/Al2O3
0.84
Na+K
1.64
Zone
Locality
reference
Table 9. Supplemental Data. Sample Localities and Summary of Glass Compositions
Appendix C
Supplemental Data
49.23
15.71
1.77
11.20
0.18
8.40
12.72
2.23
0.16
0.17
100.30
0.58
1.42
0.81
2.38
Max
NVZ
Halar (Ha)
Schiellerup, 1995
sample_id Average
n
5
SiO2
48.76
Al2O3
15.29
TiO2
1.71
FeOT
10.90
MnO
0.15
MgO
8.14
CaO
12.25
Na2O
2.13
K2O
0.14
P2O5
0.15
TOTAL
99.61
Mg#
0.57
FeO/MgO
1.34
CaO/Al2O3
0.80
Na+K
2.26
Zone
Location
Reference
48.42
14.91
1.68
10.60
0.12
7.86
12.03
2.00
0.11
0.14
98.69
0.56
1.28
0.79
2.14
Min
Average
4
47.15
14.62
2.43
12.22
0.18
7.58
11.52
2.45
0.16
0.29
98.60
0.53
1.61
0.79
2.61
47.77
15.12
2.45
12.36
0.19
7.77
11.78
2.96
0.18
0.33
99.01
0.53
1.66
0.82
3.13
Max
NVZ
Seljahjalli (Se)
Schiellerup, 1995
46.68
14.30
2.41
12.00
0.17
7.44
10.92
2.26
0.14
0.27
98.21
0.52
1.59
0.74
2.43
Min
Average
8
51.21
13.63
2.06
13.40
0.20
5.93
10.20
2.39
0.48
0.15
99.63
0.44
2.34
0.75
2.87
52.68
14.72
2.81
15.54
0.30
7.73
12.40
3.20
0.64
0.27
100.29
0.56
3.17
0.84
3.64
Max
49.88
12.80
1.59
11.02
0.14
4.84
8.90
0.77
0.34
0.00
99.06
0.36
1.43
0.64
1.34
Min
NVZ
Askja (As)
Sigurdsson&Sparks, 1981
Average
4
49.24
14.52
1.08
9.22
0.07
8.88
14.32
2.03
0.10
0.11
99.56
0.63
1.06
0.99
2.13
49.80
15.92
1.23
9.98
0.12
10.50
14.92
2.19
0.13
0.15
99.95
0.69
1.24
1.10
2.32
Max
48.58
13.57
0.93
8.44
0.01
8.00
13.46
1.81
0.05
0.10
99.06
0.59
0.82
0.89
1.86
Min
NVZ
Hrimalda (Hr)
Hansen and Gronvold, 2000
Table 9 continued.
Continued
155
50.35
16.58
1.19
10.25
0.16
9.98
14.71
2.35
0.11
0.14
101.15
0.70
1.33
0.96
2.42
Max
48.90
14.28
0.37
7.58
0.11
7.73
13.55
1.84
0.05
0.01
98.81
0.57
0.77
0.82
1.93
Min
NVZ
Gigoldur (Gi)
Hansen and Gronvold, 2000
sample_id Average
n
12
SiO2
49.45
Al2O3
15.77
TiO2
0.80
FeOT
8.63
MnO
0.14
MgO
8.84
CaO
13.93
Na2O
2.10
K2O
0.08
P2O5
0.07
TOTAL
99.80
Mg#
0.65
FeO/MgO
0.99
CaO/Al2O3
0.89
Na+K
2.18
Zone
Location
Reference
Average
11
49.92
14.32
1.36
11.40
0.20
7.70
12.34
2.24
0.14
0.13
99.75
0.55
1.54
0.86
2.38
51.00
16.00
2.39
15.87
0.26
8.92
13.90
2.76
0.32
0.27
100.81
0.62
2.81
0.96
3.00
Max
NVZ
Sprengisandur (Sp)
Meyer et al., 1985
48.90
12.45
0.77
9.38
0.15
5.65
10.06
1.90
0.05
0.00
99.14
0.39
1.12
0.68
1.99
Min
Average
10
48.34
15.87
0.99
9.18
0.16
9.71
13.80
1.78
0.07
0.06
99.95
0.65
0.95
0.87
1.85
48.98
16.28
1.11
9.49
0.18
10.56
14.09
1.87
0.09
0.18
100.70
0.68
1.02
0.93
1.96
Max
NVZ
Kistufell (Ki)
Breddam, 2002
47.70
15.10
0.87
8.69
0.15
9.27
13.49
1.71
0.05
0.00
99.22
0.64
0.85
0.84
1.77
Min
NVZ
kverkfjoll (Kv)
Hansen and Gronvold, 2000;
Hoskuldsson et al. 2006
