geology and petrogenesis of lavas from an …geology and petrogenesis of lavas from an overlapping...
TRANSCRIPT
1
GEOLOGY AND PETROGENESIS OF LAVAS FROM AN OVERLAPPING SPREADING CENTER: 9°N EAST PACIFIC RISE
By
V. DORSEY WANLESS
A DISSERTATION PRESENTED TO THE GRADUATE SCHOOL OF THE UNIVERSITY OF FLORIDA IN PARTIAL FULFILLMENT
OF THE REQUIREMENTS FOR THE DEGREE OF DOCTOR OF PHILOSOPHY
UNIVERSITY OF FLORIDA
2010
2
© 2010 V. Dorsey Wanless
3
To my family and friends
4
ACKNOWLEDGMENTS
I would like to thank my advisor (Michael Perfit), collaborators (W. Ian Ridley,
Emily Klein, Scott White, Paul Wallace, John Valley, and Craig Grimes) and supervisory
committee (Paul Mueller, Ray Russo, George Kamenov, and David Richardson) for
their mentoring and encouragement throughout this study. Additionally, I would like to
thank the Captain, officers and crew of the R/V Atlantis for all their help during cruise
AT15-17, the MEDUSA2007 Science party (including S. White, K. Von Damm, D.
Fornari, A. Soule, S. Carmichael, K. Sims, A. Fundis, A. Zaino, J. Mason, J. O’Brien, C.
Waters, F. Mansfield, K. Neely, J. Laliberte, E. Goehring, and L. Preston) for their
diligence in collecting data and samples for this study. I thank the ROV Jason II
shipboard and shore-based operations group for their assistance in collecting these
data and HMRG for processing all DSL-120 side scan and bathymetry data collected
during this cruise. Discussions with S. White and A. Goss are gratefully acknowledged.
Thanks to G. Kamenov and the UF Center for Isotope Geosciences for laboratory
assistance and to the Department of Geological Sciences staff for all of their help.
Finally, I thank my friends and family for all their support over the years. This research
was supported by the National Science Foundation (grants OCE-0527075 to MRP and
OCE-0526120 to EMK).
5
TABLE OF CONTENTS
ACKNOWLEDGMENTS ...................................................................................................... 4
page
LIST OF TABLES ................................................................................................................ 8
LIST OF FIGURES .............................................................................................................. 9
ABSTRACT........................................................................................................................ 12
CHAPTER
1 INTRODUCTION ........................................................................................................ 14
2 DACITE PETROGENESIS ON MID-OCEAN RIDGES: EVIDENCE FOR OCEANIC CRUSTAL MELTING AND ASSIMILATION ............................................ 17
Abstract ....................................................................................................................... 17 Introduction ................................................................................................................. 18 Geologic and Tectonic Setting ................................................................................... 22
9°N East Pacific Rise – Overlapping Spreading Center ..................................... 22 Juan de Fuca Ridge Propagating Ridge Tip and Axial Seamount ..................... 24 Galápagos Spreading Center- Extinct OSC or Propagating Ridge Tip? ........... 25
Petrography ................................................................................................................ 26 Geochemical Methods ................................................................................................ 26 Geochemical Results .................................................................................................. 27
Major Element Results ......................................................................................... 27 Trace Element Results......................................................................................... 28 Isotopic Data ........................................................................................................ 29
Petrogenetic Models For High-Silica Lavas ............................................................... 29 Crystal Fractionation ............................................................................................ 30
Rayleigh fractional crystallization .................................................................. 30 Crystal-melt segregation model .................................................................... 33 In situ crystallization calculations .................................................................. 34
Partial Melting (Anatexis) ..................................................................................... 35 Assimilation Fractional Crystallization ................................................................. 37
Discussion ................................................................................................................... 41 Petrogenesis of High Silica Lavas ....................................................................... 41
Geochemical evidence of partial melting ...................................................... 41 The need for crystallization, assimilation and altered crust in dacite
petrogenesis ............................................................................................... 44 Isotopic signature of assimilation .................................................................. 45
AFC Processes and Tectonic Setting ................................................................. 48 Model for Formation of MOR Dacites .................................................................. 49 Relationship of Melt Lens to Dacites at 9°N........................................................ 50
6
Effects of Assimilation on Typical MORB Compositions .................................... 51 Conclusions ................................................................................................................ 52 Acknowledgements..................................................................................................... 53
3 ROLE OF ASSIMILATION IN THE PETROGENESIS OF LAVAS ON MID-OCEAN RIDGES INFERRED FROM Cl, H2O, CO2 AND OXYGEN ISOTOPE VARIATIONS .............................................................................................................. 79
Introduction ................................................................................................................. 79 Geologic Setting ......................................................................................................... 81 Analytical Methods and Results ................................................................................. 83
Major and Trace Elements................................................................................... 83 Volatile Elements ................................................................................................. 84 Oxygen Isotope Analyses and Results ............................................................... 85
Discussion ................................................................................................................... 85 Magma Crystallization versus Assimilation-Fractionation-Crystallization
Processes ......................................................................................................... 85 Evidence for Assimilation from Oxygen Isotopes ............................................... 87 CO2 and H2O Degassing, Magma Ascent Rates and Depth of Assimilant ........ 89 Source of Assimilant ............................................................................................ 91 Oceanic Plagiogranites ........................................................................................ 92
Conclusions ................................................................................................................ 93
4 CRUSTAL DIFFERENTIAION AND SOURCE VARIATIONS AT THE 9°N OVERLAPPING SPREADING CENTER; EAST PACIFIC RISE ............................ 102
Introduction ............................................................................................................... 102 Background, Tectonic Setting and Geology of the 9oN OSC .................................. 104
Overlapping Spreading Centers ........................................................................ 104 Tectonic Setting and Previous Studies of 9°N OSC ......................................... 104 9°N OSC Geology .............................................................................................. 106 8°37’ N EPR Deval ............................................................................................. 108
Geochemical Methods .............................................................................................. 109 Geochemical Results ................................................................................................ 110
Geochemistry of East Limb Lavas ..................................................................... 110 Basalts ......................................................................................................... 111 Basaltic andesites/low-P2O5 andesites ....................................................... 111 High-P2O5 andesites.................................................................................... 112 Dacites ......................................................................................................... 112 Isotopic compositions of East Limb lavas ................................................... 113
West Limb Lavas................................................................................................ 113 8°37’ EPR Lavas ................................................................................................ 114
Discussion ................................................................................................................. 115 Shallow-level Processes Involved in the Petrogenesis of Ferrobasalts, FeTi
Basalts and Basaltic Andesites at the 9°N OSC ........................................... 115 Fractional crystallization versus assimilation in the petrogenesis of
andesites .................................................................................................. 118
7
Formation of west limb lavas by fractional crystallization .......................... 119 Where Are the South Tip and West Limb Dacites? .......................................... 120 Composition of the Melt Lens Beneath the East Limb...................................... 122 E-MORB Distribution at 9°N OSC ..................................................................... 123 9°N OSC as a Division in Mantle Components ................................................. 125
Conclusions .............................................................................................................. 126
5 CONCLUSIONS........................................................................................................ 161
LIST OF REFERENCES ................................................................................................. 164
BIOGRAPHICAL SKETCH.............................................................................................. 181
8
LIST OF TABLES
Table
page
2-1 Dacite major and trace element data..................................................................... 55
2-2 Radiogenic isotopes for 9°N OSC lavas................................................................ 57
2-3 Starting compositions for geochemical modeling.................................................. 59
2-4 Partition coefficients for Rayleigh fractional crystallization, partial melting and AFC calculations .................................................................................................... 60
2-5 AFC modeling parameters ..................................................................................... 61
3-1 Geochemical data .................................................................................................. 95
4-1 East limb major element data 9°N OSC .............................................................. 128
4-2 Trace element data 9°N OSC .............................................................................. 136
4-3 West limb major element data 9°N OSC ............................................................. 140
4-4 West limb trace element data 9°N OSC .............................................................. 142
4-5 West limb isotopic data 9°N OSC ........................................................................ 143
4-6 Major and trace element data from 8°37'N EPR ................................................. 144
4-7 Radiogenic isotope ratios 8°37'N EPR ................................................................ 145
9
LIST OF FIGURES
Figure
page
2-1 Bathymetric maps showing the tectonic setting of the MOR dacites ................... 62
2-2 Comparison of major and trace element compositions in MOR high-silica andesites and dacites ............................................................................................ 65
2-3 Bathymetric map of the 9°N OSC showing locations of samples collected during the MEDUSA2007 cruise ............................................................................ 66
2-4 Photographs of MOR high-silica lavas .................................................................. 67
2-5 Major element variations versus MgO (wt.%) for dacites from the 9°N OSC on the East Pacific Rise ......................................................................................... 68
2-6 Variation diagram showing Cl (wt%) versus MgO (wt%) for OSC lavas. ........... 69
2-7 Trace element variations versus Zr (ppm) in 9°N OSC lavas ............................... 70
2-8 Normalized trace element ratio diagrams.............................................................. 71
2-9 Radiogenic isotope ratios of 9°N OSC lavas......................................................... 72
2-10 Elemental variation diagrams showing in situ crystallization and melt-segregation models ................................................................................................ 73
2-11 Partial melting model showing a range of possible parents ................................. 74
2-12 Mantle-normalized diagrams showing results of 1-15% partial melting of an altered basalt .......................................................................................................... 75
2-13 Diagram showing the calculated effects of varying wall rock temperature on incompatible trace element composition (La) during AFC .................................... 76
2-14 Mantle-normalized trace element diagram showing results of the best-fit AFC model ...................................................................................................................... 77
2-15 Cartoon showing a possible scenario for dacite formation on MOR .................... 78
3-1 Bathymetric map of the East Pacific Rise showing the location of the 9°N OSC, the Clipperton and Siqueiros transform faults. ............................................ 96
3-2 H2O (wt%), Cl (wt%), and CO2 (ppm) versus MgO (wt%) for glasses from the 9°N OSC. ........................................................................................................ 97
3-3 H2O/Ce and Cl/K2O ratios versus MgO (wt%) for glasses from the 9°N OSC..... 98
10
3-4 Cl/K2O versus H2O/K2O for glasses from the 9°N OSC ....................................... 99
3-5 δ18O versus MgO for glasses from the 9°N OSC ................................................ 100
3-6 CO2–H2O vapor saturation diagram .................................................................... 101
4-1 Bathymetric map of the northern EPR. ................................................................ 146
4-2 Bathymetric map of the 9°N OSC with the location of rock samples ................. 147
4-3 Bathymetric map of the 9°N OSC with the melt sills ........................................... 148
4-4 Side scan sonar mosaic from data collected on the MEDUSA2007 cruise ....... 149
4-5 FeO versus MgO for glasses collected from the east limb of the 9°N OSC ...... 150
4-6 Major element variations versus MgO (wt%) for glasses collected from the east limb of the 9°N OSC ..................................................................................... 151
4-7 P2O5/TiO2 versus MgO (wt%) for east limb glasses ........................................... 152
4-8 Trace element concentrations versus Zr for glasses .......................................... 153
4-9 Incompatible trace element ratios versus Zr for glasses erupted at the OSC. 154
4-10 Radiogenic isotope ratios showing the variation in sources ............................... 155
4-11 Major element concentrations and ratios versus MgO comparing the east and west limb of the OSC ........................................................................................... 157
4-12 Trace element concentrations versus Zr comparing east and west limb ........... 159
4-13 Primitive mantle normalized diagram showing variations in andesites and basaltic andesites erupted at the OSC. ............................................................. 160
11
LIST OF ABBREVIATIONS AFC assimilation fractional crystallization
Cpx clinopyroxene
EPR East Pacific Rise
GSC Galapagos Spreading Center
HREE heavy rare earth elements
Ilm Ilmenite
JdFR Juan de Fuca Ridge
LLD liquid line of descent
MOR mid-ocean ridge
MORB mid-ocean ridge basalt
N-MORB normal mid-ocean ridge basalt
ODP Ocean Drilling Program
Ol olivine
OSC overlapping spreading center
Plag plagioclase
QFM quartz-fayalite-magnetite
REE rare earth elements
ROV remotely operated vehicle
Sp spinel
12
Abstract of Dissertation Presented to the Graduate School of the University of Florida in Partial Fulfillment of the Requirements for the Degree of Doctor of Philosophy
GEOLOGY AND PETROGENESIS OF LAVAS ON AN OVERLAPPING SPREADING
CENTER: 9°N EAST PACIFIC RISE
By
V. Dorsey Wanless
August 2010
Chair: Michael Perfit Major: Geology
In contrast to relatively homogeneous mid-ocean ridge basalt (MORB)
compositions typically erupted on fast-spreading oceanic ridges, a wide range of rock
types from basalts to dacites have been recovered at overlapping spreading centers
(OSC). This study focuses on the petrogenesis of lavas erupted at the 9°N OSC on the
East Pacific Rise in order to better understand the complex magmatic plumbing system
beneath a ridge discontinuity. Lavas that span the entire compositional range observed
on the global mid-ocean ridge (MOR) system, including basalts, ferrobasalts, FeTi
basalts, basaltic andesites, andesites and dacites have erupted along the eastern,
propagating limb of the OSC. Major and trace element analyses, radiogenic (Pb, Sr, Nd)
and oxygen isotopic ratios, volatile contents (Cl, H2O, CO2) and geochemical modeling
are used to determine the petrogenesis of MORB and genetically related high-silica
magmas.
The formation of high-silica dacites on MOR remains a petrologic enigma despite
eruption on several different ridges. They are characterized by elevated U, Th, Zr, and
Hf; relatively low Nb and Ta; and Al2O3 and K2O concentrations that are higher than
expected from fractional crystallization. Additionally, high Cl and H2O concentrations
13
and relatively low δ18O values in dacitic glasses require contamination from a seawater-
altered component. Extensive petrologic modeling of MOR dacites suggests that
fractional crystallization of a MORB parent combined with partial melting and
assimilation of altered ocean crust can generate magmas with geochemical signatures
consistent with MOR dacites. This suggests that crustal assimilation is a much more
important process on ridges than previously thought and may be significant in the
generation of evolved MORB in general.
Petrologic models indicate that ferrobasalts and FeTi basalts erupting at the OSC
can be explained by low-pressure fractional crystallization of a primitive MORB parent;
however, both fractional crystallization and magma mixing produce intermediate
compositions. Geochemical analyses suggest that there are two distinct populations of
andesites erupted at the OSC. Andesites with high-P2O5 are the most evolved MOR
compositions produced through fractional crystallization. In contrast, low-P2O5 andesites
and basaltic andesites appear to have formed primarily through mixing of ferrobasaltic
and dacitic magmas.
14
CHAPTER 1 INTRODUCTION
Mid-ocean ridges (MOR) are comprised of a series of segments that can be
subdivided at a variety of scales, ranging from tens of meters to hundreds of kilometers
between 1st order discontinuities marked by transform faults (Sempere and Macdonald,
1986; Macdonald et al., 1988). Overlapping spreading centers (OSC) are 2nd order
discontinuities that form between widely spaced transform faults on fast to intermediate
spreading ridges (Macdonald and Fox, 1983; Sempere and Macdonald, 1986, Carbotte
and Macdonald, 1992). These offsets provide both a physical and a geochemical
segmentation of the ridge, which may result from variations in mantle melting and/or
separation of crustal magma reservoirs between segments (e.g. Macdonald et al.,
1988).
Lavas erupted along fast to intermediate spreading centers, such as the northern
East Pacific Rise (EPR), may produce a range of basaltic lavas (e.g., Batiza and Niu,
1992), but they rarely erupt compositions with MgO concentrations <5 wt%. This
relatively limited compositional diversity compared to other tectonic settings is
commonly attributed to shallow-level fractional crystallization of primitive magmas within
an axial magma chamber buffered by relatively frequent recharge with more primitive
melts (Klein, 2005). Additionally, geochemical variations in MORB may result from
variable mantle melting parameters and/or mantle sources (Klein & Langmuir, 1987;
Langmuir et al., 1992).
In contrast, lavas erupted at ridge segment ends, such as an OSC, can have
highly variable compositions compared to a relatively limited range of basaltic
compositions erupted from magmatically robust segment centers (e.g. Christie and
15
Sinton, 1981; Langmuir et al., 1986). These rock types can range from typical MORB
lavas to iron-enriched FeTi basalts, andesites, and high-silica dacites. This
geochemical variability is commonly attributed to lower magma supply and cooler crust
(cold edge effect) at the end of ridge segments, which causes increased magmatic
fractionation prior to eruption (Christie and Sinton, 1981; Sinton et al., 1983; Perfit et al.,
1983; Perfit and Chadwick, 1998; Rubin and Sinton, 2007). While crystal fractionation
is undoubtedly a primary process in magma differentiation at MOR, it may not be the
only process involved in the petrogenesis of evolved lavas on MOR.
Elevated Cl concentrations in many MORB suggest that partial melting and
assimilation of seawater-altered material may be an important, but often overlooked,
process in MOR magmatism (Michael and Schilling, 1989; Michael and Cornell, 1998).
Evidence of these processes at MOR is supported by textural observations in ophiolites,
which show melting of overlying crustal material at the top of the magma chamber
(Coogan et al., 2003). Additionally, experimental results suggest that hydrous partial
melting of altered ocean crust can produce high-silica plagiogranite veins, which are a
small (<2%), but ubiquitous part of the ocean crust (Koepke et al., 2004; Koepke et al.,
2007). Despite the clear evidence of assimilation in these lavas, many models for the
magmatic plumbing system at MORs do not include partial melting and assimilation.
Dacites have been sampled from several different spreading centers; including,
the East Pacific Rise (EPR), the Juan de Fuca Ridge (JdFR), and the Galapagos
Spreading Center (GSC); however, there is no consensus on how these high-silica
lavas form on MOR. To answer these questions, we undertook a 35-day research cruise
(AT15-17) in the Spring of 2007 to the 9°N OSC, during which we surveyed 200 sq. km.
16
with DSL-120A for side scan sonar, mapped and sampled with the ROV Jason II (~7000
digital photographs, recovery of >280 rock samples, sampling of hydrothermal vent
waters and biota); and took photographs with the WHOI TowCam (~10,000 digital
photographs, CTD and MAPR data), which constitiute one of the most detailed data
sets from an OSC. The compositions of lavas recovered from the OSC exhibit
remarkable diversity, ranging from basalt to dacite: 33% of OSC lavas have SiO2 > 52
wt.%, compared to ~3% for ocean ridge basalts worldwide. This study utilizes this data
to explore the roles of crystal fractionation, partial melting and assimilation in the
petrogenesis of high-silica lavas on MOR. In addition, geochemical data are used to
determine which processes are involved in the formation of the range of compositions
(basalts, ferrobasalts, FeTi basalts, basaltic andesites, andesites and dacites) erupted
at the 9°N OSC to better understand the anatomy of a 2nd order MOR discontinuity.
17
CHAPTER 2 DACITE PETROGENESIS ON MID-OCEAN RIDGES: EVIDENCE FOR OCEANIC
CRUSTAL MELTING AND ASSIMILATION
Abstract
While the majority of eruptions at oceanic spreading centers produce lavas with
relatively homogeneous mid-ocean ridge basalt (MORB) compositions, the formation of
tholeiitic andesites and dacites at mid-ocean ridges (MOR) is a petrologic enigma.
Eruptions of MOR high-silica lavas are typically associated with ridge discontinuities and
have produced regionally significant volumes of lava. Andesites and dacites have been
observed and sampled at several different locations along the global MOR system;
including propagating ridge tips at ridge-transform intersections on the Juan de Fuca
Ridge and eastern Galápagos spreading center, and at the 9°N overlapping spreading
center on the East Pacific Rise. Despite the formation of these lavas at various different
ridges, MOR dacites show remarkably similar major element trends and incompatible
trace element enrichments, suggesting that similar processes are controlling their
chemistry. Although most geochemical variability in MOR basalts is consistent with low-
pressure fractional crystallization of various mantle-derived parental melts, our
geochemical data from MOR dacitic glasses suggest that contamination from a
seawater-altered component is important in their petrogenesis. MOR dacites are
characterized by elevated U, Th, Zr, and Hf, low Nb and Ta concentrations relative to
the rare earth elements (REE) and Al2O3, K2O, and Cl concentrations that are higher
than expected from low-pressure fractional crystallization alone. Petrologic modeling of
MOR dacites suggests that partial melting and assimilation are both integral to their
petrogenesis. Extreme fractional crystallization of a MORB parent combined with partial
melting and assimilation of amphibole-bearing altered crust produces a magma with
18
geochemical signatures consistent with MOR dacites. This supports the hypothesis that
crustal assimilation is an important process in the formation of highly evolved MOR
lavas and may be significant in the generation of evolved MORB in general.
Additionally, these processes are likely to be more common in regions of episodic
magma supply and enhanced magma-crust interaction such as at the ends of ridge
segments.
Introduction
Fast to intermediate oceanic spreading centers typically erupt geochemically
diverse basaltic lavas (e.g. Klein, 2005); however, a much more extensive range of lava
compositions, including ferrobasalts and FeTi basalts as well as rarer high-silica
andesites and dacites have been recovered from several different ridges (Perfit et al.,
1983; Langmuir et al., 1986; Natland et al., 1986; Natland & MacDougall, 1986;
Regelous et al., 1999; Smith et al., 2001). These great variations in compositions are
commonly attributed to low magma supply and/or cooler crust at ridge segment ends, or
the cold edge effect, which promote greater differentiation of magmas prior to eruption
(Christie & Sinton, 1981; Perfit et al., 1983; Sinton et al., 1983; Perfit & Chadwick, 1998;
Rubin & Sinton, 2007).
The formation of highly evolved, silicic magmas in non-ridge settings (e.g. ocean
islands, arc volcanoes, and continental interiors) have been attributed to several
different processes, including crystal fractionation, partial melting of overlying crust,
and/or assimilation of crustal material into an evolving magma chamber. On mid-ocean
ridges (MOR), many studies have documented the dominant role crystal fractionation
plays in magma differentiation (e.g. Clague & Bunch, 1976; Bryan & Moore, 1977;
Byerly, 1980) whereby extensive crystallization of olivine, plagioclase, pyroxene and Fe-
19
Ti oxides leads to the generation of highly evolved melts enriched in SiO2 and depleted
in MgO, FeO and TiO2 (e.g. Juster et al., 1989). However, while crystal fractionation is
undoubtedly a primary process involved in the differentiation of most MORB magmas, it
may not be the only mechanism involved in the formation of high-silica MOR andesites
and dacites.
Partial melting (or anatexis) of basaltic crustal material may produce evolved
compositions, particularly in settings where magma-rock interactions are likely, such as
the top of an axial magma chamber (e.g. Coogan et al., 2003b; Gillis, 2008). This
process has been suggest as the origin for high-silica lavas erupted on many ocean
islands (Iceland; O'Nions & Gronvold, 1973; Sigurdsson & Sparks, 1981; Galapagos
Islands; McBirney, 1993; Socorro Island; Bohrson & Reid, 1997; Bohrson & Reid, 1998)
and may explain the formation of high-silica lavas in back arc settings, most recently
along the Lau Spreading Center (e.g. Kent et al., 2002) and Manus Basin (Sinton et al.,
2003); although the presence of a subduction zone makes this tectonic setting much
more complicated. Evidence from ophiolites suggests that the top of the axial magma
chamber on MOR is a dynamic boundary where magmas may interact with and melt
different layers of crustal material; including both gabbros and sheeted dikes (Coogan et
al., 2003b). Recent experimental evidence suggests that partial melting of hydrous
gabbroic rock similar to that in the lower ocean crust can form silicic compositions
(Koepke et al., 2004; Kvassnes & Grove, 2008) and may explain the presence of highly
evolved plagiogranite veins in the ocean crust (Koepke et al., 2004; Brophy, 2009).
Other studies indicate that low degrees of dehydration partial melting of altered basalt,
similar in composition to dikes of the upper ocean crust, can produce dacitic melts
20
(Beard & Lofgren, 1991). These experimental studies suggest that oceanic crust will
begin to melt at temperatures as low as 850° to 900°C and <10% melting of the crust
will yield dacitic or tonalitic melts (Beard & Lofgren, 1991; Koepke et al., 2004;
Kvassnes & Grove, 2008). Kvassnes and Grove (2008) state that mineral pairs
(plagioclase-olivine and plagioclase-augite) similar to oceanic gabbros from the lower
crust will melt quickly and easily at temperatures similar to that of primitive MOR
magmas (1220-1330°C). All of these studies indicate that high level partial melting of
ocean crust can produce high-silica melts on MOR, but the role that this process may
play in the formation of voluminous extrusive silicic lavas on the seafloor has not yet
been assessed.
The compositional variability observed in arc and continental volcanics is
commonly ascribed to the associated processes of assimilation and fractional
crystallization (AFC; e.g. Bowen, 1928; De Paolo, 1981) , but similar processes may
also occur where thickened oceanic crust leads to magma-crust interaction, for
instance, within Icelandic volcanoes (e.g. Nicholson et al., 1991). On smaller scales,
the combined effects of these processes have been observed in ophiolites, where sub-
axial intrusive magmas have been in contact with and have partially melted the
overlying sheeted dikes (Gillis & Coogan, 2002; Coogan, 2003; Gillis, 2008). AFC
processes have also been invoked to explain high Cl concentrations observed in some
MORB magmas (Michael & Schilling, 1989; Michael & Cornell, 1998; le Roux et al.,
2006). During this process, a magma undergoes crystal fractionation, and the resultant
latent heat of crystallization provides the heat needed to partially melt the surrounding
wall rock. These melts are then assimilated into, and homogenized with, the
21
fractionating magma reservoir. AFC processes can produce a wide range of rock types
depending on the initial composition of the intruding magma, the degree of crystal
fractionation, the initial wall rock composition, and the amount of melting and
assimilation.
High-silica compositions are found throughout the ocean crust and are commonly
observed as intrusive or plutonic material. As mentioned above, plagiogranites veins
are a ubiquitous component of the ocean crust and have been observed in ophiolites
(e.g. Pedersen & Malpas, 1984), drill cores from the ocean crust (Casey, 1997; e.g.