Average Max
Min
9
49.82
51.87
48.09
12.90
13.58
12.02
3.25
3.72
2.11
14.75
15.79
13.53
0.24
0.34
0.04
5.12
6.70
4.54
9.63
10.58
8.93
2.35
2.84
1.26
0.63
0.72
0.43
0.53
0.83
0.21
99.23
100.32
97.98
0.38
0.47
0.35
2.95
3.31
2.04
0.75
0.83
0.69
2.98
3.53
1.69
Table 9 continued
Continued
156
50.60
14.67
4.16
15.12
0.25
8.59
13.60
3.38
0.63
0.44
100.88
0.61
3.01
0.94
4.01
Max
EVZ
Bardabunga (Ba)
Meyer et al., 1985
sample_id Average
n
18
SiO2
49.39
Al2O3
13.54
TiO2
2.21
FeOT
12.83
MnO
0.20
MgO
6.80
CaO
11.50
Na2O
2.52
K2O
0.27
P2O5
0.19
TOTAL
99.46
Mg#
0.49
FeO/MgO
1.94
CaO/Al2O3
0.85
Na+K
2.79
Zone
Location
Reference
47.31
12.27
0.93
9.89
0.17
5.03
9.74
1.84
0.08
0.00
98.46
0.37
1.15
0.74
1.92
Min
Average
11
49.78
13.52
2.33
13.00
0.21
6.66
11.17
2.44
0.29
0.22
99.64
0.47
2.14
0.82
2.73
50.86
14.60
3.76
16.09
0.27
8.67
13.80
3.11
0.56
0.40
100.23
0.61
3.56
0.95
3.67
Max
EVZ
Grimsvotn (Gr)
Meyer et al., 1985
49.06
12.08
0.98
10.09
0.16
4.51
8.74
1.87
0.08
0.00
99.07
0.33
1.16
0.71
1.95
Min
Average
29
49.54
13.54
2.24
12.88
0.20
6.75
11.39
2.48
0.28
0.21
99.50
0.48
2.01
0.84
2.76
50.86
14.67
4.16
16.09
0.27
8.67
13.80
3.38
0.63
0.44
100.88
0.61
3.56
0.95
4.01
Max
47.31
12.08
0.93
9.89
0.16
4.51
8.74
1.84
0.08
0.00
98.46
0.33
1.15
0.71
1.92
Min
EVZ
Bardabunga-Grimsvotn
Compiled
Average
9
49.58
14.95
1.32
9.79
0.11
7.89
12.79
2.26
0.15
0.12
98.96
0.59
1.31
0.85
2.41
50.18
15.88
2.24
12.66
0.18
9.26
14.08
2.70
0.32
0.22
99.90
0.67
2.07
0.90
3.02
Max
48.03
12.93
0.82
7.95
0.01
6.10
10.96
1.82
0.05
0.06
98.00
0.46
0.87
0.80
1.90
Min
EVZ
Veidivotn (Vd)
Hansen and Gronvold, 200
Table 9 continued
Continued
157
51.75
14.38
2.68
13.51
0.26
6.43
11.71
3.12
0.61
0.22
99.71
0.47
2.38
0.87
3.73
Max
EVZ
Veidivotn (Vd)
Mork, 1984
sample_id Average
n
11
SiO2
50.30
Al2O3
13.76
TiO2
2.01
FeOT
12.86
MnO
0.22
MgO
5.80
CaO
10.67
Na2O
2.68
K2O
0.42
P2O5
0.19
TOTAL
98.92
Mg#
0.45
FeO/MgO
2.23
CaO/Al2O3
0.78
Na+K
3.10
Zone
Location
Reference
49.05
13.27
1.79
11.95
0.19
5.20
9.97
2.12
0.26
0.16
97.79
0.43
2.00
0.70
2.48
Min
Average
4
49.63
13.64
1.53
12.73
0.22
7.36
11.32
2.21
0.21
0.15
98.99
0.51
1.79
0.83
2.42
50.46
13.94
1.97
14.11
0.26
9.38
11.73
2.27
0.26
0.20
99.66
0.61
2.07
0.84
2.52
Max
48.00
13.30
0.78
10.56
0.19
6.42
10.86
2.11
0.15
0.05
97.40
0.46
1.13
0.82
2.31
Min
EVZ
Veidivotn (Vd)
Thordason et al., 2003
Average
24
49.92
14.19
1.67
11.69
0.18
6.84
11.58
2.44
0.28
0.16
98.95
0.51
1.81
0.81
2.73
51.75
15.88
2.68
14.11
0.26
9.38
14.08
3.12
0.61
0.22
99.90
0.67
2.38
0.90
3.73
Max
EVZ
Veidivotn (Vd)
Compiled
48.00
12.93
0.78
7.95
0.01
5.20
9.97