Dick et al., 2000; Wilson et al., 2006), and as xenoliths in Icelandic lavas (Sigurdsson,
1977). The origin of these veins remains unclear but two main hypotheses are: 1) partial
melting of gabbroic crust (e.g. Koepke et al., 2004; Koepke et al., 2007; Nunnery et al.,
2008) and 2) extreme crystal fractionation of tholeiitic magmas(Coleman & Donato,
1979; Beccaluva et al., 1999; Niu et al., 2002). There are also many examples of
evolved plutonic rocks from slower spreading centers (e.g. Mid-Atlantic Ridge,
Aumento, 1969), which may suggest AFC or partial melting processes are occurring on
much smaller scales, deeper in the ocean crust.
In this study we examine the geochemistry of high-silica lavas from three different
MOR, including the East Pacific Rise, Juan de Fuca Ridge, and Galapagos Spreading
Center (Figure 2-1) and show they have remarkably similar major and trace element
compositions (Figure 2-2), suggesting that similar sources and processes control their
petrogenesis. More specifically, we examine the roles that crystal fractionation, partial
melting, and AFC may have played in the formation of an exceptional suite of high-silica
lavas from the 9°N overlapping spreading center (OSC) on the East Pacific Rise, and
22
evaluate if these results apply generally to the formation of high-silica lavas on other
MOR. We focus on the petrogenesis of dacites at the 9°N OSC because it is the most
complete and geologically well-constrained data set available but descriptions of the
geologic settings of high-silica lavas in the other environments are important in order to
ascertain the role tectono-magmatic settings may have on their petrogenesis.
Geologic and Tectonic Setting
The MOR system is over 70,000 km long (Macdonald et al., 1991) and the crust
formed at these ridges is overwhelmingly basaltic in nature. High-silica lavas, however,
have erupted on several fast- and intermediate-spreading ridges and are commonly
associated with specific tectonic settings; including propagating ridge tips (Christie &
Sinton, 1981; Fornari et al., 1983; Perfit & Fornari, 1983); OSC (Christie & Sinton, 1981;
Perfit et al., 1983; Sinton et al., 1983; Bazin et al., 2001); regions of ridge-hotspot
interaction (Chadwick et al., 2005; Haase et al., 2005) and at 10°30 N on the East
Pacific Rise near the ridge-transform intersection (Regelous et al., 1999). Below, we
describe the geologic setting of the 9°N OSC and the three other ridges (East Pacific
Rise, Juan de Fuca Ridge, and Galapagos Spreading Center) where highly evolved
lavas have erupted, to elucidate the relationship of magmatism to different MOR
tectono-magmatic environments.
9°N East Pacific Rise – Overlapping Spreading Center
The 9°N OSC (Figure 2-1a) is located on the East Pacific Rise between the
Clipperton and Siqueiros transform faults. It consists of two north-south trending ridges
that overlap by ~27 km and partly enclose a large overlap basin (Sempere &
Macdonald, 1986) The limbs are separated by ~8 km (Singh et al., 2006). The eastern
23
limb is propagating to the south into older crust at a rate of ~42 km Myr-1 (Carbotte &
Macdonald, 1992).
The 9°N OSC has been the focus of several geophysical studies (Detrick et al.,
1987; Harding et al., 1993; Kent et al., 1993; Kent et al., 2000; Bazin et al., 2001; Dunn
et al., 2001; Tong et al., 2002); including the first MCS 3D survey of a MOR (Kent et al.,
2000) and a 3D seismic refraction study (Dunn et al., 2001). These studies resulted in
the first 3D image of a subsurface magma chamber along a MOR, which showed a
shallow melt lens lies beneath both limbs of the OSC with an anomalously large melt
lens in the interlimb region, north of the overlap basin. This suggests that the region
currently has an unusually high magma supply rate for a ridge segment end (Kent et al.,
2000).
High-silica andesites and dacites were recovered from the eastern limb during the
Medusa2007 cruise (AT15-17) using the ROV Jason2 (Wanless et al., 2008; White et
al., 2009). Several high-silica lavas were also recovered from this area during dredging
operations in the late 1980’s (Langmuir et al., 1986). The siliceous lavas are primarily
confined to the northern section of the neo-volcanic zone on the eastern, propagating
limb, along the eastern edge of the melt lens (Figure 2-3). Morphologically, the dacites
form large individual bulbous to elongate pillows that can be several meters in diameter
(Figure 2-4). The pillows are highly striated and have a coarse bread crust surface
texture. Typically, the pillows are stacked into mounds, which can be several meters
high or constructional domes. Dacites largely occur in two regions: as a nearly linear
pillow mound in the center of the east limb neo-volcanic zone and large, elongate pillow
24
lavas on the flanks of the axial graben (Figure 2-3). Their axial and near-axial positions,
low sediment cover and unaltered nature suggest they are relatively young.
Juan de Fuca Ridge Propagating Ridge Tip and Axial Seamount
The Cleft segment is the southernmost segment of the Juan de Fuca Ridge
(Figure 2-1b). It terminates at 44°27’N at a ridge-transform intersection, where it
intersects and overlaps the Blanco Transform Zone (Embley et al., 1991; Embley &
Wilson, 1992; Smith et al., 1994). This intersection is characterized by a series of
curved ridges that overshoot the Blanco Transform Zone onto the older Pacific plate
(Stakes et al., 2006).
High-silica andesites and dacites comprise two small constructional domes on the
Pacific plate, where the axial ridge intersects, and is believed to propagate past the
Blanco Transform Zone into the older ocean crust (~6.3 ma) that was created at the
Gorda Ridge (Embley & Wilson, 1992; Stakes et al., 2006). The domes are ~20 to 30 m
high and 200 to 500 m in diameter and were sampled using rock core and the ROV
Tiburon during research cruises in 2000 and 2002 (Cotsonika, 2006). High-resolution
bathymetric maps show there are numerous other constructional domes in the region
but they have not yet been sampled, although we surmise that they are also composed
of high-silica lavas. Rare andesites have also been recovered within the axis and along
the bounding faults of the southern Cleft segment (Stakes et al., 2006).
Seismic studies of the southern Juan de Fuca Ridge indicate the presence of an
axial magma chamber beneath most of the Cleft segment (Canales et al., 2005).
However, an axial magma chamber reflector is absent south of 44°38’N where the high
silica lavas were recovered, suggesting the presence of small melt volumes resulting
from weak melt supply to the ridge-transform intersection. Zircon thermochronology and
25
U-series data indicate that the dacites erupted less than 30 ka ago (Schmitt et al.,
submitted).
The Axial segment of the Juan de Fuca Ridge is a second order ridge segment
that currently overlies the Cobb hot spot, which has produced a chain of seamounts
trending NW away from the ridge axis (Chadwick et al., 2005). The Juan de Fuca Ridge
is migrating NW at a rate of 3.1 cm/yr and has been situated above the Cobb hot spot
for the last ~0.2 to 0.7 Myr, creating a large on-axis seamount, known as Axial
seamount (Karsten & Delaney, 1989).
Axial seamount is the largest feature on this segment of the Juan de Fuca Ridge
and has a large summit caldera (Embley et al., 1999) underlain by a large seismically-
imaged axial magma chamber (West et al., 2001). It has two prominent rift zones,
extending to the north and south, which create bathymetric highs. Extensive sampling
of the main edifice shows that it is composed of moderately evolved and slightly
enriched MORB (Chadwick et al., 2005). The rift zones have linear ridges that appear
to accommodate extensive diking from the main caldera system (Chadwick et al., 2005).
Rare high-silica andesites sampled by three rock cores are located east of the northern
rift zone and may be associated with dike propagation from the main axial magma
chamber into older ridge crust.
Galápagos Spreading Center- Extinct OSC or Propagating Ridge Tip?
High-silica lavas were sampled at the eastern end of the Galápagos Spreading
Center at ~85° W. The area was extensively studied though dredging and Alvin
exploration in the early 1980s (Fornari et al., 1983; Perfit & Fornari, 1983; Perfit et al.,
1983; Embley et al., 1988). The evolved lavas erupted within the axial valley and along
26
the axis-bounding faults of the Galápagos Spreading Center approximately 20 km east
of the ridge-transform intersection with the Inca transform fault (Figure 2-1c).
Bathymetric data reveal two curved ridges surrounding a depression within this region,
which has been interpreted as an old, small, extinct OSC or deviation from axial linearity
(Perfit et al., 1983; Embley et al., 1988) that has been rifted away from the neo-volcanic
zone. Most of the evolved lavas (~63% SiO2) at the Galápagos Spreading Center were
found off-axis along the bounding faults associated with the southern portion of the
extinct OSC.
Petrography
The 9°N OSC dacites are glassy and predominantly aphyric, with sparse
microphenocrysts and very rare, small phenocrysts of plagioclase and clinopyroxene.
The small phenocrysts of plagioclase commonly have resorbed edges and sieve
textures. Several of the samples contain small clots of basaltic xenoliths comprised of
subophitic plagioclase and clinopyroxene surrounded by dacitic or andesitic glass.
Geochemical Methods
Glass chips from the outer rims of 18 dacites collected at the 9°N OSC were
analyzed on a JEOL 8900 Electron Microprobe for major element concentrations at the
USGS in Denver, Colorado (Table 2-1). Eight to ten individual points were analyzed per
sample. USGS mineral standards were used to calibrate the microprobe and secondary
normalizations were done to account for instrument drift using the JdF-D2 glass
“standard” (Reynolds, 1995), University of Florida in-house standard ALV 2392-9 from
the East Pacific Rise (Smith et al., 2001) and USGS standard dacite glass GSC (for
more detail on methods see (Smith et al., 2001)). The probe diameter during glass
analyses was 20 μm, and an accelerating voltage of 15keV and a beam current of 20nA
27
were used. High-precision chlorine and potassium concentrations were also
determined by microprobe on seven of the dacites using 200-second peak/100 second
background counting times.
Small glass fragments (10-50 mm) were handpicked, avoiding microphenocrysts
and alteration, cleaned, and dissolved for trace element and isotope analyses following
methods described in (Goss et al., 2010). Fourteen dacites from the 9°N OSC were
analyzed at medium resolution for trace element concentrations on a high precision
Element2 Inductively Coupled Plasma Mass Spectrometer (ICP-MS) at the University of
Florida (Table 2-1). Radiogenic isotope ratios (Pb, Sr, and Nd) were determined for 10
dacites using a Nu-Plasma multi-collector ICP-MS at the University of Florida Center for
Isotope Geoscience (Table 2-2). For a detailed description of sample preparation,
dissolution procedures, standards, and errors, see (Kamenov et al., 2007; Goss et al.,
2010; Goss et al., in prep). External calibration was done to quantify results using a
combination of in-house basalts (ENDV – Endeavour and ALV 2392-9) and USGS
(AGV-1, BIR-1, BHVO-1, BCR-2 and STM-1) rock standards (Kamenov et al., 2007;
Goss et al., 2010).
Geochemical Results
Major Element Results
Major element compositions of the 9°N OSC high-silica andesites and dacites are
presented (along with the trace element abundances) in Table 2-1. The major element
geochemistry of the 9°N OSC lavas is similar to Juan de Fuca Ridge and Galápagos
Spreading Center lavas (Figure 2-2). Here, only data from the 9°N OSC is discussed in
detail but it is important to note that the major and trace element contents and elemental
28
trends in high-silica lavas from all three ridges are similar. New analyses of some of the
high-silica samples previously analyzed and discussed by Perfit et al., (1983; 1999) and
representative samples from the Juan de Fuca Ridge are presented in Supplementary
Data. All high-silica samples from the 9°N OSC appear unweathered with minimal
amounts of Fe-Mn oxide coating and are essentially aphyric.
9°N OSC tholeiitic andesites through dacite samples exhibit increasing SiO2 with
decreasing FeO, TiO2, and MgO (Figures 2-2 and 2-5) with the most differentiated
dacites having ~67 wt% SiO2 and <1 wt% MgO. Al2O3 concentrations in the dacites
(12.9 to 13.3 wt%), however, do not show a large decrease compared to the OSC
basalts (Figure 2-5). The dacites have high incompatible major element concentrations
(K2O> 0.90 wt % and Na2O > 3.4 wt %; Figure 2-5) but low P2O5 concentrations (< 0.26
wt%; Figure 2-5) compared to basalts. Chlorine concentrations in the dacites range
from 0.24 to 0.70 wt% compared to < 0.01 to 0.04 wt% in the OSC basalts (Figure 2-6).
Trace Element Results
9°N OSC dacites are enriched in incompatible trace elements compared to 9°N
OSC basalts (Figure 2-7), the latter having compositions typical of normal, incompatible
trace element-depleted mid-ocean ridge basalts (N-MORB) from the northern East
Pacific Rise. For example, Zr and Hf concentrations in the dacites range from 622 to
1050 ppm and 18 to 25 ppm, respectively (Table 2-1). The dacites also contain high
concentrations of Rb, Ba, U and Th but relatively low Sr and Eu contents. Compared to
East Pacific Rise N-MORB the dacites have relatively flat REE patterns (Figure 2-2). On
mantle-normalized diagrams the dacites have positive Zr, Hf, U, and Th anomalies, and
negative Nb and Ta anomalies (Figure 2-2). Consequently, the dacites also have
29
slightly lower Nb/La and higher Zr/Dy ratios compared to OSC basalts (Figure 2-8) and
Ce/Yb and Nd/Y ratios increase from basalt to dacites (Figure 2-8). Compatible trace
elements are low in the dacites, with Ni concentrations ranging from 9.8 to 4.9. ppm
(Figure 2-7) and Cr concentrations from 13 to 1 ppm. Most incompatible trace elements
have negative correlations with MgO; however, compatible elements (i.e. Ni and Cr) and
Nb/La are positively correlated. U/Nb, Nd/Y and Ce/Yb ratios are negatively correlated
with MgO in the dacites.
Isotopic Data
The 9°N OSC dacites have very limited ranges of Pb, Sr and Nd isotopic
compositions, which lie within the general field of East Pacific Rise MORB lavas (Table
2-2; Figure 2-9). 87Sr/86Sr ratios range from 0.70246 to 0.70258, with an average of
0.70250. These values are well within the range of N-MORB East Pacific Rise lavas
from 9-10° N (Sims et al., 2002; Sims et al., 2003; Goss et al., 2010; Goss et al., in
prep) and similar to N-MORB lavas from 9°N OSC. 143Nd/144Nd ratios are also similar to
East Pacific Rise N-MORB and range from 0.513140 (εNd = 9.8) to 0.513196 (εNd =
10.9). Pb isotopes ratios for the 9°N OSC dacites are indistinguishable from 9°N OSC
basalts and other East Pacific Rise lavas, having averages of 206Pb/204Pb = 18.268,
207Pb/204Pb = 15.473, and 208Pb/204Pb = 37.679. Pb isotopes for 9° N dacites form a
tight cluster in the center of the field defined for other 9° N lavas (Figure 2-9).
Petrogenetic Models For High-Silica Lavas
We now examine the results of various models of fractional crystallization, partial
melting and assimilation and compare the results with the geochemical data described
above to evaluate their relevance to the formation of MOR dacites. Specifically, we
30
focus on physically reasonable models that are consistent with the highly differentiated
major element concentrations, high concentrations of incompatible elements, distinct
trace element patterns, and N-MORB-like isotopic signatures. The models must also be
able to explain relatively high U, Th, Zr and Hf, and low Nb and Ta as well as the flat
REE patterns. Additionally, markedly high Cl, K (and high Cl/K), Al2O3, and low P2O5
must be accounted for in successful petrogenetic schemes.
Crystal Fractionation
Several petrologic models, including MELTS thermodynamic modeling (Ghiorso &
Sack, 1995), Rayleigh crystal fractionation, crystal-melt segregation, and in-situ
crystallization are investigated here to determine if various processes of crystal
fractionation can account for major and trace element compositions of MOR dacites.
Rayleigh fractional crystallization
The program MELTS (Ghiorso & Sack, 1995) provides a useful framework to
evaluate if major element compositions of MOR dacites can be produced by crystal
fractionation. Petrologic modeling was also carried out using the program PETROLOG
(Danyushevsky, 2001) with similar results for high-silica lavas, although the programs
generate somewhat different results for intermediate compositions. Several 9°N N-
MORB’s were used as starting parental melt compositions (Table 2-3) to determine if a
moderately evolved magma could partially crystallize to produce a dacite. These
included a slightly evolved MORB (265-113), a ferrobasalt (265-43), and a FeTi basalt
(264-08). Pressures for each MELTS run were set at 1 kbar to simulate an approximate
minimum depth of crystallization in the shallow oceanic crust, the oxygen fugacity was
set at the quartz-fayalite-magnetite, and the H2O concentrations varied from 0.2 to 0.35
wt% depending on the parent melt composition. Liquid lines of descent were also
31
calculated for higher pressures (up to 5 kbar) to simulate depths of crystallization within
the nascent layer 3 and the shallow mantle. However, the liquid lines of descent
converge on similar end-member compositions at low MgO and high SiO2.
Liquids with MgO and SiO2 compositions similar, though not identical to 9°N OSC
dacites can be produced by ~75-85 % crystal fractionation of a ferrobasaltic parent.
Results predict a crystallization sequence of Ol, followed by Ol + Plag, Ol + Plag + Cpx,
Plag + Cpx + Sp (titanomagnetite), and in some models, late stage crystallization of
apatite. No pigeonite or orthopyroxene crystallization is predicted, in contrast to
experimental results (Juster et al., 1989) but pigeonite is observed in some MOR
andesites and dacites. The calculations suggest temperatures of <980°C are reached
when residual liquids attain compositions similar to dacites. The models are in
agreement with anhydrous experimental results that indicate residual dacitic liquids form
after ~87% at temperatures of ~1050°C (Juster et al., 1989).
For several major elements (TiO2, FeO, and SiO2), compositions similar to 9°N
OSC dacites are obtained through crystal fractionation of a MORB magma; the total
amount of crystallization varies slightly depending on whether the starting composition
was a moderately evolved basalt, ferrobasalt, or FeTi basalt (Figure 2-5; maximum of
~85 % crystallization). In contrast, calculated abundances of K2O, P2O5, Al2O3, and Cl
do not match the dacite end-member composition using any of the parents; modeled
residual liquids have higher P2O5 by factors of 5-10, lower K2O by factors of 1.5-2.5,
lower Al2O3 by factors of 1.4-1.5 and lower Cl by factors of 10 to 12 (Figure 2-5 and 2-
6). Although MELTS does not predict apatite saturation in andesitic liquids the
decreasing P2O5 contents in some andesites and very low values in dacites strongly
32
suggest apatite crystallization. Juster et al., (1989) calculated apatite saturation would
occur at approximately 0.7 wt% P2O5 in Galápagos Spreading Center andesites.
Additional tests of Rayleigh fractionation include modeling the behavior of trace
elements using as input parameters the degree of crystallization and mineral modal
proportions determined from the MELTS modeling. Basaltic partition coefficients are
used in the trace element modeling up to ~ 57% SiO2, and andesitic partition
coefficients are used for > 57% SiO2 (Table 2-4). The Rayleigh fractionation equation
(Cl/Co=F^(D-1)) is used to simulate a continuously evolving magma chamber in which
phenocrysts are immediately separated from the liquid. The starting composition was a
ferrobasalt from the 9°N OSC (265-43).
As shown in Figure 2-7, the trace element concentrations observed in the dacites
cannot be reproduced using Rayleigh fractionation (Figure 2-7) with the constraints
imposed by major element variations. This model reproduces some dacite incompatible
element compositions but does not reproduce the observed enrichments in most of the
incompatible elements. Although the most incompatible elements (Rb, Ba, Th, U) show
the greatest difference between the observed and calculated compositions, even less
incompatible elements (Nb, Zr, Y, Hf) require > 90% crystal fractionation. For instance,
maximum calculated Zr and Nb concentrations are 705 ppm and 13 ppm, respectively,
compared to an average of 870 and 15 ppm in the dacites. In addition, the UN/NbN is
predicted by modeling to be <1, whereas the measured dacite values are >1 and the
modeling does not reproduce the high ZrN/DyN and CeN/YbN and low NbN/LaN ratios in
the 9°N OSC lavas (Figure 2-8). The middle to heavy REE concentrations (i.e., Nd, Sm,
Eu, Dy, Yb, and Lu) can only be generated by > 90% crystal fractionation. Regardless,
33
such extreme degrees of crystallization are inconsistent with major element model
calculations.
Summarily, the calculated liquid lines of descent do not provide a good fit to the
observed major and minor compositions of the high-silica lavas, and trace element
models parameterized from the MELTS calculations do not reproduce measured trace
element abundances or trace element ratios. Thus, we conclude that extensive low-
pressure crystal fractionation is unlikely to be the sole mechanism to explain the
genesis of the high-silica lavas at the 9oN OSC.
Crystal-melt segregation model
Bachmann & Bergantz, (2004) suggest that intermediate liquids (andesites and
dacites) will separate from crystals (via filter pressing) when a magma has undergone
>40-50 volume percent crystallization. The segregated melt, which is more evolved
than the original melt, will once again crystallize until it reaches 40-50% phenocrysts,
when again the new evolved melt will separate from the phenocrysts. While this
segregation model applies strictly to systems of intermediate compositions (Bachmann
& Bergantz, 2004), here we evaluate whether a basaltic magma can evolve
geochemically, through a series of segregation events, to form compositions similar to
the MOR dacites.
To simulate these conditions a 9°N OSC ferrobasalt was allowed to undergo
equilibrium crystallization to andesitic compositions, using MELTS thermodynamic
calculations (Ghiorso & Sack, 1995) and starting conditions described for Rayleigh
fractionation models. At this composition, the liquid separates from the phenocrysts
(Bachmann and Bergantz, 2004), creating a new parent melt. This parent composition
34
becomes the new starting concentration (an andesite) for the next run, which
subsequently crystallizes 50% by volume. This process was repeated (3 times) until
MgO and SiO2 concentrations similar to the 9°N OSC dacites were obtained (Figure 2-
10). This occurred after three segregation events or 87.5 wt% crystallization, however,
as was noted during the MELTS calculations, several dacitic major and minor element
concentrations (i.e. Al2O3, K2O and P2O5) could not be produced. In addition, the
calculated incompatible trace element abundances produced using this process were
also lower than those observed in the MOR dacites (Figure 2-10).
In situ crystallization calculations
A different approach to magma crystallization is in situ crystallization, where
phenocrysts do not separate from interstitial melt until a small remaining volume of
liquid is pressed from the crystallizing mush and mixed with the main body of melt (e.g.
Langmuir, 1989; Reynolds & Langmuir, 1997; Pollock et al., 2005). This process
assumes that crystallization occurs along a temperature gradient within a solidification
front (or boundary layer) and interstitial melt evolves independently from the main melt
body. The mixing of interstitial melt back into the main magma body, gradually causes
bulk increases in incompatible element abundances and changes in trace element
ratios (Langmuir, 1989). This process will cause an increase in highly incompatible
elements compared to Rayleigh crystal fractionation, because these elements are
continually returned to the residual magma body (Langmuir, 1989).
In situ crystallization was evaluated following Reynolds and Langmuir (1997), with
a starting composition of a 9°N OSC ferrobasalt (265-43) using the same partition
coefficients as for Rayleigh fractional crystallization. Crystallizing phases include Ol,
35
Plag, Cpx, and eventually, Fe-oxides. In the modeled system, the boundary layer is
always 5% of the liquid magma chamber volume and the boundary layer crystallizes
until 35% interstitial liquid remains (Reynolds & Langmuir, 1997). All of the residual
liquid mixes back into the magma chamber during each iteration of boundary layer
crystallization. As the magma chamber continues to crystallize by this mechanism, the
boundary layer moves inward leaving restite crystals behind and the volume of the
magma body decreases. Theoretically, this process will continue to modify the liquid
magma chamber composition until an infinitesimally small amount of melt is left.
After 85% in situ crystallization (Figure 2-10) calculated major and incompatible
trace element concentrations do not match those measured in the MOR dacites. This
processes can only account for Zr concentrations in the melt that are less than 3 times
the original concentration, reaching values of 320 ppm. Even > 95% in situ
crystallization cannot reproduce the enriched trace element signatures of the MOR
dacites. Although in situ crystallization does enrich the residual melt in incompatible
elements compared to Rayleigh crystal fractionation, the failure to reproduce the
observed enrichments in MOR dacites suggests that the latter cannot result from this
process either.
Partial Melting (Anatexis)
Experimental studies (Beard & Lofgren, 1991; Koepke et al., 2004) suggest that
<10% partial melting of altered oceanic crust will produce melt with major element
concentrations consistent with MOR dacite compositions. Of particular note are the high
SiO2, Al2O3 and K2O and low FeO, TiO2, and P2O5 concentrations produced by partial
melting of amphibolite facies or greenschist facies minerals because these are also the
chemical characteristics of MOR dacites. Basaltic rocks are known to undergo partial or
36
complete alteration and recrystallization on-axis due to pervasive high temperature
hydrothermal alteration (Alt et al., 1986; Gillis & Roberts, 1999). Such altered rocks are
an attractive starting composition for anatexis because their solidus temperatures are
much lower than fresh MORB.
Altered oceanic crust can have a wide range of trace element concentrations,
depending on the degree of alteration (Alt et al., 1986). To better evaluate the
composition of wall rock involved in the formation of dacites on the 9°N OSC, we use
the batch melting equation (Cl=Co/(Dbulk [1 - F] + F) and andesitic partition coefficients
(Table 2-4) to solve for a range of possible parental wall rock compositions (Co) and
then compare these results to compositions of fresh and altered MORB. Trace element
patterns generated from the calculations suggest altered basalt provides a better fit than
fresh MORB as a source (wall rock) composition for the 9°N OSC dacites (Figure 2-11).
Consequently, we model partial melting of an altered MOR basalt (Nakamura et al.,
2007) to determine if this process could produce the geochemical characteristics
observed in the 9°N OSC lavas (Figure 2-11).
Altered basaltic wall rock is melted using two different modal mineralogies,
including one with amphibole (Haase et al., 2005) and one without (Koepke et al.,
2004). Calculated trace element concentrations resulting from 1 to 15% partial melting
of ocean crust are shown in Figures 2-7, 2-8, and 2-12. Partial melting in the absence
of amphibole (19% Ol, 30% Cpx, 50% Plag, 1% Ilm) can reproduce some, but not all, of
the trace element enrichments observed in the 9°N OSC dacites (Figure 2-12a). In
particular, the HREE in the 9°N OSC dacites are higher than the calculated
abundances. The concentrations derived from partial melting of altered crust with
37
amphibole (20% Cpx, 25% Opx, 49% Plag, 5% Amph, 1% Fe-Oxide; based on modal
proportions from (Haase et al., 2005)) are much closer to the concentrations of
incompatible elements in the 9°N OSC dacites, suggesting that amphibole is an
important component in the melting source rock (Figure 2-12b).
The best-fit model relies on <10% partial melting of amphibole-bearing altered
oceanic crust to produce incompatible trace element compositions comparable to those
in 9°N OSC dacites. In particular, the important characteristics of this model are melts
with positive Zr and Hf anomalies, negative Nb and Ta anomalies on mantle normalized
diagrams (Figure 2-12), relatively high U and Th concentrations, and high UN/NbN and
CeN/YbN ratios (Figure 2-8).
Melting of altered basalt can also produce elevated Cl concentrations similar to
those observed in the MOR dacites (Figure 2-6). Although Cl partition coefficients are
poorly constrained, we use estimates to model Cl partitioning during melting (Gillis et
al., 2003). Using the median Cl concentration from altered basalts in ODP hole 504B
(350 ppm) as a starting composition and modal proportions described above, we
calculate a range of possible Cl enrichment for 1-15% melting to be from 0.2 wt% to >
1.0 wt%. This spans the range of Cl concentrations observed in MOR dacites (Figure 2-
6).
Assimilation Fractional Crystallization
The Energy Constrained – Assimilation Fractional Crystallization (EC-AFC)
formulation of Bohrson & Spera (2001) is used to assess the role that these associated
processes play in the formation of MOR dacites. The amount of crystallization required
to produce enough heat to melt the surrounding crust is calculated and in turn, this
38
produces a specific mass of melt of a specific composition. Several physical parameters
are required as inputs to EC-AFC calculations (Table 2-5). These include the liquidus of
the magma (~1200°C based on results of MELTS modeling of OSC lavas), the
temperature and solidus of the wall rock, and the temperature of equilibrium between
the wall rock and the magma. The initial magma composition is assumed to be N-
MORB and the assimilant is amphibole-bearing altered ocean crust with the modal
composition described above.
The wall rock may span a range of temperatures (800°C to 40°C) depending on
the age of the ocean crust (Maclennan, 2008). Higher initial wall rock temperatures
allow melting to begin earlier in the evolution of the magma reservoir because less
additional heat is required to raise the wall rock above its solidus temperature, i.e.,
smaller amounts of crystallization are required to initiate melting and assimilation.
However, this lowers the overall amount of incompatible trace element enrichment in
the resulting magma because the initial magma is less chemically evolved during
assimilation and larger masses of anatectic melt may be produced (Figure 2-13). For
instance, crust with an initial temperature of 800°C will begin melting after 50 to 55%
crystallization and the resultant magma has a maximum of ~25 ppm La (Figure 2-13).
Antithetically, lowering the wall rock temperature decreases the total mass of wall rock
assimilated while increasing the amount of crystallization needed to initiate melting.
Consequently, this causes an increase in the overall incompatible element
concentration possible in the melt. Thus, crust with an initial temperature of 50°C
requires >85% crystallization to begin melting, but results in concentrations of ~35 ppm
39
La in the magma (Figure 2-13). Based on these competing processes, the best-fit wall
rock temperature to generate the 9°N OSC dacites is between 650 and 720°C.
The local solidus, as described by Bohrson and Spera (2001), is the solidus of the
assimilant, in this case, amphibolitized basalt. Several experimental studies have
examined dehydration melting of altered oceanic crust and amphibolites (Hacker, 1990;
Rapp et al., 1991; Wolf & Wyllie, 1994; Johannes & Koepke, 2001) but few studies were
performed under conditions comparable to those expected at MOR (Beard & Lofgren,
1991). These experiments determined that the solidus temperatures of altered basalts
are between 850oC and 900°C (Beard & Lofgren, 1991). Gillis & Coogan (2002)
discuss the effects of melting altered crust at the roof of an axial magma chamber and
suggest a solidus temperature of 875°C. Ti-in-zircon thermometry (905±34°C at TiO2
activity of 0.32±0.02 estimated from coexisting Fe-Ti oxides) from phenocrysts in the
Juan de Fuca Ridge dacites is broadly consistent with zircon saturation thermometry
(average of 824± 15°C)) and Fe-Ti oxide temperatures (~830°C Schmitt et al., in prep).
Based on these combined results, 875°C was used as the local solidus temperature for
AFC calculations.
The final thermal input parameter is the temperature of equilibrium, which is
defined as the final equilibrium temperature of the magma and wall rock. Generally, this
temperature should correspond to the temperature of the erupted lava. Although the
temperature of the erupted 9°N OSC dacites is uncertain, the temperatures of other
dacitic magmas have been estimated. The temperature of crystallization of Galapagos
Spreading Center andesitic magma was calculate to be as low as ~910 to 940°C based
on coexisting titanomagnetite and ilmenite grains (Perfit et al., 1983) and experimental
40
partial melts resulting from melting of oceanic gabbros showed dacitic melts at
temperatures of approximately 900°C (Koepke et al., 2004). Consequently, we use
900°C as the input equilibrium temperature. The partition coefficients are the same as
those used for partial melting and fractional crystallization models. Basaltic partition
coefficients are used for the fractionating magma, while andesitic bulk Kd values were
used for the assimilant in the absence of a comprehensive dataset of dacite Kd values.
Results of EC-AFC calculations suggest that combinations of 73 to 85% crystal
fractionation of a basaltic magma and assimilation of 5 to 20% by mass of partially
melted wall rock produces melts that have trace element compositions consistent with
the 9°N OSC dacites. In the best-fit model, melting and assimilation begins after 68%
crystallization and a further 5-17% crystallization occurs as the wall rock melt is
assimilated.
Consequently, EC-AFC trace element calculations suggest that many of the
incompatible trace element concentrations and ratios observed in the 9°N OSC dacites
can be explained through this combination of processes (Figures 2-7, 2- 8, 2-14). Of
particular importance are negative Nb and Ta anomalies (relative to La), an increase in
Zr and Hf concentrations (relative to HREE), relatively flat mantle-normalized HREE
patterns, and ratios of light to heavy REE and middle to heavy REE that are similar to
those observed in MOR dacites (Figure 2-14). For instance, Zr concentrations in the
AFC models are 852 ppm and Nb concentrations are 16 ppm compared to an average
9°N OSC dacite concentrations of 870 and 15 respectively. Although the overall fit of
the model data to the observed data is encouraging, the model values for Ba, Th, U and
Hf are slightly under-enriched (Figure 2-14). However, it should be pointed out that
41
some of the input parameters to these calculations, such as the actual degree of
alteration (and hence its composition) of the crustal assimilant, and the temperature of
the surrounding wall rock, are not well constrained and are very likely not constants.
Discussion
Petrogenesis of High Silica Lavas
Extreme crystal fractionation, partial melting of crustal material, and/or AFC
processes have been proposed as explanations for the formation of highly silicic
compositions in continental interiors, arc and ocean island settings, but only a few
studies have focused on the formation of high-silica lavas at MOR (Byerly & Melson,
1976; Perfit et al., 1983; Juster et al., 1989; Haase et al., 2005). The petrogenetic
calculations presented above demonstrate that crystal fractionation alone is not a viable
mechanism for the formation of high-silica MOR lavas, despite using a range of starting
compositions and several different end-member models (Figures 2-5, 2-7, 2-8). Instead,
results emphasize the importance of partial melting and assimilation of altered material
in the formation dacites on MOR.
Geochemical evidence of partial melting
Will crustal anatexis create geochemical signatures similar to those observed in
MOR dacites? Based on elemental systematics (e.g. Cl, U/Nb; Figures 2-6 and 2-7)
and partial melting calculations presented, it appears that partial melts of altered
oceanic crust may be involved in the generation of the MOR dacites. Partial melts of
hydrothermally altered crust produces distinct signatures compared to those of
unaltered oceanic crust, due to changes in mineralogy and bulk composition during
hydrothermal circulation. Hydrothermal circulation in layer 2B or the top of layer 3 may
cause alteration to greenschist or amphibolite assemblages, where Ca-rich plagioclase
42
is replaced by sodic plagioclase and pyroxenes form rims or overgrowths of amphibole
(Alt et al., 1986; Coogan et al., 2003b). The degree to which this occurs depends on
temperature, water/rock ratios and fluid chemistry. To explain the geochemical
signatures in the MOR dacites, our melting assemblage must include amphibole.
Melting of amphibole-bearing assemblages, a common component in altered layer
2B (Alt et al., 1986; Coogan, 2003; Coogan et al., 2003b), can explain the anomalously
high Al2O3 concentrations observed in the oceanic dacites (Figure 2-5). Dehydration
partial melting experiments, where water exists only as hydrous phases within the rock,
provide a better fit to MOR dacite compositions than hydrous partial melting results
(Koepke et al., 2004). Dehydration melting experiments produce a range of Al2O3
concentrations (Beard & Lofgren, 1991) that are similar to or higher than MOR dacites
(Figure 2-5). Comparatively high Na2O concentrations in the dacites may result from
the melting of albitic plagioclase. Elevated Na2O concentrations are not observed in all
experimental results of Beard and Lofgren (1991) but appear to be a function of the
degree of albitization of the starting material. Variable P2O5 concentrations are also
observed in the experimental melts suggesting that P2O5 contents are very low in the
source rock or that it is a residual phase in the melting residue. Similar conclusions can
be applied to the MOR dacites, which have low phosphorous contents (Figure 2-5);
however, this may also be a function of apatite crystallization during AFC processes
(see next section). Additionally, low FeO and TiO2 suggest that Fe-oxides are not a
primary melting component in the source rock. This is consistent with our proposed
source rock comprised of olivine, plagioclase, cpx, amphibole, ± Fe-oxides.
43
High Cl concentrations in MOR dacites also support the role of partial melting of
rocks altered by seawater-derived fluids (Michael & Schilling, 1989; Michael & Cornell,
1998; Coogan et al., 2003b; Gillis et al., 2003). Although Cl behaves incompatibly
during crystallization many MOR lavas show over-enrichments compared to other
elements with similar compatibilities. For example, after ~85 % fractional crystallization
of a MORB parent (with 0.01 wt% Cl), there is less than a ten-fold enrichment in Cl,
resulting in concentrations of ~0.07 wt%, compared to an order of magnitude more
(~0.7 wt% Cl) in the dacites. Analyses of altered basalt from sheeted dikes in drill holes
show that Cl concentrations span a range from 49-650 ppm (Sparks, 1995). Partial
melting (1-15%) of an amphibole-bearing wall rock (with 350 ppm Cl) results in anatectic
melts with 0.9 to 0.3 wt% Cl – covering the range observed in dacites (Figure 2-6).
Hydrothermal alteration and metamorphism are known to cause increased
concentration of some trace elements, including U, Th, Rb, and Ba, as well as Cl (e.g.
Alt & Teagle, 2003). Observed positive anomalies of some highly incompatible
elements (e.g. U and Th) relative to other incompatible elements with similar distribution
coefficients are consistent with partial melting of hydrothermally altered ocean crust
(Figure 2-12). Partial melting may also explain some of the anomalies in the high field
strength elements whose concentrations are not affected by alteration or metamorphism
but can be fractionated due to mineralogic effects. For example, the relatively low
abundances of Nb and Ta (moderately compatible during melting) and the relatively
high Zr (highly incompatible) concentrations in the high-silica lavas are a consequence
of partial melting of altered crustal material (Figure 2-12). Additionally, melting and AFC
44
models above point to the importance of amphibole in the melting assemblage to
explain the HREE (compare Figure 2-12a and 2-12b).
The need for crystallization, assimilation and altered crust in dacite petrogenesis
We propose that the partial melting and assimilation of oceanic crust plays a
significant role in the formation high-silica MOR lavas; however, the we stress that most
of the heat required to melt the wall rock comes from extensive fractional crystallization.
Coogan et al., (2003b) show that the latent heat of crystallization from the formation of 4
km thick gabbro sequence provides enough energy to heat ~1.3 km of overlying crust
from 450 to 1150°C, which promotes partial melting. It is the assimilation of these partial
melts into a fractionally crystallizing magma reservoir that produces the highly evolved
melts with enriched incompatible trace element signatures (e.g. De Paolo, 1981; Bédard
et al., 2000).
Major element compositions of MOR dacites often lie between experimental partial
melts of altered basalt (Beard & Lofgren, 1991) and liquids produced by moderate to
large extents of crystal fractionation (Figure 2-5). The major element compositions of
magmas produced by AFC may therefore, lie between partial melts of altered oceanic
crust and fractionated basaltic magmas. This is particularly apparent in Al2O3 and may
explain why very few lavas at the 9°N OSC (including ferrobasalts and basaltic
andesites) lie on the calculated liquid lines of descent (Figure 2-5).
Results from EC-AFC calculations (Bohrson & Spera, 2001) confirm that
assimilation of anatectic melts into a residual fractionated magma can explain a wide
range of trace element concentrations in MOR dacites (Figure 2-7, 2-8, 2-14). The
best-fit model for 9°N OSC dacite compositions requires significant crystal fractionation
45
(73-85 wt%) of a ferrobasalt parental magma in combination with 10 to 25% (by mass)
anatectic melt, which provides the additional incompatible element enrichments
observed.
Our models also indicate that in order to explain the relative enrichments in Rb,
Ba, Th and U concentrations present in the MOR dacites assimilation of low degree
partial melts of hydrothermally altered oceanic basalt are required. Relatively low Nb
and Ta concentrations are a consequence of both the removal of phenocrysts during
late-stage crystal fractionation and residual iron-titanium oxides in the partially melted
wall rock. Elevated Zr and Hf concentrations result from little to no zircon crystallization
in the fractionating magma and/or no residual zircon in the melting assemblage. It is
important to note that the extreme Cl enrichments in MOR dacites require a seawater
component that can be derived by AFC process and that Cl over-enrichments observed
in many MORB have been explained either by small amounts of assimilation of either
hydrothermally altered ocean crust or Cl-rich brines stored in the crust (Michael &
Schilling, 1989; Michael & Cornell, 1998; Perfit et al., 1999; Coogan et al., 2003a; le
Roux et al., 2006).
Isotopic signature of assimilation
Given that the 9°N OSC dacites have radiogenic isotopes similar to those in
spatially related basalts (Figure 2-9) what effects might assimilation, particularly of
altered crust, have on derivative melts? AFC processes can change radiogenic isotopic
ratios if the assimilant has relatively high concentrations of the element in question and
significantly different isotopic ratios than the original magma reservoir (e.g. Taylor,
1980; De Paolo, 1981). In general however, the effect on the resultant isotopes is less
dramatic in AFC processes compared to partial melting because AFC processes create
46
mixtures of altered and fresh material, while melting alone will retain the isotopic
signature of the altered crust.
Assimilation of altered oceanic crust may increase Sr isotope ratios depending on
the amount of assimilation and extents of fluid/rock interaction (e.g. Alt & Teagle, 2003).
Altered ocean crust can have a range of Sr concentrations (from less than to greater
than typical MORB compositions) depending on the type/degree of alteration (Alt &
Teagle, 2003). Assuming 80% fractional crystallization (which will not change the
isotopic ratios) and assimilation of 10 mass percent wall rock, we can calculate the
isotopic composition of the resulting melt using a ratio of 2:1 (fractionated melt to
anatectic melt). Using reasonable values for the isotopic ratios of altered sheeted dike
lavas (0.7028; average of basal dikes from Pito Deep; Hess Deep and Hole 504B;
Barker et al., 2008) and initial MORB parental magma (0.7025) and their respective Sr
concentrations (100 ppm and 120 ppm; a typical altered East Pacific Rise
concentration) mass balance calculations indicate assimilation of altered crust will
cause an increase of ~0.0001 in the 87Sr/86Sr ratio of the final magma. EC-AFC
modeling of Sr isotopes produces similar results but requires less crystal fractionation
(73%) and a ratio of crystallization to assimilation of 0.07 to produce the most
radiogenic 87Sr/86Sr signatures observed in the MOR dacites (0.70258). Therefore, this
process has the potential to slightly affect the Sr isotope ratios in the dacitic magma, but
will not result in ratios as elevated those generated directly from partial melting of
altered oceanic crust (~0.7028). Additionally, slightly elevated Sr isotope ratios in high-
silica lavas from the Galapagos Spreading Center are consistent with AFC processes
(Perfit et al., 1999). In comparison, Nd isotopes are unaffected during fractional
47
crystallization and are relatively immobile during hydrothermal alteration (Michard &
Albarède, 1986; Delacour et al., 2008). Nd isotopes from the OSC dacites are similar to
basalts from the region.
9° N OSC dacites form a tight cluster in Pb isotopic composition compared to 9°N
OSC basalts (Figure 2-9). Pb isotopes are not significantly affected by hydrothermal
alteration provided sediments (which are not abundant in this environment) are not
involved in the alteration process. We suggest that the similarity in isotopic ratios in the
dacites compared to the basalts represents an overall averaging of isotopic values from
basalts in the region due to melting and assimilating a range of different MORB
compositions at the base of the sheeted dike layer.
Relatively low oxygen isotope ratios observed in MOR dacites from the Galapagos
Spreading Center (Perfit et al., 1999) and the 9° N OSC (Wanless et al., 2009; Wanless
et al., in prep-a) also support these conclusions. Fresh MORB will have mantle oxygen
isotope values (~5.5), however, seawater alteration (seawater δ18O = 0) will decrease
this ratio (Gillis et al., 2001), while fractional crystallization of Fe-oxides, and to a lesser
extent olivine and pyroxene, will cause an increase (Matsuhisa et al., 1973). Therefore,
partial melting and assimilation of altered basalt should produce melts with lower
oxygen isotope ratios than predicted by fractional crystallization calculations
(Muehlenbach & Clayton, 1972). MOR dacites have oxygen isotope ratios similar to
MORB values (Perfit et al., 1999; Perfit et al., 2007; Wanless et al., 2009; Wanless et
al., in prep-a), suggesting that fractional crystallization alone cannot explain the
formation of dacites on MOR. Taylor (1968) suggests that to a first approximation, the
effect of assimilation on oxygen isotopes can be determined using mass balance.
48
Assuming an evolved magma has a δ18O ratio of 6.8 (largely due to fractionation of
silicates and iron oxides) and an assimilant has a δ18O ratio of 3.5 (due to seawater
alteration), the resultant oxygen isotope ratio of the AFC magma would be ~6. This
value is similar to those observed in the MOR dacites and is less than predicted by
fractional crystallization alone.
AFC Processes and Tectonic Setting
The remarkable geochemical similarity of dacites erupted at the three different
MOR discussed here indicates similar processes are controlling their petrogenesis
(Figure 2-2) and we propose these processes are linked to tectono-magmatic settings.
Andesites and dacites have erupted on several ridges, often in regions of propagation,
such as propagating ridge tips and OSC (Christie & Sinton, 1981; Perfit et al., 1983;
Sinton et al., 1983; Sinton et al., 1991). High-silica lavas have also been found along
the Pacific-Antarctic Rise (Haase et al., 2005), at Axial Seamount on the Juan de Fuca
Ridge (Chadwick et al., 2005) adjacent to a large axial magma chamber, and at the end
of a first and second-order ridge segments on the Northern East Pacific Rise (~ 8°37’ N
and 10°30’N (Langmuir et al., 1986). Collectively, these lavas erupted on intermediate
to fast spreading ridges, in settings where magma reservoirs have the potential to
undergo extensive fractional crystallization and interact with colder, and variably altered
crust.
A key result of this study is that the geochemical signatures of MOR dacites
require assimilation of partial melts of hydrothermally altered crust into an extremely
fractionated magma (that provides the heat needed to melt the oceanic crust). The
extensive amounts of fractional crystallization required suggest episodic or sporadic
49
magma supply to magma reservoirs, which may not be characteristic of more “steady-
state” ridge environments. These requirements are met at the ends of ridge segments,
where magma reservoirs may have a sporadic magma supply. In these regions,
magmas are considered to be fed intermittently to the ridge tip through dike propagation
from a more robust central region (Christie & Sinton, 1981). Between diking events the
ridge tip magma supply is cut off, allowing for increased extents of crystal fractionation
and interaction of the melt with older, altered crust. This may increase the likelihood of
eruption of high-silica lavas through AFC processes. This is not to suggest that AFC
processes do not occur in “steady-state” ridge environments but that the high-silica
melts may not be preserved or erupted in these regions (see section on Effects of
Assimilation on Typical MORB).
Model for Formation of MOR Dacites
Based on the similarity of composition of high-silica lavas from three MOR,
petrologic modeling calculations applied to dacites at the 9oN OSC, and published
experimental results, we suggest that MOR dacites form under specific conditions that
include: 1. A tectono-magmatic setting in which magma injection is episodic, allowing for
extensive crystal fractionation; 2. The presence of altered crust, which facilitates
geochemical enrichments observed in the MOR dacites. Based on these two
requirements, the tectonic setting and available geophysical information, we propose
the following model for dacite formation on MOR (Figure 2-15):
1. Injection of basaltic magma by lateral dike propagation. Formation of axial magma reservoirs.
2. Magma supply to the region is cut off or reduced, allowing for extensive fractional crystallization of the magma reservoir.
50
3. During extensive fractional crystallization the released latent heat of crystallization initially heats then partially melts the surrounding altered wall rock, which might be layer 2b dikes or high level altered gabbros.
4. The anatectic melts are assimilated into the fractionating magma body. The AFC melts may be of dacitic composition depending on crustal temperatures, extent of fractional crystallization, amount of anatectic melt, and efficiency of assimilation.
This model may account for the formation of highly evolved magmas at OSC,
propagating ridge tips, ridge-transform intersections and along dikes associated with the
down-rift volcanism on Axial seamount. This situation may also be analogous to Krafla
volcano in Iceland, where the imaged melt lens is thought to be primarily composed of
iron-rich basalts but high-silica lavas are associated with the edges of the caldera rim,
where increased magma-rock interactions may be likely (Nicholson et al., 1991).
Relationship of Melt Lens to Dacites at 9°N
Dacitic lavas at 9°N OSC erupted on-axis, over the eastern edge of the large,
seismically imaged melt lens (Kent et al., 2000; Dunn et al., 2001). Despite the eruption
of young, fresh high-silica lavas in the neo-volcanic zone, the underlying melt lens is not
assumed to be dacitic in composition. The composition of the basalts overlying the axial
magma chamber suggests it has undergone a moderate amount of crystallization (to
ferrobasalts). This suggests that the melt lens is composed primarily of basaltic magma
that has mixed to varying degrees with a highly evolved end-member on axis (Wanless
et al., in prep-b). There is also evidence of relatively young off-axis basaltic volcanism
over the main body of the imaged melt lens (north of the overlap basin; (Nunnery et al.,
2008).
The presence of a large, seismically imaged melt lens at 9°N OSC does not
contradict the episodic magma supply requirement for dacite formation. Instead, it may
enable and enhance AFC processes in the region. AFC modeling suggests that
51
extensive fractional crystallization is required to produce dacitic compositions, which
suggests low magma supply. Somewhat antithetically, the 9°N region has an
anomalously large basaltic axial magma chamber suggesting the current melt lens may
only be indirectly related to the dacites. We envision that the large melt lens is acting as
a mobilizer for the eruption of the dacites that formed in isolated magma pockets or sills
on the eastern edge of the axial magma chamber. Additionally, small-scale local mixing
of highly evolved compositions with the moderately evolved basaltic melt lens may
account for the range of compositions erupting at the OSC (Figure 2-15).
Effects of Assimilation on Typical MORB Compositions
The formation of dacite compositions on MOR requires assimilation of anatectic
melts into residual fractionated magmas, however, AFC processes may also explain
slightly elevated incompatible elements observed in MORB lavas from all sections of
MOR. The geochemical signatures of anatectic melts may be subtle in less evolved
magmas, however, elevated Cl, Al2O3, and K2O are common (Michael & Schilling, 1989;
Michael & Cornell, 1998; Perfit et al., 1999; Coogan et al., 2003a; le Roux et al., 2006).
At more magmatically active ridge sections, wall rock may have a higher initial
temperature, which would allow for melting and assimilation to begin earlier in the
evolution of the magma body and require less fractional crystallization to occur prior to
melting (Figure 2-13). For instance, at 800°C assimilation begins after only 53%
fractional crystallization compared to 68% crystallization for 720°C crust. Assimilation of
anatectic melts into a magma reservoir that has undergone less crystallization produces
less evolved compositions and therefore, cannot produce MOR dacites. It does,
however, increase the incompatible element abundances in the melt phase and may
52
explain the commonly noted anomalous incompatible element enrichments, low FeO
and elevated Al2O3 concentrations in some “normal” MORB lavas.
Conclusions
The majority of eruptions at spreading centers produce basalts with relatively
limited chemical variability; however, high-silica lavas have been sampled at several
ridges. Eruptions of andesites and dacites are typically associated with ridge
discontinuities and produce significant volumes of lava at a local scale. Limited
amounts of these lavas have been sampled at the southern terminus of the Juan de
Fuca Ridge, along the eastern Galápagos spreading center, at 8°37’N and off-axis at
~10°30’N on the East Pacific Rise. We have documented more voluminous eruptions of
high –silica lavas including highly evolved dacite on the propagating eastern limb of the
9°N overlapping spreading center (OSC) on the East Pacific Rise. Collectively, the
dacites appear to represent an end-member composition that shows similar major
element trends and incompatible trace element enrichments, suggesting similar
processes controlled their petrogenesis.