1.82
0.05
0.05
97.40
0.43
0.87
0.70
1.90
Min
Average
72
49.75
12.82
3.23
14.72
0.23
5.26
10.01
2.69
0.51
0.34
99.56
0.39
2.84
0.78
3.20
50.51
13.70
4.37
17.69
0.29
6.96
11.70
3.01
0.90
0.73
100.19
0.50
4.31
0.86
3.46
Max
48.81
11.45
2.17
12.17
0.17
4.10
8.81
2.21
0.32
0.19
98.07
0.29
1.75
0.73
2.83
Min
EVZ
Laki (Lk)
Thordason et al., 1996
Table 9 continued
Continued
158
SFZ
Katla (Kt)
Meyer et al., 1985; Lacasse et
al 2007
sample_id Average Max
Min
n
19
SiO2
48.74
50.54
46.60
Al2O3
13.07
13.94
11.60
TiO2
3.47
5.32
1.75
FeOT
14.65
16.99
11.81
MnO
0.23
0.29
0.19
MgO
5.43
7.23
4.03
CaO
10.43
12.36
8.77
Na2O
2.74
3.22
1.50
K2O
0.58
0.98
0.21
P2O5
0.24
0.70
0.00
TOTAL
99.56
101.43
98.56
Mg#
0.40
0.52
0.30
FeO/MgO
2.80
4.12
1.63
CaO/Al2O3
0.80
0.89
0.69
Na+K
3.32
4.02
1.88
Zone
Location
Reference
SFZ
Hekla (Hk)
Meyer et al., 1985; Moune et al,
2007
Average Max
Min
25
47.57
50.00
46.40
13.75
15.90
11.60
3.44
5.27
2.36
13.55
17.09
9.23
0.22
0.36
0.15
5.93
7.73
3.83
10.74
13.60
8.25
2.88
3.88
2.31
0.56
1.29
0.34
0.35
0.95
0.00
99.00
100.35
97.11
0.44
0.57
0.31
2.41
4.04
1.35
0.78
0.89
0.66
3.44
4.73
2.65
50.54
15.90
5.32
17.09
0.36
7.73
13.60
3.88
1.29
0.95
101.43
0.57
4.12
0.88
4.73
46.40
11.60
1.75
9.23
0.15
3.83
8.25
1.50
0.21
0.00
97.11
0.30
1.35
0.66
1.88
49.43
16.18
1.02
9.88
0.14
9.73
14.18
2.18
0.04
0.01
100.48
0.65
1.14
0.94
2.22
Max
Average
3
48.80
15.38
0.92
9.64
0.13
9.15
13.70
2.07
0.02
0.00
99.83
0.63
1.06
0.89
2.09
Min
Average
44
48.08
13.45
3.45
14.02
0.22
5.72
10.60
2.82
0.57
0.30
99.24
0.42
2.58
0.79
3.39
Max
WVZ
Langjokull (La)
Meyer et al., 1985
SFZ
Hekla-Katla
Compiled
47.69
14.80
0.82
9.46
0.11
8.68
12.95
2.00
0.01
0.00
99.00
0.61
0.97
0.80
2.02
Min
Table 9 continued
Continued
159
50.00
15.30
2.68
14.40
0.25
7.85
12.70
2.68
0.27
0.21
101.34
0.55
2.23
0.91
2.95
Max
WVZ
Hlodufell (Hl)
Moore and Calk, 1991
sample_id Average
n
19
SiO2
49.22
Al2O3
14.30
TiO2
1.87
FeOT
12.46
MnO
0.22
MgO
7.03
CaO
12.17
Na2O
2.36
K2O
0.19
P2O5
0.16
TOTAL
99.97
Mg#
0.50
FeO/MgO
1.79
CaO/Al2O3
0.85
Na+K
2.54
Zone
Location
Reference
48.50
13.10
1.55
11.20
0.20
6.26
11.40
2.15
0.12
0.13
99.07
0.44
1.48
0.81
2.30
Min
Average
12
48.46
15.00
1.76
11.44
0.20
7.74
12.18
2.22
0.28
0.22
99.49
0.55
1.48
0.81
2.49
48.80
15.60
1.94
12.10
0.22
7.99
12.40
2.27
0.31
0.25
100.13
0.56
1.63
0.88
2.56
Max
WVZ
Raudafell (Ra)
Moore and Calk, 1991
48.00
13.90
1.66
11.20
0.19
7.44
11.60
2.17
0.26
0.18
98.95
0.52
1.40
0.76
2.44
Min
Average
30
48.08
14.40
1.97
12.03
0.20
7.65
12.52
2.21
0.18
0.24
99.47
0.53
1.58
0.87
2.39
48.80
15.60
2.80
13.70
0.22
8.62
13.00