The formation of highly evolved lavas on MOR requires a combination of partial
melting, assimilation and crystal fractionation. The highly enriched incompatible trace
element signatures cannot be produced through crystal fractionation alone and appears
to require partial melting of altered ocean crust. EC-AFC modeling suggests significant
amounts (>75%) crystallization of a MORB parent magma and modest amounts (5-
20%) of assimilation of hydrothermally altered ocean crust can produce geochemical
signatures consistent with dacite compositions. The AFC process explains trace
element abundances in high-silica lavas and accounts for several major and minor
element concentrations (i.e. Al2O3, K2O and Cl).
53
An important constraint provided by AFC calculations is the temperature of the
assimilant. Varying the wall rock temperature can change the amount of
crystallization/assimilation that occurs and the overall enrichments observed in
incompatible trace element concentrations. The formation of dacites at the 9°N OSC
requires temperatures of surrounding crust to be 650 - 720 °C, which requires >68%
fractional crystallization ferrobasalt magma before melting can begin. While this amount
of crystallization is unlikely in regions of high/constant magma supply, the surrounding
wall rock in typical ridge settings may be much warmer than at the ends of ridge
segments, allowing for assimilation at much lower percents of crystallization. At wall
rock temperatures of 800 °C, calculations suggest that assimilation begins after ~53
wt% fractional crystallization. This suggests that while conditions are not appropriate for
the petrogenesis of dacites at typical ridge settings, assimilation of crustal material may
be common but geochemically cryptic.
The formation of high-silica lavas on MOR appears to require a unique tectono-
magmatic setting, where episodic magma supply allows for extensive crystal
fractionation, partial melting and assimilation. These conditions are met in regions of
ridge propagation, such as OSC and propagating ridge tips, where diking allows for
episodic injection of magma into older, altered ocean crust. Here, the magma
undergoes extreme crystallization without repeated replenishment, creating enough
latent heat of crystallization to melt and assimilate surrounding wall rock.
Acknowledgements
We thank the Captain, officers and crew of the R/V Atlantis for all their help
during cruise AT15-17, the MEDUSA2007 Science party (including White, S., Von
Damm, K., Fornari, D., Soule, A., Carmichael, S., Sims, K., Zaino, A., Fundis, A.,
54
Mason, J., O’Brien, J., Waters, C., Mansfield, F., Neely, K., Laliberte, J., Goehring, E.,
and Preston, L.) for their diligence in collecting data and samples for this study. We
thank the Jason II shipboard and shore-based operations group for their assistance in
collecting these data and HMRG for processing all DSL-120 sidescan and bathymetry
collected during this cruise. Discussions with S. White and A. Goss are gratefully
acknowledged and contributed to this research. Thanks to G. Kamenov and the UF
Center for Isotope Geoscience for laboratory assistance. This research was supported
by the National Science Foundation [grants OCE-0527075 to MRP and OCE-0526120
to EMK].
55
Table 2-1. Dacite major and trace element data Sample #
266-58
265-65
265-64
265-67
266-50
266-53
265-84
265-63
266-47
265-85
266-46
265-94
SiO2 63.0 63.8 64.0 64.1 64.3 64.3 64.4 64.4 64.5 65.0 65.0 65.2 TiO2 1.10 1.26 1.28 1.34 1.07 1.06 1.13 1.29 0.99 1.06 0.94 0.97 Al2O3 13.1 13.2 13.1 13.3 13.2 13.3 13.2 13.3 13.2 13.1 12.9 13.0 FeO 8.43 8.14 8.27 8.49 8.08 8.06 8.18 8.22 7.74 7.99 7.17 7.90 MnO 0.16 0.15 0.16 0.16 0.14 0.14 0.15 0.15 0.14 0.16 0.15 0.14 MgO 1.75 1.34 1.60 1.49 1.27 1.12 1.23 1.29 1.02 1.18 1.41 1.13 CaO 4.34 4.21 4.45 4.41 3.78 3.73 3.92 4.21 3.53 3.78 3.71 3.54 Na2O 3.63 3.84 3.46 3.93 4.23 4.16 3.41 3.71 4.94 3.67 4.76 4.29 K2O 0.96 0.97 0.97 0.95 1.10 1.09 1.19 0.99 1.22 1.22 1.19 1.14 P2O5 0.26 0.22 0.20 0.23 0.24 0.25 0.22 0.21 0.23 0.20 0.17 0.23 Cl 0.24 0.65 0.64 Total 96.69 97.17 97.55 98.42 97.35 97.20 96.98 97.76 97.54 97.41 97.43 97.61 Trace Elements (ppm) Li 32 34 32 30 29 30 31 31 Sc 17 20 18 15 15 17 14 13 V 122 140 121 61 102 121 73 63 Cr 13 12 12 4 2 10 4.6 1 Co 15 17 16 13 14 15 12 11 Ni 9.0 9.8 9.2 5.8 7.0 8.4 6.6 4.9 Cu 17 19 18 18 21 17 19 17 Zn 110 124 122 109 100 113 108 105 Ga 30 35 28 29 29 31 28 30 Cs 0.11 0.12 0.12 0.14 0.13 0.10 0.15 0.13 Rb 9.1 10 10 13 13 9.5 14 12 Ba 51 57 53 65 60 50 68 60 Th 1.7 1.8 1.6 2.3 2.3 1.7 2.4 2.3 U 0.59 0.65 0.64 0.86 0.82 0.59 0.92 0.84 Nb 13 15 13 16 15 13 15 16 Ta 0.81 1.1 0.92 1.0 1.4 0.99 1.1 1.2 La 23 26 23 29 27 24 29 29 Ce 67 77 68 82 78 68 82 84 Pr 10 11 10 12 11 10 12 12 Sr 76 90 89 86 70 76 81 68 Nd 46 53 46 54 51 47 52 55 Zr 735 842 622 856 968 745 872 1050 Hf 19 21 18 22 23 19 23 25 Sm 14 15 15 16 14 14 16 16 Eu 3.0 3.3 3.1 3.2 2.8 3.0 3.0 3.0 Tb 3.2 3.4 3.4 3.6 3.1 3.1 3.6 3.4 Dy 21 23 23 24 20 21 24 22 Y 132 148 142 157 133 132 154 146 Ho 4.6 4.9 4.9 5.2 4.3 4.4 5.1 4.8 Er 14 15 15 16 13 13 16 14 Yb 14 15 16 17 14 14 17 15 Lu 2.1 2.3 2.4 2.6 2.1 2.1 2.6 2.3 Tm 2.2 2.3 2.4 2.5 2.1 2.1 2.5 2.3 Gd 17 18 17 19 17 16 19 18 Pb 3.3 5.0 3.4 3.2 4.9 5.8 3.6 5.7
56
Table 2-1. Continued Sample #
264-09
265-70
265-42
265-83
265-95
265-40
SiO2 65.8 66.3 66.5 67.5 67.5 TiO2 0.89 0.87 0.94 0.76 0.77 Al2O3 13.2 13.2 13.0 13.3 13.1 FeO 7.03 7.17 7.92 6.68 6.47 MnO 0.13 0.14 0.16 0.13 0.12 MgO 1.06 0.80 0.89 0.67 0.94 CaO 3.48 3.23 3.50 2.98 3.01 Na2O 4.24 4.08 3.99 3.88 4.43 K2O 1.21 1.33 1.20 1.37 1.21 P2O5 0.20 0.19 0.21 0.16 0.15 Cl 0.58 0.70 0.51 0.67 Total 97.78 97.27 98.91 97.37 97.67 Trace Elements (ppm) Li 27 34 32 32 31 31 Sc 12 12 14 11 10 12 V 46 45 58 32 52 51 Cr 4 4 3 3 3 4 Co 10 10 11 8 8 10 Ni 5.7 5.0 5.9 4.7 5.4 5.3 Cu 16 16 15 14 15 17 Zn 89 106 119 103 98 103 Ga 28 29 30 29 30 30 Cs 0.13 0.17 0.16 0.17 0.13 0.13 Rb 13 15 14 15 12 12 Ba 66 73 70 76 62 60 Th 2.6 2.7 2.4 2.8 2.4 2.5 U 0.98 1.04 0.91 1.05 0.86 0.84 Nb 15 16 16 16 17 17 Ta 1.0 1.1 1.1 1.1 1.2 1.0 La 29 31 29 31 29 29 Ce 84 88 83 87 84 85 Pr 12 12 12 12 12 12 Sr 78 78 83 76 61 73 Nd 53 55 53 54 55 57 Zr 824 934 816 922 985 1013 Hf 23 25 22 25 25 25 Sm 16 17 17 16 16 17 Eu 3.0 3.1 3.4 3.1 2.9 3.3 Tb 3.6 3.7 3.8 3.6 3.3 3.7 Dy 24 25 25 25 22 24 Y 151 160 160 159 145 146 Ho 5.1 5.4 5.5 5.3 4.7 5.2 Er 16 16 16 16 15 15 Yb 17 18 18 18 15 16 Lu 2.6 2.7 2.7 2.7 2.3 2.4 Tm 2.5 2.6 2.6 2.6 2.3 2.5 Gd 19 19 20 19 18 20 Pb 2.8 3.8 3.5 3.8 4.1 3.6
57
Table 2-2. Radiogenic isotopes for 9°N OSC lavas
Sample 208Pb/ 204Pb 2σ error
207Pb/ 204Pb 2σ error
206Pb/ 204Pb 2σ error
208Pb/ 206Pb 2σ error
E. Limb Basalts 264-04 37.699 1.74E-03 15.476 7.89E-04 18.279 8.93E-04 2.062 2.87E-05 265-18 37.677 1.94E-03 15.472 7.33E-04 18.275 7.80E-04 2.062 3.56E-05 265-35 37.664 1.60E-03 15.470 5.59E-04 18.250 7.06E-04 2.064 3.08E-05 265-43 37.661 1.51E-03 15.467 5.45E-04 18.249 6.11E-04 2.064 3.17E-05 265-113 37.695 1.67E-03 15.477 6.33E-04 18.275 6.83E-04 2.063 2.45E-05 266-01 37.642 1.95E-03 15.469 7.30E-04 18.235 9.22E-04 2.064 2.17E-05 266-33 37.683 2.19E-03 15.474 8.55E-04 18.277 8.56E-04 2.062 2.76E-05 265-05 37.687 1.96E-03 15.478 6.26E-04 18.294 7.24E-04 2.060 3.51E-05 E. Limb Basaltic Andesites 265-24 37.683 1.68E-03 15.473 5.87E-04 18.264 6.10E-04 2.063 3.34E-05 265-56 37.681 1.49E-03 15.474 5.67E-04 18.270 6.18E-04 2.062 3.06E-05 265-91 37.683 1.40E-03 15.475 5.25E-04 18.266 5.05E-04 2.063 3.09E-05 265-103 37.674 2.06E-03 15.472 7.56E-04 18.261 8.79E-04 2.063 2.74E-05 265-109 37.673 1.79E-03 15.471 6.69E-04 18.259 7.04E-04 2.063 2.64E-05 265-125 37.672 1.90E-03 15.471 7.32E-04 18.265 7.80E-04 2.063 2.57E-05 264-20 37.690 1.54E-03 15.476 6.14E-04 18.269 5.94E-04 2.063 3.47E-05 265-49 37.689 1.60E-03 15.474 5.77E-04 18.268 6.22E-04 2.063 3.25E-05 E. Limb Andesites 264-14 37.673 2.08E-03 15.472 9.65E-04 18.262 1.11E-03 2.063 5.32E-05 265-25 37.661 1.69E-03 15.469 6.99E-04 18.251 7.47E-04 2.064 2.44E-05 265-90 37.675 1.52E-03 15.472 6.01E-04 18.262 6.60E-04 2.063 2.75E-05 265-100 37.680 1.46E-03 15.474 5.81E-04 18.265 6.24E-04 2.063 2.68E-05 266-54 37.688 1.69E-03 15.475 6.59E-04 18.272 7.73E-04 2.063 3.29E-05 E. Limb Dacites 264-09 37.676 1.78E-03 15.472 7.14E-04 18.268 8.10E-04 2.062 2.88E-05 265-40 37.674 2.05E-03 15.471 8.24E-04 18.265 8.86E-04 2.063 3.58E-05 265-42 37.678 1.72E-03 15.472 6.53E-04 18.270 7.40E-04 2.062 3.40E-05 265-64 37.682 2.51E-03 15.476 9.76E-04 18.267 1.13E-03 2.063 3.32E-05 265-70 37.676 1.96E-03 15.471 7.58E-04 18.266 7.97E-04 2.063 3.64E-05 265-83 37.681 1.51E-03 15.473 5.94E-04 18.271 6.65E-04 2.062 2.49E-05 265-84 37.689 1.67E-03 15.476 6.55E-04 18.274 7.49E-04 2.062 2.59E-05 265-85 37.677 2.05E-03 15.471 7.79E-04 18.267 8.31E-04 2.063 3.02E-05 265-95 37.680 1.67E-03 15.475 6.95E-04 18.263 8.27E-04 2.063 3.80E-05 266-53 37.678 1.32E-03 15.472 5.01E-04 18.268 5.10E-04 2.063 3.16E-05 2 sigma error reflects in-run machine error. Long term reproducibility estimates are: 87Sr/86Sr = ± 0.00003, 143Nd/144Nd = ± 0.000018, 206Pb/204Pb = ± 0.0034 (205 ppm), 207Pb/204Pb = 0.0028 (184 ppm), 208Pb/204Pb = ± 0.0086 (234 ppm) *Unknowns were normalized to an 143Nd/144Nd value of 0.511215 ± 0.000007 for JNdi-1, which is reported by Tanaka et al. (2000) relative to a La Jolla 143Nd/144Nd value of 0.511858 (Lugmair and Carlson, 1978).
58
Table 2-2. Continued
Sample 87Sr/86Sr 2σ error 143Nd/144Nd* 2σ error Eps Nd E. Limb Basalts 264-04 0.70250 0.000011 0.513163 0.000011 10.2 265-18 0.70249 0.000018 0.513190 0.000013 10.8 265-35 0.70247 0.000012 0.513164 0.000005 10.3 265-43 0.70250 0.000021 0.513154 0.000009 10.1 265-113 0.70249 0.000016 0.513172 0.000005 10.4 266-01 0.70246 0.000015 0.513158 0.000007 10.1 266-33 0.70244 0.000013 0.513160 0.000005 10.2 265-05 0.70243 0.513191 0.000008 10.8 E. Limb Basaltic Andesites 265-24 0.70256 0.000013 0.513187 0.000005 10.7 265-56 0.70253 0.000012 0.513187 0.000004 10.7 265-91 0.70253 0.000012 0.513179 0.000004 10.6 265-103 0.70245 0.000012 0.513162 0.000004 10.2 265-109 0.70249 0.000016 0.513145 0.000007 9.9 265-125 0.70246 0.000045 0.513154 0.000006 10.1 264-20 0.70244 0.513179 0.000004 10.6 265-49 0.70244 0.513196 0.000007 10.9 E. Limb Andesites 264-14 0.70243 0.000014 0.513153 0.000005 10.0 265-25 0.70254 0.000018 0.513152 0.000005 10.0 265-90 0.70250 0.000011 0.513159 0.000006 10.2 265-100 0.70249 0.000012 0.513179 0.000004 10.6 266-54 0.70246 0.513193 0.000005 10.8 E. Limb Dacites 264-09 0.70254 0.000019 0.513171 0.000008 10.4 265-40 0.70247 0.000011 0.513149 0.000008 10.0 265-42 0.70246 0.000015 0.513140 0.000006 9.8 265-64 0.70258 0.000015 0.513185 0.000007 10.7 265-70 0.70248 0.000024 0.513141 0.000005 9.8 265-83 0.70253 0.000015 0.513147 0.000007 9.9 265-84 0.70251 0.000011 0.513179 0.000007 10.6 265-85 0.70248 0.000019 0.513148 0.000005 9.9 265-95 0.70247 0.000019 0.513165 0.000006 10.3 266-53 0.70247 0.000015 0.513175 0.000010 10.5
59
Table 2-3. Starting compositions for modeling Sample # 264-08 265-43 265-113 SiO2 50.1 50.5 51.9 TiO2 2.68 1.92 2.17 Al2O3 12.7 13.9 13.4 FeO 14.1 11.6 12.8 MnO 0.26 0.21 0.23 MgO 5.69 6.98 5.93 CaO 9.58 11.14 9.48 Na2O 3.27 2.86 3.28 K2O 0.21 0.13 0.22 P2O5 0.28 0.19 0.28 Cl 0.07 0.01 0.02 Total 99.24 99.67 99.72 Trace Elements (ppm) Li 10 8 11 Sc 42 42 38 V 450 347 323 Cr 18 109 32 Co 43 41 39 Ni 34 54 34 Cu 60 59 51 Zn 116 94 103 Ga 21 18 21 Cs 0.03 0.01 0.03 Rb 2.1 1.1 2.2 Ba 17 8 16 Th 0.34 0.18 0.40 U 0.13 0.08 0.15 Nb 5.5 3.1 5.7 Ta 0.37 0.21 0.36 La 6.8 4.5 7.6 Ce 21 14 24 Pr 3.3 2.3 3.9 Sr 126 120 111 Nd 17 13 20 Zr 180 126 229 Hf 4.8 3.4 5.7 Sm 6.0 4.4 6.7 Eu 1.9 1.5 2.0 Tb 1.5 1.1 1.6 Dy 10 7 10 Y 59 44 61 Ho 2.0 1.5 2.2 Er 5.9 4.4 6.3 Yb 6.0 4.4 6.2 Lu 0.93 0.68 0.95 Tm 0.91 0.67 0.96 Gd 7.7 5.8 8.5 Pb 0.61 0.39 0.95
60
Table 2-4. Partition coefficients for Rayleigh fractional crystallization, partial melting and AFC calculations Andesite Partition Coefficients Basalt Partition Coefficients Element Olivine CPX Plag Apatite Ilmenite Amphibole Olivine CPX Plag Ilmenite Rb 0.01 0.02 0.025 0.001 0.034 0.04 0.0003 0.0004 0.056 0.034 Ba 0.01 0.02 0.155 0.12 0.00034 0.1 0.00001 0.0003 1.45 0.00034 Th 0.01 0.01 0.19 1.28 0.00055 0.15 7.00E-06 0.0021 0.13 0.00055 U 0 0 0.34 1.4 0.0082 0.008 9.00E-06 0.001 0.051 0.0082 Nb 0.00017 0.005 0.033 0.0011 2 0.28 0.00005 0.0089 0.045 2 Ta 0.00002 0.014 0.11 0.003 1.7 0.27 0 0.013 0.066 1.7 La 0.00006 0.062 0.082 11.4 0.000029 0.027 0.0002 0.054 0.13 0.000029 Ce 0.00006 0.116 0.072 12.9 0.000054 0.0293 7.00E-05 0.086 0.11 0.000054 Sr 0.00217 0.08 2.7 4.3 0 0.28 0.00004 0.091 1.4 0 Nd 0.00015 0.33 0.045 32.8 0.00048 0.0325 0.0003 0.19 0.066 0.00048 Zr 0.00450 0.14 0.0009 0.042 0.29 0.26 0.001 0.26 0.048 0.29 Hf 0.00370 0.21 0 0.014 0.38 0.43 0.0029 0.33 0.051 0.38 Sm 0.00044 0.41 0.033 16.1 0.00059 0.024 0.0009 0.27 0.054 0.00059 Eu 0.00056 0.57 0.55 25.5 0.009 0.0498 0.0005 0.43 0.65 0.009 Dy 0.00250 0.94 0.034 34.8 0.01 0.0136 0.0027 0.44 0.024 0.01 Y 0.00380 0.9 0.01 7.1 0.0045 0.0196 0.0082 0.47 0.013 0.0045 Yb 0.05600 0.63 0.014 15.4 0.17 0.102 0.024 0.43 0.0079 0.17 Lu 0 0.605 0.039 3.92 0.084 0 0.016 0.56 0.06 0.084
Primary Reference
Zanetti et al., 2004
Klein et al., 2000
Dunn and Sen, 1994
Prowatke and Klemme, 2006
Zack and Brumm, 1998
Bottazzi et al., 1999
Halliday et al., 1995
Halliday et al., 1995
Dunn and Sen., 1994
Zack and Brumm, 1998
Secondary Reference
Gill, 1979 Gill, 1979
Rollinson, 1993
Fujimaki, 1986
Rollinson, 1993
Haase et al., 2005
Rollinson, 1993
Rollinson, 1993
Rollinson, 1993
Rollinson, 1993
61
Table 2-5. AFC modeling parameters Parameter Abbreviation Value Units Magma Liquidus Temp tlm 1200 deg C Magma Temp tmo 1200 deg C Assm. Liquidus Temp tla 1100 deg C Country Rx Temp tao 720 deg C Solidus Temp ts 875 deg C
Magma Spec. Heat cpm 1484 J/Kg K
Assm. Spec. Heat cpa 1388 J/Kg K
Crsyt. Enthal Hcry 396000 J/Kg Fusion Enthal Hfus 354000 J/Kg Equilibration Temp Teq 900 deg C
62
A
Figure 2-1. Bathymetric maps showing the tectonic setting of the MOR dacites (data from GeoMapApp; Carbotte et al., 2004). Boxes show the general locations of high-silica lavas on each ridge. Dacites are commonly associated with the ends of ridge segments, such as overlapping spreading centers (OSC) and propagating ridge tips at ridge transform intersections. A) 9°N OSC on the East Pacific Rise (EPR) B) Propagating Ridge Tip on the Juan de Fuca Ridge (JdFR) and Axial seamount and C) possible OSC near the Propagating Ridge Tip on the Galapagos Spreading Center (GSC).
63
B
Figure 2-1. Continued
64
C
Figure 2-1. Continued
65
Figure 2-2. Comparison of major and trace element compositions in MOR high-silica andesites and dacites from the East Pacific Rise, Juan de Fuca Ridge, and Galapagos Spreading Center. MOR dacitic lavas have similar major and trace element compositions, while andesites have more variable compositions that lie between dacites and highly evolved MORB. A) Mantle-normalized diagram showing similarities in trace element compositions between dacites from three different ridges and an andesite from Axial Seamount. Average composition for N-MORB from the 9° 17’ – 10°N segment of the East Pacific Rise is shown for comparison. MOR dacites are characterized by low Nb and Ta and high U, Th, Zr and Hf relative to elements of similar incompatibilities. B) and C) Major element plots showing the range of compositions of MOR dacites compared to East Pacific Rise MORB. Gray field shows the range of compositions from >1200 analyses of MORB glasses from the East Pacific Rise north of the 9°N OSC (from PetDB, Sims et al., 2002, 2003 and Perfit unpublished data).
66
Figure 2-3. Bathymetric map of the 9°N OSC showing locations of samples collected during the MEDUSA2007 cruise. Samples are divided into rock types based on silica content (dacite >62wt% SiO2; andesite 57-62 wt % SiO2; basaltic andesite 52-57 wt% SiO2; basalts <52wt% SiO2 and FeTi basalts <52wt% SiO2 and >12 wt% FeO). Dacites are primarily found on-axis on the eastern limb of the OSC. 50 m contour intervals are shown.
67
Figure 2-4. Photographs of MOR high-silica lavas from A) and B) 9°N OSC East Pacific Rise, C) Galapagos Spreading Center and D) Juan de Fuca Ridge. Morphologically, dacites typically form blocky angular flows and large elongate pillow lavas with roughly corrugated or striated surfaces.
68
Figure 2-5. Major element variations versus MgO (wt.%) for dacites from the 9°N OSC on the East Pacific Rise (filled squares). Gray field represents all lavas collected in 2007 from the 9°N OSC. Dacites are compared to three low-pressure fractional crystallization trends (calculated using MELTS; Ghiorso & Sack, 1995) using parental compositions from OSC basalts (see text for modeling parameters and details). Not all dacite major element variations can be explained by fractional crystallization alone (e.g. Al2O3, K2O, and P2O5). Experimental compositions from partial melting of altered basalt (Beard and Lofgren, 1991) are shown for comparison (*).
A B
C D
E F
69
Figure 2-6. Variation diagram showing Cl (wt%) versus MgO (wt%) for OSC lavas. Superimposed are three liquid lines of descent (calculated using Melts; Ghiorso & Sack, 1995), showing that the maximum amount of Cl enrichment due to extensive fractional crystallization cannot produce the high Cl concentrations in the OSC dacites. The dashed rectangle represents the range of compositions that can be produced through 1-15% partial melting of an altered basalt with 350 ppm Cl (purple star). 350 ppm Cl is the median value of Cl analyzed in sheeted dikes from ODP Hole 504B with minimum and maximum values (49 and 650 ppm) shown with an error bar. Partition coefficients for Cl are from Gillis et al., (2003). The range of MgO values for partial melts were taken from experimental partial melts of less than 15% (Beard & Lofgren, 1991).
70
Figure 2-7. Trace element variations versus Zr (ppm) in 9°N OSC lavas. Superimposed on the diagrams are calculated trends for fractional crystallization (model 1 and 2 using the Rayleigh fractionation equation), 1 to 15 % partial melting (assuming batch melting), and AFC simulations (EC-AFC; Bohrson and Spera, 2001). See text for model parameters. Kinks in models represent changes in crystallizing phases.
71
Figure 2-8. Normalized trace element ratio diagrams showing the range of dacite compositions at the 9°N OSC. Fractional crystallization, partial melting and AFC trends are shown as in Figure 2-4. Tick marks on trend lines for partial melts are in 1% increments, while fractional crystallization tick marks represent 10% intervals. Concentrations are mantle-normalized (Sun & McDonough, 1989).
72
Figure 2-9. Radiogenic isotope ratios of 9°N OSC lavas. A) Pb-isotope ratios showing 9°N OSC dacites with 208Pb/204Pb and 206Pb/204Pb ratios similar to OSC N-MORB basalts and northern East Pacific Rise N-MORB (gray field; data from (Sims et al., 2002; Sims et al., 2003; Goss et al., 2010). B) Nd and Sr isotopes with 9°N OSC lavas that are similar to EPR and OSC N-MORB.
73
Figure 2-10. Elemental variation diagrams showing 9°N OSC dacites versus calculated liquid lines of descent produced during two alternative types of crystallization. Crosses show the evolution of a melt during in situ crystallization (Langmuir, 1989). Circles show the liquid line of descent of a magma that undergoes 50% fractional crystallization and is then separated (via filter pressing) from the phenocrysts (melt-segregation model). The resulting magma undergoes an additional 50% crystallization, until it once again separates from the crystals following the model of Bachman and Bergantz (2004). Kinks in models represent changes in crystallizing phases.
74
Figure 2-11. Partial melting model showing a range of possible parents (gray lines) that could produce the 9°N OSC dacites (bold red line) from 1-15% partial melting. Sheeted dikes and the upper parts the gabbroic layer may be composed of a range of compositions depending on the composition of the starting material and degree of alteration. Possible parental compositions were calculated using the batch melting equation and solving for the initial parent composition (Co). Three possible wall rock compositions are superimposed on the calculated parental range. Altered basalt 1 (squares; (Nakamura et al., 2007) provides the best match to the calculated source compositions and is used as the parent rock in subsequent partial melting and AFC models.
75
Figure 2-12. Mantle-normalized diagrams showing results of 1-15% partial melting of an altered basalt (see Figure 2-11) using two different modal minerologies. Partial melts were calculated using the batch melting equation. See text for more detail. A) Partial melting results using a non-amphibole-bearing gabbro assemblage (19% Ol, 30% Cpx, 50% Plag, 1% Ilm). B) Partial melting results using an amphibole-bearing gabbro (20% Cpx, 25% Opx, 49% Plag, 5% Amph, 1% Fe-Oxide; Haase et al., 2005). The amphibole-bearing gabbro provides the best fit to the 9°N OSC dacites.
76
Figure 2-13. Diagram showing the calculated effects of varying wall rock temperature on incompatible trace element composition (La) during AFC. A) Higher wall rock temperatures cause earlier onset of assimilation compared to lower initial temperatures. Higher initial temperatures produces lower overall incompatible element abundances compared to lower wall rock temperatures because the amount of fractional crystallization is lower. The average La concentration of the 9°N OSC dacites is ~28ppm, which can be produced by assimilating partial melts of a wall rock at starting temperatures of 650 to 720°C. B) Plot showing the ratio of assimilation to crystallization (Ma*/Mc) for various temperatures of wall rock. Lower ratios produce higher incompatible trace element concentrations in the hybrid melts.
77
Figure 2-14. Mantle-normalized trace element diagram showing results of the best-fit AFC model (thin black lines). This model requires a total of 73-85% fractional crystallization in combination with 5-20% assimilation of partially melted wall rock to produce trace element compositions similar to the 9°N OSC dacites (bold red line). Fractional crystallization is the dominant process until 68% of the magma has crystallized. This is followed by 5-20% assimilation of partial melts in conjunction with an additional 5-17% crystallization.
78
Figure 2-15. Cartoon showing a possible scenario for dacite formation on MOR. A) Injection of basaltic magma into a ridge segment end through dike propagation. B) Magma supply rates diminish at segment end, abandoning pockets of magma and allowing for extensive fractional crystallization. The latent heat of crystallization begins to heat up and partially melt the hydrothermally altered wall rock. C) Partial melts of wall rock are assimilated or mixed into the evolving magma chamber, resulting in dacitic magmas.
79
CHAPTER 3 ROLE OF ASSIMILATION IN THE PETROGENESIS OF LAVAS ON MID-OCEAN
RIDGES INFERRED FROM CL, H2O, CO2 AND OXYGEN ISOTOPE VARIATIONS
Introduction
Crustal assimilation has been proposed as an important process in the
petrogenesis of mid-ocean ridge (MOR) magmas (e.g., O'Hara, 1977; Michael &
Schilling, 1989; Michael & Cornell, 1998), but it is largely ignored as a primary igneous
process in ridge settings for several reasons. First, contamination is commonly
overlooked because most geochemical variations in MOR lavas can be readily
explained by fractional crystallization, variations in mantle melting parameters, or
differences in mantle source compositions. Second, the low volume of assimilated melt
compared to the more voluminous mid-ocean ridge basalt (MORB) magmas makes
assimilation a difficult process to identify in the erupted lavas. Third, the magma and
wall rock may have similar major and trace element compositions, resulting in liquids
(assimilants) that are geochemically difficult to discriminate from typical MORB lavas.
However, variable degrees of hydrothermal alteration of basaltic crust can produce
significant changes in fluid mobile element concentrations and isotopic ratios depending
on the water/rock ratios (e.g., Alt & Teagle, 2000). Therefore, components that are
particularly sensitive to seawater interaction, such as Cl, U, H2O, and oxygen isotopes,
can be used to determine the extent to which crustal assimilation is involved in MOR
magmatism.
Contamination of MOR magmas by a seawater-altered component was first
proposed based on excess Cl in MORB glasses (Michael & Schilling, 1989; Michael &
Cornell, 1998). The elevated Cl concentrations compared to elements of similar
incompatibility (K, Nb, Ti) in fresh MORB glass cannot be explained by post-eruption
80
alteration or fractional crystallization and are instead attributed to assimilation of a
seawater-derived component such as saline brines or altered ocean crust (Michael &
Schilling, 1989; Michael & Cornell, 1998). Consequently, Cl over-enrichment has been
identified in many submarine settings, including MOR (Perfit et al., 1999; Coogan et al.,
2003b; le Roux et al., 2006; Wanless et al., 2010; Wanless et al., accepted), back-arc
basins (Kent et al., 2002; Sun et al., 2007) and ocean islands (e.g., Kent et al., 1999).
Despite the clear evidence of assimilation in these lavas, many models for the
magmatic plumbing system at MORs continue to ignore this process.
An alternate approach to identifying crustal contamination on MORs is through
analyses of oxygen isotope ratios. Low oxygen isotope ratios relative to mantle values
are observed in many lavas from Icelandic volcanoes (e.g., Gautason & Muehlenbach,
1998). Lower than mantle values have also been observed in melt inclusions from
Hawaiian glasses and are attributed to assimilation of altered Pacific crust (Eiler et al.,
1996; Garcia et al., 1998). Decreases in δ18O relative to primary mantle-like values are
attributed to the temperature dependent fractionation of oxygen isotopes between
seawater and mineral phases in ocean crust during high-temperature hydrothermal
alteration (Muehlenbach & Clayton, 1972; Alt et al., 1996). Therefore, assimilation of
crust altered at high-temperatures near MORs should lower the oxygen isotope ratios of
the resulting magma. Unfortunately, this signal may be hard to identify in many MORB
magmas because the mass of assimilant relative to parent magma is generally not
sufficient to appreciably change the oxygen isotope ratios. However, it may be an
important discriminator in magmas that have undergone significant assimilation.
81
Here, we examine Cl, H2O, and CO2 concentrations and oxygen isotope ratios
from a suite of lavas collected at the 9°N overlapping spreading center (OSC) on the
East Pacific Rise (EPR) to place constraints on the role of crustal assimilation at MORs.
This suite includes a nearly continuous range of compositions from basalts to dacites,
including one of the most evolved lava compositions sampled on a MOR (>67 wt%
SiO2). Major and trace element data indicate that assimilation of altered ocean crust is a
critical process for the formation of MOR dacites (Wanless et al., accepted) and as
such, it should be reflected in the Cl, H2O, and CO2 concentrations and oxygen isotope
ratios of these lavas. Additionally, we use geochemical results to better define the
source and depth of assimilation processes beneath MORs and use CO2 concentrations
to provide information on degassing of magmas during extensive fractional
crystallization and the ascent rates of high-silica magmas.
Geologic Setting
The 9°N OSC is located between the Clipperton and Siqueiros transform faults on
the EPR (Figure 3-1). It is a second order ridge discontinuity consisting of limbs that
overlap by ~27 km and offset the ridge axis ~8 km from east to west (Sempere &
Macdonald, 1986). The OSC has been migrating southward down the ridge axis at a
rate of approximately 42 km/Myr as the eastern limb propagates into older crust and the
western limb recedes or dies (Macdonald & Fox, 1983; Carbotte & Macdonald, 1992).
The 9°N OSC is one of the most extensively studied OSC on the MOR system. It
has been the focus of several geophysical studies (Detrick et al., 1987; Harding et al.,
1993; Kent et al., 1993; Kent et al., 2000; Bazin et al., 2001; Dunn et al., 2001; Tong et
al., 2002), which have produced the first 3D multi-channel seismic survey of a mid-
ocean ridge (Kent et al., 2000) and a 3D seismic refraction study (Dunn et al., 2001).
82
These studies reveal a shallow melt lens beneath both limbs of the OSC and in the
interlimb region north of overlap basin (Kent et al., 2000). The western, receding limb
melt lens is narrow and shows no significant variation in depth along axis (Kent et al.,
2000), while the melt lens beneath the eastern, propagating limb shows variations in
both width and depth. Beneath the southern portion of the east limb the melt lens is
narrow and significantly deeper than the rest of the eastern ridge axis, plunging ~500 m
southward over ~6 km (Kent et al., 2000). North of the overlap basin, the melt lens is
anomalously wide (> 4 km) and is not centered directly below the ridge axis, instead
extending from the axis ~4 km to the west (Kent et al., 2000; Tong et al., 2002).
Although the depth of the lens varies along axis, the top of the melt lens appears to
follow the base of the sheeted dikes, at approximately 1.5-2 km beneath the seafloor
(Kent et al., 2000; Tong et al., 2002).
The first lava sampling in this region occurred during the CHEPR dredging and
wax coring cruise that recovered several high-silica lavas, along with basalts and FeTi
basalts (Langmuir et al., 1986). More recently, the 9°N OSC was the focus of the
MEDUSA2007 research cruise (AT15-17), which completed detailed mapping using the
DSL-120A side-scan system (White et al., 2009), and the WHOI TowCam (Fornari,
2003) and extensive sampling using the ROV Jason2, (Wanless et al., accepted;
Wanless et al., in prep-b) of the region. Results of this cruise (White et al., 2009;
Wanless et al., accepted; Wanless et al., in prep-b) revealed a range of rock types from
basalts to dacites, but the high-silica lavas are confined to the eastern propagating limb
of the OSC (Wanless et al., accepted). Additionally, the basalts erupted at the OSC are
83
dominantly ferrobasalts, in contrast to the dominantly MORB lavas erupted from the
9°15 to 10° N section of the EPR (Perfit et al., in prep).
Analytical Methods and Results
Major and Trace Elements
During the MEDUSA2007 research cruise, 275 glassy samples were collected
from the 9°N OSC. Methods, standards and results from all major and trace element
analyses are discussed in detail in Wanless et al., (accepted) and Wanless et al., (in
prep.). Major elements were analyzed on a JEOL 8900 Electron Microprobe at the
USGS in Denver, Colorado. High-precision Cl and K2O concentrations were determined
using 200-second peak/100 second background counting times. Samples were
analyzed for trace element concentrations on an Element2 Inductively Coupled Plasma
Mass Spectrometer (ICP-MS) at the University of Florida. Concentration data and ratios
used in this paper are listed in Table 3-1.
Here, we use the general term “basalt and basaltic” to include basalts sensu
stricto, ferrobasalts and FeTi basalts. The intermediate lavas include both basaltic
andesites and andesites, with SiO2 concentrations that range from 52–57 wt% and 57-
62 wt%, respectively. Dacites have >62 wt% SiO2. Cl concentrations range from 0.01 to
0.07 wt% in the basalts and are generally higher in the basaltic andesites (0.01 to 0.31
wt %), andesites (0.20 to 0.42 wt %) and dacites (0.23 to 0.70; Figure 3-2). K2O
concentrations range from 0.13 to 1.37 wt % and generally increase with increasing
SiO2 content. Basalts range from 0.13 to 0.21 wt%, basaltic andesites range from 0.26
to 0.60 wt%, andesites range from 0.63 to 0.83 wt% and dacites range from 0.89 to 1.37
wt% (Figure 3-2). Cl/K2O ratios range from 0.04 to 0.60 (Figures 3-3 and 3-4).
84
Volatile Elements
A representative subset of 20 samples (covering the range of rock types) from the
9°N OSC was selected for volatile analyses (Table 3-1). Several glass chips were
handpicked from each sample, avoiding alteration and microphenocrysts. Samples were
analyzed for H2O and CO2 concentrations by Fourier transform infrared (FTIR)
spectroscopy at the University of Oregon (Johnson et al., 2009). Water concentrations
were calculated either using the fundamental OH- stretching vibration at 3559 cm-1 or
from the average of the two molecular water peaks (1630 cm-1 and 5200 cm-1) and the
4500 cm-1 OH- peak. An absorption coefficient of 63 L/mol cm was used for the 3550
cm-1 peak (Dixon et al., 1995b; Dixon et al., 1995a), and absorption coefficients for the
near-IR peaks were calculated based on major element concentrations following
methods in Mandeville et al., (2002). CO2 concentrations were measured using the
carbonate peaks at 1515 and 1430 cm-1, using background subtraction procedures
described in Johnson et al., (2009) and absorption coefficients calculated from Dixon &
Pan (1995).
Volatile concentrations are highly variable in the OSC lavas, but are consistent
with major element trends (Table 3-1). H2O concentrations range from 0.23 to 0.39 wt%
in basaltic lavas and from 0.24 to 1.56 in basaltic andesite samples (Figure 3-2).
Andesites have H2O concentrations ranging from 0.99 to 1.50 wt% and dacites range
from 1.53 to 2.35 wt%. CO2 concentrations in basalts range from 131 to 256 ppm
(Figure 3-2). Two of the basaltic andesites have CO2 concentrations (232 and 184
ppm) but otherwise all others had CO2 concentrations below the detection limits of ~25
ppm. H2O/Ce ratios, involving elements of similar magmatic incompatibility, generally
increase with increasing silica and range from 0.014 to 0.024 (Figure 3-3), while
85
H2O/K2O ratios vary from 0.095 to 2.59, with the highest ratios observed in the basaltic
andesites (Figure 3-4).
Oxygen Isotope Analyses and Results
Oxygen isotope ratios (δ18O, per mil notation) were determined on 26 fresh,
microphenocrysts-free glass chips that cover the range of rock types (Table 3-1) at the
CO2-laser-fluorination laboratory at the University of Wisconsin, Madison, following
methods described in Valley et al., (1995). Aliquots of 2.4-3.2 mg were treated with BrF5
overnight, and then individually heated with a CO2 laser in the presence of BrF5.
Measurements were standardized with 4-5 analyses of UWG-2 garnet standard per day
(δ18O = 5.8‰; Valley et al. 1995), and are reported in standard δ-notation relative to
Standard Mean Ocean Water (SMOW). Reproducibility of the standard during each
session was better than ±0.15‰ (2SD). The δ18O values range from 5.31 to 6.19 ‰ in
the OSC lavas with an average of 5.79 ‰ (Figure 3-5). The basalts and basaltic
andesites have similar oxygen isotope ratios, ranging from 5.51 to 5.79 ‰ and 5.31 to
5.92 ‰ respectively. The andesites and dacites have variable δ18O (5.38 to 6.19 ‰),
with a mean of 5.86 ‰.
Discussion
Magma Crystallization versus Assimilation-Fractionation-Crystallization Processes
Liquid lines of descent (LLD’s) were calculated using the MELTS thermodynamic
modeling program to simulate fractional crystallization at fO2 = QFM and P = 1 kbar
(Ghiorso & Sack, 1995). These results suggest that fractional crystallization alone
cannot account for the high Cl and H2O concentrations observed in the MOR dacites,
andesites or basaltic andesites (Figure 3-2). H2O concentrations in the dacites are as
86
much as two times greater than model predictions. Similarly, Cl concentrations
observed in the dacites are more than ten times greater than model predictions (Figure
3-2). H2O/Ce ratios, which should not change over a wide range of fractional
crystallization, are generally higher than expected in the MOR dacites, having almost
two times higher ratios than the basalts, whereas basaltic andesites and andesites
show a wider range of ratios (Figure 3-2a).
The high Cl concentrations, however, are similar to those produced by 1-15%
partial melting of altered basaltic crust (Wanless et al., accepted). While both H2O and
Cl exhibit over-enrichments compared to calculated fractional crystallization trends
(Figure 3-2), the Cl over-enrichment is much greater than that of H2O. The difference
between these enrichment factors may be due to the higher Cl contents in the altered
crustal source and, to a lesser extent, variable amounts of H2O degassing during
magma ascent and eruption on the seafloor (see degassing section below). In contrast
to H2O, Cl remains soluble in most basaltic and andesitic lavas at eruption depths
greater than 700 meters below sea level (Unni & Schilling, 1978; Webster et al., 1999)
and therefore, does not degas.
Le Roux et al., (2006) used Cl/Nb ratios to assess the role of assimilation in
MORB magmas because Cl and Nb have similar partition coefficients in basaltic
systems. Cl is enriched in seawater-altered crust while Nb remains immobile during
seawater alteration, which allows for discrimination between the effects of fractional
crystallization and assimilation. However, the advanced fractional crystallization in the
OSC lavas results in precipitation of Fe-Ti oxides in which Nb is a compatible element,
so the Cl/Nb ratio has limited applicability in the OSC lavas. Here, we use Cl/K2O
87
(Figure 3-3) because these elements also have broadly similar incompatibilities over a
wide range of crystallization (e.g., Kent et al., 1999) and this ratio has been used in
several studies to identify crustal contamination (e.g., Michael & Cornell, 1998; Kent et
al., 1999). K2O concentrations are not affected by high temperature alteration in the
lower sheeted dikes and upper gabbros (Alt et al., 1996). Many of the MOR dacites
have Cl/K2O ratios that are greater than five times the ratios observed in the spatially
related basalts (Figure 3-3). Additionally, many andesites and basaltic andesites
erupted at the OSC have Cl/K2O ratios three to seven times higher than the basalts
erupted in the region and higher than mantle values (0.065) predicted by Michael and
Cornell (1998), suggesting they have also been contaminated by crustal material.
These observations clearly suggest the operation of assimilation in the formation of
basaltic andesites, andesites and dacites on MOR.
Evidence for Assimilation from Oxygen Isotopes
Crystallization of Fe-Ti oxides leads to a significant increase in oxygen isotope
ratios of evolving magmas (Taylor, 1968) because these phases and to a lesser extent
Fe-Mg silicate phases preferentially incorporate 16O relative to 18O compared to the
remaining melt (Taylor, 1968; Anderson et al., 1971; Muehlenbach & Byerly, 1982).
During crystallization of MORB magma there is a slight increase in the δ18O as the
magma crystallizes ferromagnesian silicates, followed by a dramatic increase when Fe-
Ti oxides precipitate (Matsuhisa et al., 1973). Fractional crystallization can result in an
increase of δ18O values by 1-1.5‰ (e.g., Muehlenbach & Byerly, 1982).
Using modal mineral proportions calculated from MELTS (Ghiorso & Sack, 1995)
and oxygen isotope fractionation factors from Bindeman et al., (2008), we have
88
calculated the LLD for δ18O during fractional crystallization of a MORB parental magma
(Figure 3-5). These calculations suggest that if the MOR dacites formed through
fractional crystallization the δ18O should be ~6.8‰; however, the measured ratios are
~1‰ lower (dacite average = 5.86‰), supporting the conclusion that the dacites could
not form from fractional crystallization alone. Similar observations have been made for
dacites erupted at the Galapagos Spreading Center (δ18O = 3.9–6.2‰; Perfit et al.,
1999).
The temperature-dependent fractionation of oxygen isotopes between seawater
and mineral phases causes decreases in δ18O of ocean crust relative to primary mantle-
like values during high-temperature hydrothermal alteration, but increases δ18O during
low-T alteration (<200-250°C; Muehlenbach & Clayton, 1972; Alt et al., 1996).
Additionally, as seawater migrates through and reacts with the crust at temperatures <
200-250°C, the δ18O of that water evolves to lower values. This results in an overall
decrease in δ18O with depth in the crust. Profiles of oxygen isotope ratios in exposed
sections of ocean crust (Hess Deep; Agrinier et al., 1995) and drill holes (e.g., ODP Site
504; Alt et al., 1996) show that δ18O values are higher than unaltered MORB
compositions (~5.6 +- 0.2 ‰; Eiler, 2001) in the upper volcanic section of the crust, but
decrease to lower than the mantle values in the lower sheeted dikes and gabbros (for
profiles see Alt & Teagle, 2000). Therefore, if assimilation occurs at the base of the
sheeted dikes or in the upper gabbros, it should lower the oxygen isotope ratios of the
resulting magma.
High temperature hydrothermal alteration of the ocean crust results in an observed
decrease in the δ18O of the sheeted dike and gabbro layers to ~ 4‰ (Alt et al., 1996)
89
compared to mantle values of 5.6‰ (e.g., Alt et al., 1986; Eiler, 2001). Therefore,
partial melting of this material will produce melts with oxygen isotope ratios less than
mantle-like values, whereas assimilation of this material into fractionating magma will
produce magmas with oxygen isotope ratios between altered basalt and that expected
from fractional crystallization alone. The δ18O value of the MOR andesites and dacites
range from 5.38 to 6.19‰, which is much more variable than the spatially related
basalts and lower than calculated fractional crystallization trends (Figure 3-5).
Taylor (1968) suggested that to a first approximation, the effect of assimilation on
oxygen isotope ratios can be determined using mass balance equations. Using results
from Energy Constrained-AFC petrologic modeling calculations (Bohrson & Spera,
2001), the ratio of fractionating magma to assimilant in the dacites is 2:1(Wanless et al.,
accepted). Assuming the evolved magma has a δ18O of 6.8 (largely due to fractionation
of silicates and iron oxides) and a crustal assimilant with a δ18O of 4 (due to high
temperature seawater alteration; Alt et al., 1996), the resultant oxygen isotope ratio of
the AFC magma would be ~5.9. This value is similar to the average of ratios observed
in the MOR dacites and is less than predicted by fractional crystallization alone.
CO2 and H2O Degassing, Magma Ascent Rates and Depth of Assimilant
Although the geochemical data are consistent with crustal assimilation, these
signatures do not constrain where this contamination occurs. The depth of equilibration
of vapor-saturated melts can be calculated using the H2O and CO2 solubility model of
Dixon et al., (1995a,b). This model is based on experimental results at pressures and
temperatures similar to MOR magmatic conditions. A potential complication of this
modeling is that CO2 and H2O may undergo variable degassing during ascent from the
magma chamber to the seafloor, which is caused by slow diffusion of CO2 to the
90
nearest gas bubble during ascent. However, the rapid ascent rates in MORB magmas
and quenching of the lavas at the seafloor will trap the volatiles and allow for very little
degassing (e.g., Dixon et al., 1988). Therefore, these model calculations can provide
minimum depths of equilibration of the melts prior to eruption. Using this model, we can
calculate the equilibrium pressure of a vapor-saturated melt (i.e., magma chamber
depth) of a given composition prior to eruption on the seafloor using the VOLATILECALC
program (Newman & Lowenstern, 2002).
H2O and CO2 concentrations in the OSC lavas suggest a range of equilibration
pressures with a maximum of ~550 bars. Most of the basaltic lavas have equilibration
pressures that approximately equate to the top of the imaged melt lens (~1.5 km or 450
bars; Kent et al., 2000), suggesting little to no degassing during ascent and eruption on
the seafloor. In contrast, MOR andesites and dacites have completely degassed CO2,
and some may have also lost H2O (Figure 3-6). The evolved lavas have equilibrium
pressures approximately equal to that of the seafloor (~250 bars), suggesting slower
ascent rates and higher degrees of degassing prior to eruption.
The vapor saturation calculations support the hypothesis that the depth of
contamination on MOR occurs at the top of the axial magma chamber (le Roux et al.,
2006). Textural observations in ophiolites show melting and assimilation occurs at the
roof of the magma chamber, which may migrate within the ocean crust but often
coincide with the base of the sheeted dikes (Coogan et al., 2002; Gillis & Coogan, 2002;
Gillis, 2008).
The hydrostatic pressure at the seafloor results in extensive CO2 degassing from a
CO2-saturated basaltic melt, however, quenched lavas supersaturated with CO2 are
91
common in submarine settings (Dixon et al., 1988; Dixon et al., 1995b; Dixon et al.,
1995a; Saal et al., 2002; le Roux et al., 2006). This has been attributed to rapid ascent
rates of basaltic magmas, which do not allow for complete vapor exsolution. In contrast,
slow ascent rates or pooling of isolated magma batches at shallow depths could lead to
magma degassing (Dixon et al., 1988). A key factor that affects the ascent and effusion
rate of magma is viscosity. The andesites and dacites erupted on the seafloor have
higher viscosities than MORB magmas, which may allow for significant bubble
nucleation and growth prior to eruption. Magma degassing during ascent or eruption of
the high-silica dacites is supported by large elongate vesicles observed in hand
samples and low CO2 contents in the glasses.
Source of Assimilant
Despite growing evidence of crustal contamination on MORs, the source of the
assimilant and depth of the process are poorly understood. Glass compositions suggest
that assimilation of brines is responsible for the Cl contamination in some EPR MORB
(le Roux et al., 2006); however, assimilation of partially melted, altered crust may also
result in anomalous Cl enrichments (e.g., Michael & Schilling, 1989). Partial melting or
thermal breakdown of Cl-bearing amphibole in altered crust may result in elevated Cl
concentrations in the resulting magma (Michael & Schilling, 1989) and has been
suggested as a possible source of Cl enrichment in Galapagos Spreading Center
andesites and dacites (Perfit et al., 1999) and in dacites from the 9°N OSC (Wanless et
al., accepted).
As mentioned above, Cl/K2O and H2O/K2O ratios provide a means to discriminate
between sources of contamination on MOR because of the variable concentrations of
these elements in possible assimilants (Kent et al., 1999). Potential contaminants on
92
MOR include altered basaltic crust (Cl = 0.1 wt%, H2O = 5 wt%, K2O = 1 wt%), seawater
(Cl = 1.935 wt%, H2O = 97.5 wt%, K2O = 0.04 wt%); 15% NaCl brine (Cl = 9.9 wt%, H2O
= 85 wt%, K2O = 0.25 wt%); and 50% NaCl brine (Cl = 30.3 wt%, H2O = 50 wt%, K2O =
0.25 wt%; see Kent et al., 1999 and references therein). Fluids enriched with as much
as 50% NaCl have been observed in melt inclusions in several MOR gabbros (Kelley &
Delaney, 1987). These brines form from high temperature phase separation of
seawater during hydrothermal circulation (e.g., Berndt & Seyfreid, 1990) and may be
trapped along grain boundaries or pore spaces within the altered crust (Michael &
Schilling, 1989).