2.45
0.32
0.29
100.95
0.58
2.05
0.95
2.77
Max
WVZ
Efstadalsfjall (Ef)
Moore and Calk, 1991
47.40
13.50
1.68
11.20
0.18
6.68
11.60
2.08
0.14
0.20
98.90
0.47
1.30
0.78
2.25
Min
Average Max
Min
10
47.88
48.10
47.70
15.95
16.10
15.70
1.38
1.42
1.34
10.68
10.80
10.60
0.17
0.17
0.16
8.60
8.79
8.39
12.30
12.40
12.20
2.19
2.24
2.17
0.13
0.13
0.12
0.13
0.14
0.12
99.41
99.66
99.24
0.59
0.60
0.58
1.24
1.29
1.21
0.77
0.79
0.76
2.32
2.37
2.29
WVZ
Kalfstindar (Ka)
Moore and Calk, 1991
Table 9 continued
Continued
160
49.39
15.84
2.07
12.17
0.16
9.08
13.89
2.60
0.24
0.20
100.81
0.63
1.60
0.89
2.80
Max
WVZ
Thingvellir (Tg)
Meyer et al., 1985
sample_id Average
n
10
SiO2
48.79
Al2O3
15.09
TiO2
1.49
FeOT
10.87
MnO
0.14
MgO
8.33
CaO
12.82
Na2O
2.35
K2O
0.13
P2O5
0.09
TOTAL
100.09
Mg#
0.58
FeO/MgO
1.32
CaO/Al2O3
0.85
Na+K
2.48
Zone
Location
Reference
48.04
14.00
1.02
9.38
0.10
7.63
11.86
2.07
0.04
0.00
99.20
0.53
1.03
0.78
2.11
Min
Average
89
49.03
14.80
0.85
9.55
0.17
9.32
15.04
1.62
0.04
0.04
100.46
0.63
1.03
1.02
1.66
49.63
15.72
2.14
11.78
0.21
9.99
15.54
2.15
0.18
0.28
101.45
0.67
1.78
1.07
2.30
Max
47.97
14.15
0.71
8.53
0.15
6.62
13.33
1.47
0.02
0.01
99.73
0.50
0.88
0.87
1.49
Min
WVZ
Midfell (Mi)
Gurenko and Sobolev, 2006
Average
4
48.80
14.60
0.92
8.68
0.15
9.37
14.36
1.87
0.03
0.05
98.82
0.66
0.93
0.98
1.90
49.06
14.90
1.00
8.95
0.15
9.83
14.80
1.96
0.03
0.06
99.22
0.68
1.02
1.03
1.99
Max
WVZ
Maelifell (Ma)
Hansteen, 1991
48.63
14.28
0.85
8.41
0.14
8.75
13.77
1.81
0.02
0.05
98.65
0.64
0.86
0.94
1.84
Min
Average
103
48.77
14.04
2.11
12.12
0.20
6.78
12.16
2.49
0.32
0.27
99.27
0.50
1.88
0.87
2.81
53.40
15.70
3.43
15.30
0.29
9.74
15.50
3.29
0.87
1.03
101.03
0.65
4.24
1.05
4.07
Max
WVZ
Hengill (Hg)
Tronnes, 1990
47.10
12.90
0.76
8.75
0.14
3.35
7.83
1.54
0.01
0.03
97.55
0.30
0.94
0.60
1.55
Min
Table 9 continued
Continued
161
WVZ
Hengill complex (Hg)
Compiled
49.00
15.00
3.09
13.90
0.25
6.98
11.80
2.84
0.54
0.41
100.55
0.50
2.37
0.85
3.28
Max
RP
Geitafell (Ge)
Moore and Calk, 1991
sample_id Average Max
Min
Average
n
206
12
SiO2
48.89
53.40
47.10 48.78
Al2O3
14.43
15.84
12.90 13.86
TiO2
1.51
3.43
0.71
2.56
FeOT
10.88
15.30
8.41 12.97
MnO
0.18
0.29
0.10
0.23
MgO
8.01
9.99
3.35
6.44
CaO
13.48
15.54
7.83 11.49
Na2O
2.10
3.29
1.47
2.65
K2O
0.18
0.87
0.01
0.43
P2O5
0.16
1.03
0.00
0.33
TOTAL
99.81
101.45
97.55 99.74
Mg#
0.56
0.68
0.30
0.47
FeO/MgO
1.47
4.24
0.86
2.03
CaO/Al2O3
0.93
1.07
0.60
0.83
Na+K
2.28
4.07
1.49
3.09
Zone
Location
Reference
48.50
13.00
2.13
12.10
0.21
5.82
11.00
2.50
0.36
0.26
98.89
0.43
1.76
0.79
2.86
Min
Table 9 continued
162