Bulk mixing of any of these contaminants with OSC basalt cannot produce the
observed compositions of the MOR dacites (Figure 3-4). Instead, most dacites have
lower H2O/K2O ratios and higher Cl/K2O ratios compared to the possible assimilants.
However, if we assume that the dacites formed from AFC processes, then assimilation
of low-degree partial melts (5-10%) of altered basalt into a fractionating MORB magma
should produce compositions similar to the OSC dacites. This process can explain the
elevated Cl/K2O ratios observed in the MOR dacites. The higher H2O/K2O ratios
observed in several MOR dacites (Figure 3-4) suggest that the magmas may have also
interacted with small volumes of a saline brine. For instance, mixing of ~0.2 wt% of a
50% NaCl brine can produce H2O/K2O concentrations similar to the MOR dacites and
for this small amount of brine, the change in δ18O of the magma would be negligible
(<<0.01).
Oceanic Plagiogranites
Plagiogranites veins are a volumetrically small but ubiquitous component of the
ocean crust and have been observed in ophiolites (e.g. Pedersen & Malpas, 1984), drill
93
cores from the ocean crust (e.g. Casey, 1997; Dick et al., 2000), and as xenoliths in
Icelandic lavas (Sigurdsson, 1977). There are also many examples of evolved plutonic
rocks from slower spreading centers (e.g., Mid-Atlantic Ridge; Aumento, 1969), which
may suggest AFC or partial melting processes are occurring on much smaller scales,
deeper in the ocean crust. The origin of these veins remains unclear but two main
hypotheses are: 1) partial melting of gabbroic crust (e.g. Koepke et al., 2004; Koepke et
al., 2007) and 2) extreme crystal fractionation of tholeiitic magmas (Beccaluva et al.,
1999, Coleman & Donato, 1979, Niu et al., 2002).
The composition of oceanic plagiogranites varies considerably (Koepke et al.,
2007), making a direct petrologic comparison between the MOR dacites and the silicic
veins difficult. The mantle-like δ18O values of zircons collected from plagiogranite veins
in the gabbroic crust at the mid-Atlantic ridge (~5.2 ± 0.2 ‰) suggest little to no
seawater contamination in the evolved melt at these depths (Grimes et al., 2010a), but
lower δ18O values in plagiogranite veins from the Oman ophiolite (average of 4.6± 0.6
‰) are thought to represent remelting of altered ocean crust (Grimes et al., 2010b).
These studies suggest that high-silica lavas can form in a variety of different MOR
settings and may have a range of different compositions but that melting is an important
process in their petrogenesis.
Conclusions
Variations in volatile concentrations and δ18O in 9°N OSC lavas suggest that these
magmas have experienced assimilation during their petrogenesis, with the most
extreme signatures observed in high-silica andesites and dacites and little evidence in
basaltic lavas. H2O concentrations are up to two times higher in dacitic lavas compared
to calculated fractional crystallization trends, whereas Cl has excesses of seven to ten
94
times predicted values. δ18O values are on average ~1‰ lower than ratios expected
from fractional crystallization of ferromagnesian silicates and Fe-Ti oxide phases,
consistent with assimilation of an additional component or components.
The source of the excess H2O and Cl and low δ18O values is partially melted,
hydrothermally altered oceanic crust, but may also include small volumes of saline
brines produced during two-phase separation of high-temperature hydrothermal fluids.
Vapor saturation pressures calculated from H2O-CO2 data suggest that assimilation
most likely occurs at the top of the melt lens, which at the 9°N OSC, corresponds
approximately to the base of the sheeted dikes.
Basaltic lavas were supersaturated with CO2 at their eruption depths suggesting
fast ascent rates. In contrast, high-silica lavas were completely degassed CO2 and
variably degassed H2O prior to quenching on the seafloor. This suggests slower ascent
rates and/or lower effusion rates for the high-silica lavas, which is consistent with their
higher viscosities and the presence of large elongate vesicles.
95
Table 3-1. Geochemical data
Sample SiO2
TiO2
Al2O3 FeO
MgO
CaO
Na2O
K2O
P2O5 Cl H2O CO2 d18O Ce Cl/K2O
H2O/K2O
H2O/ Ce
264-08 50.45 2.70 12.86 14.07 5.69 9.51 3.33 0.21 0.28 0.07 0.39 231 5.79 20.5 0.35 1.82 0.019 265-43 50.12 1.92 13.86 11.57 6.83 11.13 2.81 0.13 0.19 0.01 0.25 224 14.2 0.10 1.87 0.018 265-82 50.77 1.95 13.81 11.70 6.93 11.21 2.69 0.14 0.19 0.01 0.24 256 5.51 13.0 0.07 1.74 0.019 265-88 50.85 2.02 13.88 11.99 6.71 11.06 2.85 0.15 0.21 0.02 0.26 131 5.71 15.0 0.13 1.72 0.017 266-51 50.55 1.94 13.85 11.59 7.33 10.97 2.86 0.13 0.19 0.01 0.23 218 13.3 0.09 1.82 0.017 265-104 51.25 1.90 13.79 11.65 6.90 10.59 3.00 0.16 0.22 0.05 0.31 219 16.9 0.29 1.91 0.018 264-18 52.15 1.77 14.04 10.89 6.05 9.94 3.27 0.26 0.20 0.01 0.24 232 5.72 13.4 0.04 0.95 0.018 265-106 53.33 1.94 13.26 12.35 5.11 8.71 3.67 0.32 0.39 0.19 0.65 184 5.77 33.9 0.59 2.02 0.019 265-103 56.50 2.01 12.33 13.74 2.74 6.45 3.77 0.52 0.65 0.30 1.24 0 5.73 60.4 0.58 2.38 0.020 265-91 56.83 2.10 12.31 14.23 2.09 6.24 3.45 0.60 0.78 0.31 1.56 0 5.31 66.1 0.52 2.59 0.024 265-125 55.59 1.66 13.63 10.73 4.89 8.30 3.64 0.41 0.20 5.92 32.7 265-90 58.08 1.91 12.43 13.69 1.74 5.76 3.51 0.66 0.74 0.34 0.99 0 73.2 0.51 1.50 0.014 265-100 58.09 1.76 12.64 12.68 1.89 5.51 3.82 0.63 0.54 5.56 72.7 266-54 59.65 1.72 13.22 11.25 2.28 5.60 3.91 0.75 0.43 0.42 1.50 0 5.83 67.9 0.56 2.01 0.022 264-14 61.75 1.30 13.46 8.71 2.47 5.54 3.94 0.83 0.22 6.06 57.0 265-69 61.02 1.72 13.62 9.97 1.91 5.38 3.89 0.80 0.27 0.20 1.44 0 55.8 0.25 1.80 0.026 265-63 64.43 1.29 13.26 8.22 1.29 4.21 3.71 0.99 0.21 5.57 68.1 265-64 64.04 1.28 13.12 8.27 1.60 4.45 3.46 0.97 0.20 0.24 1.73 0 6.19 76.5 0.25 1.79 0.023 265-65 63.79 1.26 13.25 8.14 1.34 4.21 3.84 0.97 0.22 5.86 67.5 265-66 62.81 1.43 13.13 9.05 1.99 4.98 3.86 0.89 0.24 0.23 1.81 0 5.84 84.2 0.26 2.04 0.022 265-42 66.92 0.94 13.09 8.04 0.86 3.49 0.83 1.17 0.21 0.51 5.87 82.9 0.44 265-67 64.10 1.34 13.33 8.49 1.49 4.41 3.93 0.95 0.23 5.92 68.0 265-70 66.26 0.87 13.20 7.17 0.80 3.23 4.08 1.33 0.19 0.70 2.35 0 6.08 88.1 0.53 1.76 0.027 265-83 67.46 0.76 13.27 6.68 0.67 2.98 3.88 1.37 0.16 0.67 1.53 0 5.94 87.2 0.49 1.12 0.018 265-84 64.39 1.13 13.17 8.18 1.23 3.92 3.41 1.19 0.22 5.73 77.8 265-85 65.01 1.06 13.13 7.99 1.18 3.78 3.67 1.22 0.20 0.64 1.90 0 5.95 82.5 0.52 1.55 0.023 265-94 65.22 0.97 13.04 7.90 1.13 3.54 4.29 1.14 0.23 5.90 83.9 265-95 67.46 0.77 13.10 6.47 0.94 3.01 4.43 1.21 0.15 6.07 83.6 266-53 64.28 1.06 13.31 8.06 1.12 3.73 4.16 1.09 0.25 0.66 1.74 0 5.38 82.2 0.60 1.60 0.021 266-57 62.47 1.30 13.16 9.20 1.59 4.37 4.11 0.98 0.29 0.51 2.08 0 5.65 72.4 0.52 2.13 0.029 264-09 65.97 0.90 13.19 7.02 1.05 3.47 4.32 1.20 0.21 0.58 5.73 83.9 0.48
96
Figure 3-1. Bathymetric map of the East Pacific Rise showing the location of the 9°N OSC, the Clipperton and Siqueiros transform faults.
97
Figure 3-2. H2O (wt%), Cl (wt%), and CO2 (ppm) versus MgO (wt%) for glasses from the 9°N OSC. Black crosses indicate the calculated fractional crystallization trend using MELTS (Ghiorso and Sack, 1995). All dacitic lavas and several basaltic andesites lie above the calculated trend, indicating another processes is involved in their petrogenesis.
A B
C D
98
Figure 3-3. H2O/Ce and Cl/K2O ratios versus MgO (wt%) for glasses from the 9°N OSC. Generalized trends for assimilation and fractional crystallization are shown as dashed lines. In general, the andesites, dacites and most of the basaltic andesites have higher incompatible element ratios than the basaltic lavas, suggesting that they cannot result from fractional crystallization alone. Instead, they are consistent with assimilation of an altered basalt.
A
B
99
Figure 3-4. Cl/K2O versus H2O/K2O for glasses from the 9°N OSC. Lines representing mixing of 6 possible assimilants with an OSC basalt are shown. Mixing end-members include 5 and 10% partial melts of an altered basalt, an altered basalt, a 50% and 15% NaCl brine, and seawater (see Kent et al., 1999 for references). See text for mixing end-member concentrations. A combination of partial melting of an altered basalt and <0.2 wt% of a 50% NaCl brine with a basaltic end-member can explain the formation of high-silica lavas on the OSC.
100
Figure 3-5. δ18O versus MgO for glasses from the 9°N OSC. The MOR dacites, andesites and several basaltic andesites lie below calculated fractional crystallization trends (black dashed line). The lower δ18O values are consistent with assimilation of altered oceanic crust, which has lower δ18O values due to high-temperature hydrothermal circulation. δ18O values for lavas from the Juan de Fuca ridge and Galapagos Spreading Center are shown for comparison.
101
Figure 3-6. H2O versus CO2 for 9°N OSC glasses. Superimposed on this diagram are CO2–H2O vapor saturation curves for 200 to 800 bars based on models by Dixon et al. (1995a, b). Black dashed lines show a general magma degassing trends during ascent. The gray band represents an approximate depth of the top of the imaged melt lens (Kent et al., 2000). The pressure at the seafloor is shown as a dotted line. Most of the basalts are in equilibrium with pressures consistent with the top of the imaged melt lens (~550 bars). The dacites, andesites, and two basaltic andesites have completely degassed CO2 and may have also lost H2O prior to or during eruption. H2O and CO2 concentrations in one basaltic andesite lie between the high-silica lavas and the basaltic lavas, which may indicate mixing (red dashed line).
102
CHAPTER 4 CRUSTAL DIFFERENTIAION AND SOURCE VARIATIONS AT THE 9°N
OVERLAPPING SPREADING CENTER; EAST PACIFIC RISE
Introduction
Mid-ocean ridges (MOR) can be divided into a series of segments or
discontinuities that range in length from tens of meters to hundreds of kilometers
(Sempere & Macdonald, 1986; Macdonald et al., 1988). Overlapping spreading centers
(OSC) are second-order discontinuities that form between widely spaced first-order
transform faults on fast to intermediate spreading ridges (Macdonald & Fox, 1983;
Sempere & Macdonald, 1986; Carbotte & Macdonald, 1992). Both first and second
order discontinuities delineate physical and geochemical segmentation of the ridge that
reflect sub-ridge processes, such as variations in degrees of mantle melting and/or
separate crustal magma reservoirs (e.g. Macdonald et al., 1988).
Lavas erupted along fast to intermediate spreading centers, such as the northern
East Pacific Rise (EPR), may produce a range of basaltic lavas (e.g. Batiza & Niu,
1992), but they only rarely erupt compositions with MgO concentrations <5 wt%. This
relatively limited compositional diversity compared to other tectonic settings is
commonly attributed to shallow-level fractional crystallization of primitive magmas within
an axial magma chamber that is buffered by relatively frequent recharge with more
primitive mantle melts (Klein, 2005). Additionally, geochemical variations in mid-ocean
ridge basalt (MORB) may result from variable mantle melting parameters and/or mantle
sources (Klein & Langmuir, 1987; Langmuir et al., 1992).
Lavas erupted at ridge segment ends, such as OSC's, can have a broad range of
compositions compared to the relatively limited basaltic compositions erupted from
magmatically robust segment centers (Christie & Sinton, 1981; Langmuir et al., 1986;
103
Wanless et al., accepted). This variability can be attributed to lower magma supply and
cooler crust at the end of ridge segments (cold edge effect), which cause increased
magmatic fractionation prior to eruption (Christie & Sinton, 1981; Perfit et al., 1983;
Sinton et al., 1983; Perfit & Chadwick, 1998; Rubin & Sinton, 2007). Although crystal
fractionation is undoubtedly a primary process in magma differentiation at MOR, recent
geochemical studies show that highly evolved incompatible trace element
concentrations (Wanless et al., accepted) and low oxygen isotope ratios (Wanless et al.,
submitted) in MOR dacites require partial melting and assimilation of oceanic crust. The
extent to which these processes contribute to the chemistry of more mafic magmas on
MOR remains poorly constrained.
Lavas erupted at segment ends also preserve geochemical signatures that can be
ascribed to mantle source variations. A greater proportion of enriched mid-ocean ridge
basalt (E-MORB) erupted at ridge segment ends compared to segment centers on the
northern East Pacific Rise (EPR) may represent decrease in the amount of melt feeding
the ridge axis in these regions (Christie & Sinton, 1981; Sinton et al., 1983; Langmuir et
al., 1986). Despite evidence for variations in mantle source between segments, the
spatial distribution of E-MORB lavas at OSC is not well constrained.
The present study combines geophysical, geochemical, and bathymetric data to
examine the magmatic plumbing system at 9ºN OSC on the EPR (Figure 4-1). We use
major and trace element data and isotopic ratios to explore the relative roles of crystal
fractionation, assimilation and magma mixing within the shallow crust beneath the 9°N
OSC. Trace element data and isotopic ratios are used to examine how variations in
mantle sources and magma supply can contribute to the distribution of compositions
104
observed on both limbs of the OSC. These analyses are compared to compositions of
lavas erupted from segment centers to the north (9-10°N) and south (8°37’N) of the
OSC to explore the extent of these variations beneath the EPR.
Background, Tectonic Setting and Geology of the 9oN OSC
Overlapping Spreading Centers
At OSCs the ridge axis splits into two overlapping, curvilinear, sub-parallel axes or
“limbs” that may offset the ridge by up to 15 km (Macdonald & Fox, 1983; Macdonald et
al., 1988). The ratio of offset width to overlap length is approximately 1:3 (Macdonald et
al., 1988) and the inward curving limbs surround an elongate basin (Macdonald & Fox,
1983). The limbs migrate sub-parallel to the overall ridge strike with one limb
propagating and the other dying (Hey et al., 1980; Sinton et al., 1983; Pollard & Sydin,
1984). Consequently, the propagating limb migrates into older and colder ocean crust
and the receding limb gradually becomes amagmatic. As the OSC migrates with time, it
produces offsets in bathymetry and magnetic signature of the ocean crust (Hey et al.,
1977; Carbotte & Macdonald, 1992). Average OSC migration rates can be calculated
using the fossil V-shaped bathymetric scars developed in the wake of the propagating
OSC that are left either by linking of one ridge axis with the other during propagation,
which leads to decapitation of the ridge tip, or by repeated self-decapitation along a
single limb of the OSC (Macdonald et al., 1988).
Tectonic Setting and Previous Studies of 9°N OSC
The 9°N OSC is located between the Clipperton and Siqueiros transform faults
(Figure 4-1) and is one of eight 2nd order discontinuities on the northern EPR
(Macdonald & Fox, 1983). It consists of two north-south trending ridges that overlap by
~27 km and offset the ridge by ~8 km (Sempere & Macdonald, 1986). The 9°N OSC is
105
divided into three main sections: the eastern propagating limb, the western receding
limb, and an overlap basin separating the two limbs (Figure 4-2). The eastern limb can
be further divided into the east limb ridge, the east limb tip and the northern inter-limb
region, also called the northwest flank (Figure 4-2), following nomenclature in Nunnery
et al., (2008). Based on geologic, magnetic and bathymetric data, the eastern limb has
been propagating south at a rate of ~42 km/Myr, the western limb has receded
(Macdonald & Fox, 1988; Carbotte & Macdonald, 1992).
The 9°N OSC is one of the largest and most extensively studied 2nd order
discontinuities on the global MOR system. It has been the focus of several geophysical
studies (Detrick et al., 1987; Harding et al., 1993; Kent et al., 1993; Kent et al., 2000;
Bazin et al., 2001; Dunn et al., 2001; Tong et al., 2002), including the first multi-channel
seismic 3-D survey of a MOR (Kent et al., 2000) and a 3-D seismic refraction
study(Dunn et al., 2001). Collectively, these studies reveal the presence of a shallow
melt lens beneath each of the limbs and a widening of the eastern lens below the inter-
limb region north of the overlap basin (Figure 4-3; Kent et al., 2000). The melt lens
beneath the western, receding limb is narrow and shows no discernable variation in
depth along axis, but the melt lens beneath the eastern, propagating limb is variable in
width and depth (Kent et al., 2000). Beneath the southern portion of the east limb the
melt lens is narrower and deeper than the rest of the eastern ridge axis, plunging ~500
m over ~6 km (Kent et al., 2000). It also cuts across the tectonic seafloor fabric (White
et al., 2009). To the north, the melt lens widens to more than 4 km (Figure 4-3) and is
displaced slightly off axis to the west into the inter-limb region (Kent et al., 2000; Tong et
al., 2002). Tomographic studies reveal a low velocity zone beneath the entire OSC at ~
106
9 km depth near the mantle-crust transition (Toomey et al., 2007). This zone extends
~8 km to either side of the ridge axis (Dunn et al., 2001).
Recently, the 9°N OSC was the focus of the MEDUSA2007 research cruise
(AT15-17), which carried out detailed mapping and extensive sampling of the region
using the ROV Jason2, DSL-120A side-scan system, and the WHOI TowCam (Fornari,
2003). This expedition acquired >10,000 photographs of the seafloor in combination
with the most complete and well-constrained lava sampling of any OSC (White et al.,
2009; Wanless et al., accepted). Prior to this study, limited lava sampling of the region
during the CHEPR cruise indicated that both high-silica and E-MORB lavas existed in
addition to N-MORB lavas (Langmuir et al., 1986).
9°N OSC Geology
The northern portion of the eastern ridge is characterized by an axial summit
trough (AST) ~0.9 km across and ~50 m high (Figure 4-4). The AST walls gradually
diminish in height to the south, eventually disappearing by ~9°05’ N. This topographic
change is accompanied by a gradual shift from a volcanically dominated seafloor fabric
in the north to a highly faulted and tectonized fabric observed at the southern tip that
also correlates with narrowing of the imaged melt lens, from ~4 km wide to <1 km
(Figure 4-3), and a near-absence of high-silica lavas (Figure 4-2).
Video, still photographs, and side scan collected during the 2007 cruise suggest
that over 80% of the lavas erupted within the mapped region are pillow lavas (White et
al., 2009), a morphology that is observed across all regions of the OSC, but dominates
the seafloor in the inter-limb region and overlap basin. In comparison, >80% of lavas
erupted on the EPR north of the OSC are lobate and sheet flows, and <20% are pillows
(Kurras et al., 2000; White et al., 2002; Soule et al., 2005; Fundis et al., 2010). Pillow
107
flow morphology at the OSC appears to be controlled by low effusion rates (<0.1 m3/s
at a viscosity of 100 Pa s; White et al., 2009). The youngest lavas on the northern
portion of the eastern limb, based on glassy surfaces, lack of sediment cover, and well
developed ornamentation, are confined to the neo-volcanic zone (Nunnery et al., 2008),
a narrow region of focused magmatism where zero-age lavas erupt (Perfit & Chadwick,
1998). On the southern tip of the eastern limb, the neo-volcanic zone is ill-defined;
however, the youngest lavas are primarily located along the western margin of the
eastern limb (Nunnery et al., 2008). We now describe the geology of each region of the
OSC in detail.
Despite the overwhelming abundance of pillow lavas formed at the OSC, lavas
erupted on the east limb ridge are morphologically quite diverse (sheet, hackly and
lobate flows, and small and large pillow lavas) and appear to correlate, at least to a first
degree, with composition. High-silica lavas erupted within the AST form a ~10 m high,
linear, pillow mound composed of atypically large pillow lavas. This pillow mound is
surrounded primarily by lobate lavas and to a lesser degree, sheet flows. Several large
areas of collapse, with drainback features, were observed within the AST and are
associated with basaltic sheet flows that surround the high-silica pillow mounds. The
lobate and sheet lavas are predominantly ferrobasaltic in composition; however, several
FeTi basalts were also recovered within the AST. Basaltic andesites also erupted within
the AST and form lobate flows and pillow mounds. Lavas sampled from both sides of
the AST walls at the edges of the ridge axis are primarily andesites and dacites and
form large, elongate to bulbous pillows. The only active hydrothermal vent at the OSC
was observed within the AST on the east limb axis within an andesitic to dacitic pillow
108
mound. A faulted and tectonized fabric dominates the seafloor on the southern tip of
the eastern limb. In places it is covered by elongate fresh pillow mounds that are cut by
fabric-parallel fissures. Lobate flows are sparse in this region.
Lavas erupted over the northern portion of the overlap basin and inter-limb region
primarily consist of pillow lavas. A linear pillow mound, trending approximately N-S,
defines the outer edge of this region and lies over the westernmost extents of the wide
melt lens (Kent et al., 2000). Lavas comprising this mound are younger than expected
(based on thin sediment cover and the presence of glassy buds) for their distance from
the neo-volcanic zone, suggesting that this region is the site of off-axis volcanism
(Nunnery et al., 2008; White et al., 2009). This conclusion is supported by excess 230Th
measured in several of these samples, which indicate eruption ages of <8,000 ka
(Waters, pers comm.). The southern overlap basin is primarily composed of bulbous
pillow mound fields, based on backscatter images (White et al., 2009), but photographic
and sample coverage is sparse.
The western limb of the OSC differs from the eastern limb in being primarily
comprised of sheet and lobate flows, with fewer pillow lavas (White et al., 2009) and
extensive areas of lava collapse and pillars with drain back features. Similar features
are common within AST’s on many other regions of the northern EPR. Extinct vent
fields were also observed, but no active hydrothermal venting.
8°37’ N EPR Deval
The 8°37’N deviation in axial linearity (deval) is located south of the 9°N OSC and
approximately 40 km north of the Siqueiros transform fault (Figure 4-1; Langmuir, 1986).
It is the southern extension of the western limb of the 9°N OSC. The 8°37’N deval is
structurally similar to a small OSC and, therefore, may have a similar magmatic
109
plumbing system. This deval was dredged during the CHEPR cruise, which recovered
several E-MORB lavas and a high-silica andesite (Langmuir et al., 1986). Additional
investigations of the area were conducted during a response cruise in 2003 (Zierenberg,
2007 pers comm.) to determine if seismic events recorded on hydrophones in March of
2001 were indicators of new volcanic activity (Bohnenstiehl et al., 2003). This cruise
included several ALVIN dives; however, there was no evidence of a recent eruption.
Lava morphology in the region includes sheet flows, lobates and pillow lavas. During
these dives nine samples were recovered, including several glassy E-MORBS and one
glassy andesite. Data from these samples are discussed below.
Geochemical Methods
Over 280 rock samples were collected from the 9°N OSC during the
MEDUSA2007 cruise. Of these, 275 have glassy outer rims, from which glass was
handpicked and analyzed on a JOEL 8900 electron microprobe for major and minor
element concentrations at the USGS facility in Denver, CO. Eight to ten points were
analyzed per sample. The probe diameter was routinely 20 μm, to minimize sodium
loss, with an accelerating voltage of 15 keV and a beam current of 20nA. Several
USGS minerals were used as calibration standards and secondary normalizations
involved the JdF-D2 glass “standard” (Reynolds, 1995), ALV 2392-9 (in-house
standard), and dacite glass GSC (USGS standard) to account for instrumental drift
(Smith et al., 2001). Chlorine and sulfur concentrations, as well as high-precision
potassium values were also determined on nine samples, using 200-second peak/100
second background counting times. Major element concentrations for samples from the
8°37’N deval on the EPR were determined by microprobe at UC Davis following
methods described in Schiffman et al., (2010).
110
A representative subset of fresh glasses were handpicked, cleaned in a dilute
acid, and dissolved for trace element and isotope analyses following methods described
in Goss et al., (2010). 73 samples from the 9°N OSC and seven samples from the 8°37
N deval were analyzed for high precision trace elements on an Element2 Inductively
Coupled Plasma Mass Spectrometer (ICP-MS) at the University of Florida. External
calibration was done to quantify results using a combination of internal (ENDV –
Endeavour and ALV 2392-9) and USGS (AGV-1, BIR-1, BHVO-1, BCR-2 and STM-1)
rock standards. High precision Pb, Sr, and Nd isotopic abundances on 37 samples were
determined using the Nu- Plasma multi-collector ICP-MS at the University of Florida
(Wanless et al., accepted). For detailed descriptions of sample preparation, dissolution
procedures, standards, and statistical data, see Goss et al., (2010) and Kamenov et al.,
(2007).
Geochemical Results
Lavas erupted at the 9°N OSC display a large range of compositions on both the
east and west limb (Figure 4-2). Below, we discuss the geochemical variability on each
limb by rock type to better address the petrogenesis of the OSC lavas.
Geochemistry of East Limb Lavas
The east limb of the 9°N OSC has produced basalts, ferrobasalts, FeTi basalts,
basaltic andesites, low-P2O5 and high-P2O5 andesites, and dacites (Figures 4-2, 4-5).
These lavas cover a wide compositional range and we, therefore, subdivide our results
below by rock type. Major and trace element data from the east limb are present in
Table 4-1 and 4-2, respectively. Radiogenic isotope ratios for the east limb lavas are
presented in Table 2-2.
111
Basalts
Basalts consists of all lavas with <52 wt% SiO2 and include ferrobasalts and FeTi
basalts. In comparison to more mafic basalts from the heavily sampled 9°50 bulls-eye
site on the northern EPR, basalts from the east limb of the OSC are more evolved on
average (Table 4-1), with MgO concentrations ranging from 6.16 to 7.36 wt%, high FeO
(10.83 to 13.50 wt%), and TiO2 (1.70 to 2.83 wt%) contents (Figures 4-5, 4-6). All lavas
are N-type MORB with variable K2O (0.11 to 0.27 wt%), P2O5 (0.12 to 0.27 wt%) and Cl
concentrations (0.004 to 0.05 wt%). P2O5/TiO2 ratios range from 0.08 to 0.13 (Figure 4-
7) and Cl/K2O ratios from 0.03 to 0.29 and are thus comparable to other MOR
ferrobasalts (e.g. Michael & Cornell, 1998). Incompatible trace element concentrations
are relatively high compared to MORB from the northern EPR but ratios are relatively
constant (Figures 4-8, 4-9) despite eruption over wide geographic region at the OSC
and are comparable to other ferrobasalts erupted on MOR. Rare earth elements (REE)
patterns are remarkably similar in the ferrobasalts, with an average LaN/YbN ratio of 0.77
± 0.05 (Figure 4-9). High field strength element (HFSE) ratios are also relatively limited
with Zr/Nb of 40 ± 2.8, and U/Nb of 0.03 ± 0.001. Compatible trace element
concentrations in the ferrobasalts are variable (Cr = 9 to 166 ppm; Ni = 32 to 71 ppm),
but show positive correlations with MgO.
Basaltic andesites/low-P2O5 andesites
Basaltic andesites (SiO2 >52 and <57 wt%SiO2) and low-P2O5 andesites from the
east limb have highly variable major and trace element compositions (Table 4-1). MgO
in the basaltic andesites ranges from 1.5 to 6.5 wt% and FeO ranges from 8.27 to 13.64
wt% (Figure 4-5). Minor elements are also highly variable (Figure 4-6; P2O5 = 0.18 to
0.58 wt%; K2O = 0.19 to 0.99 wt%; and Cl = 0.04 to 0.5 wt%). P2O5/TiO2 (average of
112
0.15 ± 0.04) and P2O5/K2O (average of 0.79 ± 0.28) ratios span the range between the
high-silica and ferrobasalt end-members (Figure 4-7). Ni and Cr concentrations ranging
from 6 to 50 ppm and 5 to 117 ppm, respectively (Figure 4-8) are generally, though not
exclusively, lower than in the ferrobasalts. REE and HFSE concentrations are variable,
but have relatively constant ratios with average LaN/YbN ratios = 1.08 ± 0.12, Zr/Nb = 47
± 8.3, and U/Nb ratios ≥0.03 (Figure 4 -9). The basaltic andesites lie along well-defined
trends toward high-silica lavas.
High-P2O5 andesites
Low MgO and high FeO, TiO2, P2O5, and K2O characterize the high-P2O5
andesites and distinguish them from low-P2O5 andesites discussed above. P2O5 ranges
from 0.54 to 0.78 wt%, which is up to 0.49 wt% greater than the average low-P2O5
andesite (Figure 4-6). Consequently, these lavas have higher P2O5/TiO2 (0.29 to 0.39)
and P2O5/K2O (0.86 to 1.47) ratios compared to other lavas on the east limb (Figure 4-
7). They have high average Cl concentrations of 0.3 ppm (Figure 4-6) compared to
MOR basalts. Most incompatible trace element concentrations are high in these lavas
compared to ferrobasalts and basaltic andesites, but similar to the low-P2O5 andesites
(Figure 4-8). The high-P2O5 andesites have an average Zr ~ 700 ppm, La ~22 ppm, and
Yb ~15 ppm with LaN/YbN ratios of 1.15 ± 0.11, and Zr/Nb ratios of 47 ± 3.5, (Figure 4-
9). Distinctive features are the lack of negative Nb and Ta anomalies and slightly lower
U/Nb ratios (<0.03) when compared to the low-P2O5 andesites and some basaltic
andesites.
Dacites
The geochemistry of the MOR dacites is discussed in detail in Wanless et al.,
[accepted] and are, therefore, only briefly discussed here. They have SiO2
113
concentrations up to 67.5 wt% and MgO as low as 0.67 wt%. P2O5 concentrations range
from 0.15 to 0.25 wt% (Figure 4-6). K2O and Cl concentrations average 1.17 and 0.63
wt% respectively. P2O5/TiO2 and P2O5/K2O are also low (Figure 4-7). Highly
incompatible trace element concentrations are elevated, with average Ba, U, and La
concentrations of 62 ppm, 0.8 ppm and 28 ppm respectively (Figure 4-8). They also
have high Zr (734 ppm to 1050 ppm) and Hf (18.6 ppm to 24.9 ppm) concentrations, but
low Nb (13 ppm to 16.5 ppm), Ta (0.8 ppm to 1.4 ppm) and U/Nb ratios >0.03 (Figure 4-
9). The dacites have average LaN/YbN ratios of 1.26 ± 0.10, Zr/Nb = 58 ± 4.5, and U/Nb
ratios ≥0.04.
Isotopic compositions of East Limb lavas
In contrast to the variability noted in the major and trace element compositions of
east limb, the radiogenic isotopic ratios are relatively uniform. They have 208Pb/204Pb
and 206Pb/204Pb ranging from 37.642 to 37.699 and 18.235 to 18.294 respectively
(Figure 4-10a). 87Sr/86Sr ranges from 0.70243 to 0.70258 and 143Nd/144Nd ratios range
from 0.51314 to 0.513120 (Figure 4-10b), which is similar to lavas erupted on the EPR
to the north near 9°50’N (Sims et al., 2002; Sims et al., 2003; Wanless et al., accepted).
West Limb Lavas
Fifty lava samples (Table 4-3, 4-4) were collected from the west limb of the 9°N
OSC of which 47 are basaltic, 2 are basaltic andesites and one is andesitic in
composition (Figure 4-11a). Fourty-five samples have typical N-MORB K/Ti ratios (K/Ti
= K2O/TiO2*100 <1.3; Langmuir et al., 1986; Sinton et al., 1991; Reynolds et al., 1992;
Perfit et al., 1994) and five have higher ratios typical of E-MORB (K/Ti>0.13; Figure 4-
11b). Compared to the east limb basalts, the average west limb basalt is slightly more
primitive, with MgO concentrations ranging from 6.42 to 8.72 wt% (Figure 4-11a) and
114
K2O and P2O5 concentrations from 0.07 to 0.17 wt% and 0.13 to 0.26 wt%, respectively
(Figure 4-11a). Lavas dredged in this region during the CHEPR cruise had similar
compositional variations, including E- and N-MORB (Langmuir et al., 1986). West limb
N-MORB lavas have LaN/YbN ratios of 0.72 ± 0.04 and Zr/Nb of 37 to 45 (Figure 4-12).
Highly incompatible elements and HFSE have limited ranges in concentration, for
instance, Ba ranges from 3.4 ppm to 9.2 ppm, U from 0.04 ppm to 0.07 ppm, and Nb
from 1.55 ppm to 3.22 ppm (Figure 4-12). The west limb E-MORB lava has LaN/YbN
ratios of 2.8 and Zr/Nb ratios of 9.7. The single andesite has major element
concentrations (Figure 4-11a) and trace element abundances (Figure 4-13) similar to
the high-P2O5 andesites on the east limb (Figure 4-11a). It has a SiO2 concentration of
55.95 wt% and MgO of 2.84 wt%, with high P2O5 (0.73 wt%) and K2O (0.67 wt%).
Isotopically, the west limb lavas have slightly more radiogenic values compared to
the east limb lavas (Figure 4-10; Table 4-5). The N-MORB lavas have 208Pb/204Pb and
206Pb/204Pb ratios ranging from 37.737 to 37.846 and 18.253 to 18.369, while the E-
MORB lava has higher lead isotope values (208Pb/204Pb = 38.022 and 206Pb/204Pb =
18.590) (Figure 4-10a). N-MORB 87Sr/86Sr values ranges from 0.70249 to 0.70265 with
an average of 0.70256 (Figure 4-10b) whereas the E-MORB 87Sr/86Sr ratio is also more
radiogenic (0.70282). 143Nd/144Nd ratios of the N-MORB lava range from 0.51314 to
0.51315, while the E-MORB lava has a ratio of 0.51305. The andesite has isotopic
ratios intermediate between the west limb E-MORB and N-MORB values.
8°37’ EPR Lavas
Nine samples collected from the 8°37 N deval range in composition from basalt to
andesite, with MgO and SiO2 concentrations from 7.29 to 3.08 wt % and 50.47 to 57.14
wt%, respectively (Table 4-6). P2O5 concentrations range from 0.3 to 0.5 wt.%. Two of
115
the samples are E-MORB with K/Ti values of 0.19 and 0.15. REE and HFSE
abundances are variable with average LaN/YbN ratios of 1.33 ± 0.17, Zr/Nb of 20.92 to
31.84, and U/Nb of 0.02 to 0.04.
Radiogenic isotope ratios were measured in seven samples collected from 8°37’N
(Table 4-7). Compared to basalts collected from the 9°50N region of the EPR (Sims et
al., 2002; Sims et al., 2003) and the 9°N OSC, these samples have, on average, more
radiogenic Pb (Figure 4-10a) and 87Sr/86Sr ratios (0.70255 to 0.70267) and less
radiogenic 143Nd/144Nd ratios (0.51312 to 0.51318; Figure 4-10b). Consistent with their
elevated incompatible element ratios, the samples have Sr, Nd and Pb isotopes that
plot between the field typical EPR N-MORB and more radiogenic E-MORB (Figure 4-
10).
Discussion
Shallow-level Processes Involved in the Petrogenesis of Ferrobasalts, FeTi Basalts and Basaltic Andesites at the 9°N OSC
Shallow level differentiation of MOR magma is primarily controlled by fractional
crystallization and mixing of different MORB melts (Clague & Bunch, 1976; Bryan &
Moore, 1977; Byerly, 1980). These processes are often difficult to discern from each
other on geochemical variation diagrams because mixing can produce compositions
that lie along liquid lines of descent resulting from fractional crystallization. However, our
geochemical results and petrographic studies suggest that both mixing and fractional
crystallization are important at the OSC, particularly in the generation of basaltic
andesites.
The geochemistry of most of the basaltic lavas erupted at the 9°N OSC, including,
ferrobasalts, FeTi basalts, and many basaltic andesites, can primarily be explained
116
through low-pressure fractional crystallization of typical N-MORB parental magmas
(Figure 4-5). In general, the compositions of the 9°N lavas are slightly more evolved
than the average MORB erupted along the northern EPR (Perfit et al., 1994; Goss et al.,
2010). Based on MELTS calculations (Ghiorso & Sack, 1995), ferrobasalts and FeTi
basalts can be generated by ~30% and ~55% fractional crystallization (respectively) of
ol + plag + cpx at pressures of 1 kbar from a relatively primitive MORB magma (2392-9
from the 1991 EPR eruption at 9o50’N; Figure 4-6). These results are consistent with
previous studies of EPR lavas covering large geographic areas that collectively suggest
fractional crystallization is the dominant or exclusive process effecting the composition
of MOR magmas (e.g. Batiza & Niu, 1992). In contrast, many of the basaltic andesites
erupted at the 9°N OSC lie off of typical fractional crystallization trends.
OSC lavas are predominantly aphyric, but some lavas contain sparse phenocrysts
and microphenocrysts of olivine, clinopyroxene and plagioclase that are consistent with
shallow- level fractional crystallization of a typical MORB (Grove et al., 1992).
Petrographic studies, however, suggest some lavas have had more complex
petrogenetic histories (Zaino, 2009). Although some plagioclase phenocrysts in east
limb ferrobasalts and basaltic andesites exhibit normal zoning from core to rim, others
show reverse zoning or no zoning at all. Commonly, all three zoning patterns occur in a
single sample. This suggests that each phenocryst has undergone a distinct magmatic
history prior eruption on the seafloor and supports the role of mixing in the petrogenesis
of these lavas. Resorbed rims on olivine phenocrysts in east limb basalts coupled with
more magnesian compositions that should be in equilibrium with co-existing glasses are
also indicative of magma mixing (Zaino, 2009).
117
Many of the low MgO OSC lavas have elemental concentrations that lie off of
calculated liquid lines of descent and are inconsistent with crystal fractionation alone.
This is particularly apparent in the concentrations of FeO, TiO2, and P2O5 in many
basaltic andesites (Figure 4-6). These elements are typically enriched in MORB
systems during fractional crystallization until melts become saturated with Fe-oxides (~5
wt % MgO) and apatite (~1 wt% MgO). However, many of the evolved OSC lavas have
relatively low concentrations of these elements and appear to lie along straight lines that
suggest mixing between high-silica and ferrobasaltic magmas. Similar linear trends are
observed in incompatible and compatible trace element concentrations (i.e. Cr and Ni)
versus Zr and MgO (not shown). Similarly, trace element ratios (e.g. U/Nb) of basaltic
andesites do not plot along fractional crystallization trends but lie along mixing lines
between evolved basalts and high-silica lavas (Figures 4-8, 4-9). The high-silica mixing
end-member is likely produced during partial melting of the oceanic crust or assimilation
and fractional crystallization (AFC) processes that can produce MOR dacites (Wanless
et al., accepted). Using an OSC dacite composition as an end-member, our mass-
balance calculations confirm that many of the basaltic andesites can be produced by
bulk mixing of 25% dacitic melt with 75% ferrobasaltic melt. Similar processes can
explain the low-P2O5 andesites (see discussion below).
This scenario is consistent with the petrogenetic models of dacite formation on
OSC, which requires an episodic magma supply at the propagating limb of the OSC
(Wanless et al., accepted). In this case, a basaltic magma is injected into the OSC and
undergoes variable degrees of crystal fractionation to produce ferro- and FeTi basalt
compositions. These magmas mix to varying degrees with pre-existing high-silica melts
118
beneath the ridge axis to produce a range of compositions, including basaltic andesites
and andesites (Figure 4-5).
Fractional crystallization versus assimilation in the petrogenesis of andesites
Although fractional crystallization and magma mixing appear to be the dominant
processes involved in the differentiation of magmas at the 9°N OSC, incompatible trace
element concentrations suggest that the formation of highly evolved lavas (i.e., dacites)
on ridges requires partial melting and assimilation of altered ocean crust (Wanless et
al., accepted). Here, we assess the roles of fractional crystallization, assimilation and
magma mixing in the formation of intermediate compositions (andesites) on MOR.
There are two distinct geochemical populations of andesites at the OSC, which likely
require different petrogenetic histories (Figure 4-5).
Both populations of andesites (high-P2O5 and low-P2O5 andesites) have similar
SiO2 and MgO concentrations, but different TiO2, FeO, Al2O3, P2O5, and trace element
concentrations at a given MgO (Figures 4-5, 4-6, and 4-8). Petrologic modeling of
major element variations indicates that high-P2O5 andesite magmas formed primarily
through extensive crystal fractionation of basaltic magmas (Figure 4-6). MELTS
calculations (Ghiorso and Sack, 1995) suggest that this process involved up to 75%
fractional crystallization of ol + plag + cpx and minor amounts of Fe-oxide of a
ferrobasaltic parent (265-43) at 1kbar pressure. Oxygen fugacity at the QFM-1 buffer is
required to delay the onset of Fe-oxide crystallization, and produce the elevated FeO
and TiO2 concentrations observed (Figure 4-5). Elevated P2O5 concentrations require
that apatite crystallization has not occurred during the evolution of the melt (Figure 4-6).
This is consistent with MELTS calculations and is supported by apatite saturation
calculations for these compositions (Watson, 1979), which suggest the high liquidus
119
temperature (>1000°C) of these melt compositions inhibits apatite saturation. Trace
element ratios also suggest extensive fractional crystallization is required to explain
incompatible trace element concentrations and ratios (e.g. low U/Nb ratios). Rayleigh
fractionation calculations (Figure 4-8) suggest that ~85% fractional crystallization is
required to produce the high-P2O5 incompatible trace element abundances, compared
to the 75% calculated using MELTS for the major elements (Figure 4-6).
In contrast, low-P2O5 andesites have low FeO, TiO2, and P2O5 concentrations and
high U/Nb ratios (Figures 4-6, 4-9). The low FeO and TiO2 and high U/Th ratios are a
result of crystal fractionation of Fe-oxides and the low P2O5 is a results of apatite
crystallization. These lavas are spatially related to high-silica dacites erupted on-axis
and are likely formed from mixing of high-silica melts and ferrobasaltic magmas. The
high-silica melts are a result of AFC processes that include crystallization of Fe-oxides
and apatite (Wanless et al., accepted). This is similar to the petrogenesis of the basaltic
andesites, but in opposite proportions. Bulk mixing calculations suggest mixes of up to
25% ferrobasalt and as low as 75% dacitic melt will produce compositions similar to the
low-P2O5 andesites.
Formation of west limb lavas by fractional crystallization
The west limb of the OSC has been receding at a rate approximately equal to the
propagation of the east limb (Carbotte & Macdonald, 1992). This must reflect a
progressive decrease in magma supply on the western limb tip and, therefore, a
decrease in the heat supply and magma recharge, which could lead to greater degrees
of crystal fractionation (e.g. Christie & Sinton, 1981). Under these circumstances, one
might expect the eruption of a range of compositions, including high-silica lavas.
120
However, the lavas erupted on the west limb encompass a narrower compositional
range than that observed on the eastern, propagating limb. With the exception of a
single andesitic lava, lavas erupted on the west limb can be explained by up to ~25%
fractional crystallization of either an E- or N-MORB parent with ~8.5 wt% MgO (Figure
4-11). Trace element patterns are consistent with fractional crystallization as the
primary shallow-level differentiation process (Figure 4-12); however, some complex
zoning patterns in plagioclase phenocrysts (Zaino, 2009) are consistent with a
combination of crystallization and magma mixing.
The single west limb andesite recovered appears to have formed through
extensive fractional crystallization. Petrologic modeling suggest that ~75% fractional
crystallization can explain the major and trace element compositions observed in the
west limb andesite. When compared to the wide range of east limb andesites, the west
limb andesite has trace element patterns similar to the high-P2O5 andesites, which are
geochemically dominated by fractional crystallization (Figure 4-13). The west limb
andesite lacks negative Nb and Ta anomalies observed in the low-P2O5 andesites, has
a small negative K anomaly, and elevated phosphorus contents and P2O5/TiO2 ratios
(Figures 4-11, 4-14). Assimilation does not appear to play a significant role in the
formation of the west limb andesite, which may be due to cooler crustal conditions on
the dying limb.
Where Are the South Tip and West Limb Dacites?
A near-absence of high-silica lavas appears to be a characteristic of both the
southern tip of the eastern limb and the western limb as a whole (Figure 4-2), in contrast
to the abundant dacites erupted over the northern east limb. This lack of high-silica
samples is puzzling because the tips of both the eastern and western limbs should
121
experience low magma supply rates, which would promote greater degrees of fractional
crystallization and more evolved magma compositions (Christie & Sinton, 1981). We
argue that the subtle difference in the thermal conditions in the oceanic crust play a
major role in controlling the composition of erupted lavas.
Large amounts of heat, in part from the latent heat of crystallization and in part
from replenishment by an episodic magma supply, in the volcanically active region of
the propagating eastern limb enhances melting and assimilation of ocean crust, while
allowing for extensive fractional crystallization (Wanless et al., accepted). In contrast,
low magma recharge rates and cooler conditions in the tectonically controlled southern
east limb tip and dying west limb may result in extensive fractional crystallization, but
melting and assimilation are inhibited. The presence of an unusually large and
extensive melt lens below the northern portion of the eastern ridge axis is consistent
with higher crustal temperatures and greater magmatic activity that results in partial
melting of the oceanic crust. Additionally, the northern portion of the east limb may
experience higher rates of magma recharge than the starved eastern tip and the dying
western limb. This leads to the formation of high-silica lavas through AFC processes
(Wanless et al., accepted). In contrast, the narrower, deeper melt lenses in the
southern tip of the east limb and dying west limb (Kent et al., 2000) provide less heat to
the system, resulting in cooler crust, which may inhibit melting, assimilation and the
formation of dacitic lavas. These observations are supported by compositions of the
two andesite lavas erupted on the southern tip and the west limb (Figure 4-13), which
have incompatible trace element patterns similar to andesites produced primarily by
fractional crystallization with little evidence of assimilation. Another possibility is that
122
melting and assimilation may occur in these regions, but that the melts are
volumetrically too small to erupt, perhaps resulting in the formation of small
plagiogranite intrusions and veins within the crust rather than erupted lavas.
Composition of the Melt Lens Beneath the East Limb
Seismic evidence indicates that the shallow crustal melt lens widens to more than
4 km beneath the northern portion of the east limb (Figure 4-3) but it is centered
beneath the inter-limb region and not below the neo-volcanic zone (Kent et al., 2000;
Tong et al., 2002). Tomographic studies reveal a low velocity zone beneath the entire
OSC at ~ 9 km depth, near the mantle-crust transition (Toomey et al., 2007). Despite
this evidence for a regionally robust magmatic system, there is scant evidence for the
eruption of primitive lavas anywhere within the OSC, but, enigmatically, the lavas
erupted have evolved compositions. This suggests that even in areas with large,
imaged melt bodies, magmas that originate within the shallow mantle have complicated
and extended differentiation histories in the crust and are unlikely to erupt in pristine
condition.
It is difficult to relate lavas erupted within the inter-limb region to the currently
imaged melt lens, however, they do appear to be younger than expected, suggesting
that they did not erupt on-axis (Nunnery et al., 2008). Additionally, the observation that
lavas erupted within the inter-limb region (directly above the melt lens) are almost
exclusively composed of ferrobasalts (60 of 67 samples) with very uniform compositions
(Figures 4-6, 4-8) suggests that these reflect the relatively evolved nature of the shallow
melt lens.
The slight variability in major and trace element compositions of the inter-limb
ferrobasalts can be modeled by 20% to 30% fractional crystallization of a MORB
123
parental composition (MgO = 7 wt%). Many basalts erupted within the neo-volcanic
zone on the eastern limb are also ferrobasaltic and have incompatible trace element
concentrations that are similar to inter-limb ferrobasalts, despite the proximity of high-
silica lavas. Assuming that the these lavas are directly related to the current melt lens,
this suggests that the 4 km wide melt lens is primarily composed of ferrobasaltic magma
and that beneath the east limb neo-volcanic zone it has mixed with high-silica melts,
creating a wide range of compositions on the eastern edge of the melt lens.
Presumably a large volume of primitive mantle-derived basalts must have been
fractionally crystallized at crustal and possibly upper mantle depths to result in a large,
evolved melt lens. Crystallization of a primitive magma to form a 4 km wide melt lens
with ferrobasaltic composition would release significant amounts of heat to the
surrounding region. It is the latent heat of crystallization that provides the extra heat to
cause partial melting and assimilation of the crust leading to the formation of high-silica
magmas along the eastern limb of the OSC (Wanless et al., accepted).
E-MORB Distribution at 9°N OSC
Both E-MORB and N-MORB lavas were recovered on the western limb of the OSC
in this study as well as during the CHEPR cruise [Langmuir et al., 1986]. In contrast,
only N-MORB lavas have been recovered from the east limb. Incompatible element
enrichment corresponds with more radiogenic Sr and Pb isotopes and less radiogenic
Nd isotopes in the west limb E-MORB lavas (Figure 4-10).
The sub-ridge upper mantle is generally thought to be relatively incompatible
element-depleted and isotopically homogeneous; however, there are well-documented
cases of small-scale (veined) heterogeneities in the MOR mantle source. Proximal and
roughly coeval eruptions of E-MORB and N-MORB at MOR’s suggest that the sub-ridge
124
mantle varies in elemental and isotopic composition over small spatial/temporal scales
[e.g. Hart et al., 1973; Sun et al., 1975; White and Schilling, 1978]. Formation of E-
MORB magma is often interpreted as a result of overall lower degrees of mantle melting
of a source comprised of a greater compliment of incompatible enriched material
compared to N-MORB magma. More recently, 2 stage models involving “normal”
amounts of melting of enriched components embedded in a largely depleted sub-ridge
mantle have had success in explaining both elemental and isotopic systematics in E-
type MORB [Donnelly et al., 2004]. The enriched component is thought to be
volumetrically smaller than normal mantle and may consist of enriched veins in a
depleted mantle [e.g. Hanson, 1977]. While E-MORB may be an important component
in ridge magmas, it may be diluted or overwhelmed by the N-MORB signature during
more robust magmatic activity. Consequently, eruption of E-MORB compositions have
been linked to diminished magmatic activity along ridge segments [Reynolds and
Langmuir, 1997, Waters et al. in press].
The eruption of E-MORB lavas only on the western limb of the OSC suggests a
magmatic plumbing system that either allows for the preservation of enriched melts from
the mantle on one limb compared to the other or that the mantle sources supplying each
limb are different. The nonsymmetrical distribution of parental MORB types may result
from differences in magma supply to the two limbs. As spreading shifts from one axis to
the other at an OSC, the magma supply from the mantle will diminish on the dying limb
and progressively increase on the propagating limb. This allows for the preservation of
enriched melts on the dying limb in contrast to the magmatically robust propagating
limb.
125
9°N OSC as a Division in Mantle Components
N-MORB lavas from the western limb of the OSC and at 8°37’N have measurably
different isotope ratios compared to N-MORB lavas erupted on the east limb (Figure 4-
10). The east and west limb N-MORB lavas have similar incompatible trace element
ratios (i.e. LaN/YbN) but west limb N-MORB lavas have higher 208Pb/204Pb and
207Pb/204Pb ratios (Figure 10a) and lower εNd (Figure 10b) compared to east limb N-
MORB. These compositions cannot be explained by simple 2 component mixing of an
enriched source with a typical N-MORB component from the northern EPR (Figure 4-
10). These isotopic differences suggest that different mantle sources are feeding the
two limbs and that the OSC acts as a division between these sources.
The west limb E-MORB lavas have isotopic compositions similar to E-MORB lavas
erupted at the intersection of the EPR with the Siqueiros Transform fault to the south (~
8°N). The 8°37’ lavas, which lie geographically between the Siqueiros Transform and
the 9°N OSC, also have more radiogenic Sr and Pb than the east limb lavas (Figure 4-
10). This suggests that the mantle source below the EPR from the dying west limb to
the Siqueiros Transform fault is generally more enriched than the mantle beneath the
east limb of the OSC extending up to the Clipperton Transform [Sims et al., 2002; Sims
et al., 2003]. This is consistent with observations that the leading limb of the EPR may
tap a slightly more “enriched” mantle than the trailing limb [Carbotte et al., 2004]. While
it appears to be more enriched overall, simple 2 component mixing of depleted mantle
with an enriched end-member similar to the Siqueiros E-MORB cannot explain the
range of isotopic ratios erupted on this segment of the EPR.
126
Conclusions
The 9°N OSC is one of the most extensively studied 2nd order discontinuities on
the global MOR system and one with the most completely imaged crustal melt lens. It
has erupted a wide compositions ranging from basalts to dacites. Much of the
compositional variability can be ascribed to low-pressure (~1 kbar) fractional
crystallization of N-MORB magma or mixing of ferrobasaltic and high-silica magmas.
The west limb magmas are slightly less evolved than those on the east limb.
Andesitic compositions on the east limb can be divided into two different groups
(high- and low-P2O5) with different major and trace element characteristics. The high-
P2O5 andesites are produced dominantly by extensive fractional crystallization (~75%).
Low-P2O5 andesites are produced through extensive mixing of high-silica dacitic and
ferrobasaltic magma. Magma mixing also explains the compositions of many of the
basaltic andesites erupted on the east limb that are not consistent with calculated
MORB fractional crystallization trends.
The near-absence of high-silica lavas on the southern tip of the east limb and the
entire west limb compared to the northern east ridge axis suggests a different tectono-
magmatic environment in these settings. We believe that this is due to the cooler
temperatures of the ocean crust in these regions as a consequence of decreased
magmatic input. Dacitic lavas are produced from the combination of assimilation and
fractional crystallization in regions where the latent heat of crystallization provides
enough heat to partially melt the surrounding wall rock. The cooler crust at the dying
western limb and the southern ridge tip may allow for extensive fractional crystallization,
however, it is either not enough to heat and melt the surrounding wall rock or these
partial (anatectic?) melts are not erupted.
127
The distribution of evolved lavas and E-MORB lavas across the OSC is not
symmetric, suggesting that the 2nd order discontinuity represents a boundary??
division? in the magmatic plumbing system of the EPR. E-MORB lavas are only
observed on the dying western limb and overall the lavas are less evolved. We suggest
that the lower magma supply at the west limb allows for the preservation and eruption of
E-MORB compositions, whereas the more robust magmatic system on the propagating
east limb overwhelms this signature.
N-MORB lavas on the west limb have more radiogenic Pb and Sr and less
radiogenic Nd compared to east limb N-MORB lavas. Lavas erupted to the south of the
OSC (8°37’N) also have more radiogenic Pb and Sr isotope ratios. This suggests a
slightly different mantle is feeding this section of the EPR and that the OSC provides a
fundamental division between mantle sources beneath the ridge axis.
128
Table 4-1. East limb major element data 9°N OSC sample Rock Type SiO2 TiO2 Al2O3 FeO Mn MgO CaO Na2O K2O P2O5 Cl S Total WC-07 FeTi 49.28 2.60 12.57 14.08 0.24 5.71 9.24 3.21 0.20 0.28 97.42 WC-06 FeTi 49.73 2.49 12.69 14.02 0.24 5.67 9.35 3.28 0.19 0.27 97.94 WC-08 ferrobasalt 49.99 1.75 13.80 11.11 0.21 7.27 10.95 2.90 0.15 0.16 98.29 WC-01 ferrobasalt 50.07 1.83 13.69 11.36 0.21 7.20 10.94 2.94 0.14 0.18 98.57 265-43 ferrobasalt 50.53 1.92 13.88 11.56 0.21 6.98 11.14 2.86 0.13 0.19 0.01 0.25 99.67 WC-02 ferrobasalt 50.13 1.80 13.73 11.31 0.21 7.26 11.04 2.94 0.13 0.20 98.75 265-98 ferrobasalt 50.24 1.94 13.82 11.75 0.21 7.17 10.86 2.86 0.12 0.19 99.16 266-14 ferrobasalt 50.45 2.14 13.48 12.38 0.22 6.69 10.75 3.11 0.16 0.22 99.60 264-08 FeTi 50.10 2.68 12.73 14.06 0.26 5.69 9.58 3.27 0.21 0.28 0.07 0.32 99.24 266-15 ferrobasalt 50.51 2.38 13.09 13.30 0.23 6.27 10.11 3.27 0.18 0.26 99.61 266-51 ferrobasalt 50.55 1.94 13.85 11.59 0.21 7.33 10.97 2.86 0.13 0.19 0.01 99.62 267-14 ferrobasalt 50.55 1.92 14.04 11.57 0.21 6.96 11.05 2.86 0.13 0.21 99.51 265-20 ferrobasalt 50.68 1.85 13.92 11.31 0.22 7.21 11.34 2.95 0.13 0.18 0.01 0.22 100.0 266-48 ferrobasalt 50.57 1.92 14.09 11.56 0.21 7.34 11.00 3.03 0.13 0.19 100.0 266-38 FeTi 50.58 2.00 13.60 12.35 0.23 6.63 10.73 3.15 0.17 0.20 99.63 267-13 ferrobasalt 50.59 1.85 14.08 11.34 0.21 7.11 11.21 2.82 0.12 0.21 99.54 267-03 FeTi 50.60 2.13 13.58 12.49 0.23 6.45 10.15 3.05 0.16 0.26 99.10 264-23 ferrobasalt 50.61 1.87 13.89 11.29 0.22 7.28 11.14 2.92 0.13 0.18 99.52 266-16 ferrobasalt 50.61 2.02 13.67 12.10 0.23 6.87 10.75 3.15 0.16 0.22 99.79 265-44 ferrobasalt 50.62 1.93 14.02 11.72 0.22 6.86 11.11 2.81 0.13 0.19 99.62 266-40 ferrobasalt 50.63 1.77 13.96 11.52 0.20 7.24 11.25 2.96 0.14 0.16 99.84 266-17 ferrobasalt 50.63 1.87 14.00 11.31 0.20 7.24 11.43 3.01 0.14 0.19 100.0 265-99 ferrobasalt 50.65 1.84 13.91 11.42 0.20 7.36 11.00 2.90 0.13 0.18 99.57 266-34 ferrobasalt 50.65 1.84 14.10 11.30 0.21 7.25 11.08 2.90 0.11 0.18 99.63 265-87 ferrobasalt 50.65 1.98 13.84 11.94 0.23 6.84 11.06 2.77 0.14 0.20 99.64 266-32 ferrobasalt 50.67 1.85 14.01 11.38 0.22 7.27 11.13 2.92 0.12 0.18 99.74 265-96 ferrobasalt 50.67 1.83 14.00 11.43 0.19 7.35 11.01 2.85 0.13 0.18 99.65 267-05 FeTi 50.68 1.99 13.69 12.00 0.22 6.39 10.31 3.06 0.16 0.25 98.76 264-05 ferrobasalt 50.69 1.83 13.79 11.42 0.22 7.07 11.14 2.95 0.15 0.18 99.44 267-12 FeTi 50.70 2.20 13.58 12.75 0.22 6.38 10.40 3.03 0.16 0.25 99.67 266-39 FeTi 50.71 2.04 13.62 12.42 0.24 6.57 10.54 3.21 0.19 0.20 99.74 267-08 FeTi 50.71 1.99 13.80 12.03 0.21 6.79 10.77 2.95 0.14 0.22 99.62 265-19 ferrobasalt 50.62 1.83 14.02 11.23 0.21 7.19 11.24 2.93 0.12 0.18 0.01 0.24 99.83 266-33 ferrobasalt 50.73 1.84 14.00 11.35 0.22 7.29 11.13 2.88 0.12 0.20 99.75 266-13 ferrobasalt 50.74 2.29 13.21 12.95 0.23 6.34 10.54 3.21 0.17 0.25 99.93
129
Table 4-1. Continued sample Rock Type SiO2 TiO2 Al2O3 FeO Mn MgO CaO Na2O K2O P2O5 Cl S Total 266-52 ferrobasalt 50.74 1.93 13.92 11.63 0.21 7.30 10.95 2.89 0.13 0.18 99.87 265-14 ferrobasalt 50.75 1.83 13.92 11.20 0.21 7.16 11.19 2.97 0.12 0.18 99.54 WC-05 ferrobasalt 50.75 2.00 13.44 11.53 0.21 6.11 9.69 3.35 0.26 0.27 97.62 266-02 FeTi 50.76 2.08 13.63 12.45 0.22 6.68 10.51 3.11 0.12 0.21 99.77 266-08 ferrobasalt 50.76 1.98 13.88 11.74 0.22 7.12 10.92 2.86 0.13 0.20 99.81 265-86 ferrobasalt 50.76 1.99 13.83 11.87 0.22 6.79 11.08 2.72 0.14 0.20 99.59 267-09 ferrobasalt 50.76 1.95 13.90 11.99 0.22 6.90 10.82 2.95 0.14 0.24 99.87 265-97 ferrobasalt 50.77 2.02 13.78 12.02 0.22 7.03 10.87 2.61 0.13 0.21 99.65 265-82 ferrobasalt 50.77 1.95 13.81 11.70 0.22 6.93 11.21 2.69 0.14 0.19 0.01 99.60 265-92 ferrobasalt 50.77 1.98 13.91 11.77 0.22 7.08 10.99 2.93 0.13 0.20 99.97 265-10 ferrobasalt 50.78 1.86 13.94 11.33 0.22 7.13 11.23 2.99 0.13 0.18 99.78 265-52 ferrobasalt 50.78 1.93 14.01 11.77 0.22 6.99 11.18 2.92 0.13 0.20 100.1 265-15 ferrobasalt 50.79 1.86 13.90 11.25 0.21 7.14 11.24 2.99 0.13 0.17 99.67 267-11 FeTi 50.79 2.20 13.66 12.69 0.22 6.44 10.32 2.95 0.16 0.25 99.68 266-19 ferrobasalt 50.79 1.86 13.73 11.87 0.23 6.88 11.03 3.09 0.14 0.20 99.83 266-12 ferrobasalt 50.80 2.15 13.43 12.67 0.24 6.51 10.66 3.27 0.18 0.22 100.1 265-68 ferrobasalt 50.81 1.81 14.02 11.17 0.21 7.29 11.51 2.79 0.13 0.17 0.01 99.92 265-53 ferrobasalt 50.81 1.96 14.05 11.83 0.22 6.85 11.17 2.92 0.13 0.19 100.1 266-18 ferrobasalt 50.82 2.01 13.70 12.08 0.22 6.91 10.88 3.14 0.16 0.22 100.1 266-37 FeTi 50.82 2.03 13.57 12.26 0.22 6.73 10.58 3.02 0.15 0.20 99.59 265-62 ferrobasalt 50.82 1.95 13.93 11.72 0.22 6.92 11.19 2.90 0.13 0.19 99.97 266-11 ferrobasalt 50.82 2.23 13.30 12.93 0.24 6.43 10.34 3.26 0.18 0.27 100.0 267-07 ferrobasalt 50.82 1.96 13.94 11.84 0.20 6.96 10.86 2.94 0.14 0.22 99.89 265-13 ferrobasalt 50.83 1.85 14.01 11.22 0.22 7.21 11.28 2.93 0.12 0.18 99.83 265-11 ferrobasalt 50.84 1.80 14.00 11.11 0.21 7.23 11.29 2.95 0.12 0.17 99.72 267-01 FeTi 50.84 2.03 13.65 12.39 0.21 6.34 10.18 3.07 0.16 0.23 99.11 265-45 ferrobasalt 50.84 1.94 14.12 11.80 0.21 6.87 11.21 2.89 0.13 0.19 100.2 265-88 FeTi 50.85 2.02 13.88 11.99 0.22 6.71 11.06 2.85 0.15 0.21 0.02 99.94 266-09 ferrobasalt 50.85 1.93 14.00 11.43 0.21 7.10 11.08 2.84 0.12 0.19 99.76 265-17 FeTi 50.68 2.09 13.45 12.61 0.23 6.48 10.75 3.14 0.13 0.20 0.00 0.27 100.0 265-119 ferrobasalt 50.86 2.00 13.88 11.77 0.22 7.12 10.86 2.90 0.13 0.19 99.94 265-73 ferrobasalt 50.86 1.96 13.98 11.79 0.22 6.97 11.20 2.83 0.13 0.20 100.1 267-02 ferrobasalt 50.87 1.91 14.10 11.45 0.20 6.73 10.73 2.71 0.15 0.24 99.09 265-16 ferrobasalt 50.87 1.84 14.13 11.23 0.22 7.22 11.22 3.01 0.12 0.17 100.0 265-04 ferrobasalt 50.87 1.84 14.01 11.21 0.21 7.15 11.38 3.00 0.13 0.18 99.97
130
Table 4-1. Continued sample Rock Type SiO2 TiO2 Al2O3 FeO Mn MgO CaO Na2O K2O P2O5 Cl S Total 267-06 ferrobasalt 50.88 1.97 14.02 11.91 0.22 6.91 10.84 2.84 0.14 0.22 99.94 266-26 FeTi 50.88 2.13 13.57 12.82 0.22 6.34 10.30 3.06 0.15 0.23 99.69 265-79 ferrobasalt 50.89 1.89 13.99 11.75 0.22 6.92 11.20 2.86 0.14 0.19 100.0 266-04 FeTi 50.89 2.36 13.25 13.41 0.23 6.18 9.94 3.13 0.17 0.23 99.79 265-72 ferrobasalt 50.89 1.93 14.05 11.76 0.23 6.98 11.13 2.85 0.13 0.20 100.1 266-22 FeTi 50.90 1.92 13.87 12.05 0.21 6.79 10.73 2.97 0.13 0.21 99.79 265-39 ferrobasalt 50.90 1.95 13.80 11.57 0.21 7.13 11.19 2.95 0.13 0.18 100.0 264-06 ferrobasalt 50.90 1.83 13.75 11.37 0.22 7.06 11.14 2.96 0.15 0.17 99.56 265-46 ferrobasalt 50.91 1.94 14.09 11.88 0.22 6.80 11.20 2.99 0.14 0.20 100.4 266-23 FeTi 50.91 1.89 13.89 12.00 0.20 6.80 10.80 2.97 0.13 0.21 99.80 266-24 FeTi 50.91 1.90 13.87 12.04 0.22 6.81 10.70 2.98 0.13 0.22 99.79 265-08 ferrobasalt 50.92 1.86 13.94 11.25 0.21 7.21 11.37 2.96 0.12 0.16 100.0 265-18 ferrobasalt 50.75 1.82 14.01 11.31 0.21 7.21 11.26 2.97 0.12 0.18 0.01 0.22 100.1 265-09 ferrobasalt 50.93 1.85 13.96 11.28 0.22 7.20 11.28 2.97 0.13 0.18 99.99 266-27 FeTi 50.94 2.13 13.59 12.83 0.23 6.32 10.19 3.11 0.14 0.25 99.74 265-89 ferrobasalt 50.94 1.88 14.01 11.57 0.22 6.95 11.31 2.76 0.14 0.18 99.94 266-30 ferrobasalt 50.95 1.88 13.78 11.90 0.21 7.02 10.73 2.97 0.12 0.19 99.76 265-41 ferrobasalt 50.96 1.92 14.14 11.47 0.22 7.10 11.31 2.84 0.13 0.19 100.3 264-04 ferrobasalt 50.54 1.84 13.66 11.53 0.21 6.97 11.06 2.98 0.15 0.18 0.01 0.25 99.39 266-28 FeTi 50.97 2.15 13.52 12.86 0.23 6.34 10.27 3.07 0.15 0.25 99.80 266-35 FeTi 50.97 1.91 13.74 12.13 0.22 6.91 10.59 3.00 0.12 0.20 99.80 266-07 FeTi 50.97 2.36 13.15 13.50 0.25 6.16 9.72 3.18 0.18 0.25 99.71 267-10 ferrobasalt 50.97 1.97 13.99 11.97 0.21 6.92 10.81 2.95 0.14 0.23 100.2 265-05 ferrobasalt 50.98 1.88 13.88 11.38 0.22 6.99 11.34 3.02 0.13 0.18 99.99 265-78 ferrobasalt 50.98 1.88 14.01 11.70 0.21 6.97 11.18 2.86 0.14 0.19 100.1 264-07 ferrobasalt 50.98 1.82 13.84 11.40 0.21 7.08 11.19 2.98 0.14 0.18 99.80 267-15 ferrobasalt 50.98 2.00 13.80 11.81 0.21 6.97 10.84 2.86 0.14 0.21 99.82 266-29 FeTi 50.99 2.19 13.32 12.88 0.23 6.39 10.15 3.07 0.15 0.24 99.61 266-01 FeTi 50.99 2.00 13.51 12.39 0.23 6.76 10.71 3.04 0.12 0.17 99.91 265-71 ferrobasalt 51.00 1.89 13.96 11.68 0.22 6.96 11.19 2.82 0.14 0.19 100.0 265-80 ferrobasalt 51.01 1.92 13.95 11.70 0.22 6.98 11.24 2.87 0.13 0.19 100.2 265-74 ferrobasalt 51.01 1.99 13.93 11.92 0.22 6.81 11.05 2.90 0.14 0.20 100.2 265-02 ferrobasalt 51.04 1.89 13.98 11.39 0.22 6.99 11.37 3.03 0.12 0.18 100.2 266-36 ferrobasalt 51.05 1.89 13.83 11.88 0.22 7.04 10.66 2.97 0.13 0.20 99.87 266-21 ferrobasalt 51.06 1.85 13.95 11.93 0.22 6.94 10.80 2.96 0.13 0.19 100.0
131
Table 4-1. Continued sample Rock Type SiO2 TiO2 Al2O3 FeO Mn MgO CaO Na2O K2O P2O5 Cl S Total 265-81 FeTi 51.06 2.00 14.13 11.93 0.22 7.00 11.24 2.39 0.13 0.19 100.3 265-22 ferrobasalt 51.06 1.86 14.05 11.35 0.22 7.15 11.37 3.01 0.12 0.18 100.4 266-25 FeTi 51.06 1.98 13.68 12.26 0.22 6.57 10.61 2.98 0.14 0.22 99.72 266-41 ferrobasalt 51.06 1.86 14.09 11.30 0.22 6.84 10.74 3.22 0.21 0.18 99.71 264-03 ferrobasalt 51.07 1.84 13.89 11.34 0.21 7.07 11.19 2.98 0.15 0.17 99.90 266-10 ferrobasalt 51.07 1.93 13.95 11.50 0.21 6.97 10.97 2.89 0.14 0.21 99.84 266-03 ferrobasalt 51.07 1.77 14.24 10.83 0.21 7.04 10.86 3.06 0.16 0.21 99.44 265-03 ferrobasalt 51.08 1.89 13.89 11.40 0.22 7.02 11.27 3.03 0.13 0.19 100.1 265-76 ferrobasalt 51.08 1.89 13.98 11.61 0.22 6.98 11.25 2.84 0.14 0.19 100.2 265-06 ferrobasalt 51.09 1.85 13.89 11.44 0.21 7.00 11.32 3.00 0.13 0.17 100.1 266-31 ferrobasalt 51.10 1.88 13.78 11.90 0.22 7.09 10.72 2.96 0.13 0.18 99.95 265-07 ferrobasalt 51.11 1.90 13.98 11.47 0.22 6.96 11.24 3.02 0.13 0.19 100.2 265-38 ferrobasalt 51.13 1.89 13.91 11.34 0.22 7.09 11.11 3.00 0.15 0.19 100.0 265-30 ferrobasalt 51.13 1.87 14.10 11.45 0.22 6.91 11.32 3.01 0.15 0.18 100.3 265-29 ferrobasalt 51.17 1.88 14.04 11.44 0.22 6.88 11.27 3.04 0.15 0.18 100.2 265-121 ferrobasalt 51.17 1.89 14.05 11.44 0.21 7.12 10.93 2.91 0.14 0.19 100.1 265-26 ferrobasalt 51.17 1.94 13.73 12.03 0.22 6.84 11.09 3.08 0.16 0.19 100.5 265-51 ferrobasalt 51.17 1.92 14.19 11.66 0.22 7.11 11.27 2.93 0.13 0.19 100.8 265-21 ferrobasalt 51.18 1.83 14.04 11.33 0.21 7.20 11.37 2.98 0.13 0.17 100.4 265-36 ferrobasalt 51.18 1.93 13.90 11.42 0.22 7.10 11.12 2.96 0.14 0.19 100.2 265-35 ferrobasalt 51.21 1.94 13.78 11.57 0.22 6.95 11.14 2.99 0.15 0.20 100.2 265-28 ferrobasalt 51.23 1.79 14.12 11.19 0.21 7.18 11.37 2.99 0.15 0.18 100.4 265-33 ferrobasalt 51.23 1.97 13.93 11.48 0.22 6.83 11.16 3.02 0.15 0.18 100.2 265-37 ferrobasalt 51.23 1.93 13.95 11.36 0.22 7.11 11.15 3.00 0.14 0.19 100.3 265-104 ferrobasalt 51.25 1.90 13.79 11.65 0.22 6.90 10.59 3.00 0.16 0.22 0.05 99.69 265-27 ferrobasalt 51.27 1.79 14.12 11.13 0.21 7.10 11.47 2.99 0.14 0.17 100.4 265-34 ferrobasalt 51.30 1.99 13.93 11.52 0.22 6.92 11.11 3.01 0.15 0.19 100.3 265-107 ferrobasalt 51.30 1.92 13.83 11.67 0.21 6.85 10.51 3.07 0.16 0.23 99.76 265-93 ferrobasalt 51.32 1.97 13.90 11.82 0.22 6.84 10.62 3.02 0.15 0.21 100.1 265-23 ferrobasalt 51.32 1.88 13.78 11.92 0.22 6.92 10.98 3.02 0.13 0.19 100.4 265-01 FeTi 51.36 2.03 13.82 11.89 0.23 6.50 10.71 3.21 0.17 0.23 100.2 265-32 ferrobasalt 51.36 1.96 13.89 11.53 0.22 6.90 11.11 3.00 0.15 0.18 100.3 265-115 FeTi 51.38 2.17 13.34 12.72 0.23 6.16 9.77 3.24 0.19 0.25 99.45 265-12 ferrobasalt 51.41 1.92 13.75 11.45 0.22 6.60 10.69 3.11 0.19 0.20 99.54 265-31 FeTi 51.62 2.01 13.89 11.72 0.23 6.85 11.07 3.05 0.15 0.20 100.8
132
Table 4-1. Continued
sample Rock Type SiO2 TiO2 Al2O3 FeO Mn MgO CaO Na2O K2O P2O5 Cl S Total
265-110 ferrobasalt 51.65 1.90 13.81 11.69 0.20 6.71 10.41 3.07 0.17 0.23 99.84 265-114 FeTi 51.66 2.18 13.39 12.83 0.24 5.88 9.53 3.29 0.22 0.27 99.48 265-111 FeTi 51.78 2.17 13.38 12.65 0.22 5.73 9.40 3.35 0.23 0.30 99.21 265-112 FeTi 51.79 1.97 13.67 12.00 0.22 6.41 10.11 3.12 0.19 0.24 99.72 266-06 ferrobasalt 51.80 1.70 14.21 10.87 0.20 6.72 10.59 3.05 0.18 0.14 99.47 265-113 FeTi 51.92 2.17 13.40 12.82 0.23 5.93 9.48 3.28 0.22 0.28 99.72 WC-04 FeTi 51.94 2.00 12.96 12.63 0.25 4.88 8.59 3.57 0.33 0.44 97.59 265-61 ferrobasalt 51.96 1.81 14.06 11.26 0.21 6.43 10.62 3.05 0.19 0.18 99.76
264-18 basaltic andesite 52.15 1.77 14.04 10.89 0.20 6.05 9.94 3.27 0.26 0.20 98.77
266-45 basaltic andesite 52.24 1.79 13.86 11.28 0.22 6.46 10.18 3.33 0.27 0.17 99.80
264-11 basaltic andesite 52.36 1.72 13.79 10.75 0.20 6.55 10.47 3.13 0.24 0.19 99.40
264-17 basaltic andesite 52.61 1.77 14.12 10.95 0.20 6.23 10.07 3.27 0.25 0.20 99.66
265-105 basaltic andesite 52.77 1.90 13.33 12.06 0.22 5.58 9.26 3.44 0.27 0.32 99.13
265-116 basaltic andesite 52.86 1.84 13.77 11.64 0.21 5.73 9.38 3.37 0.26 0.28 99.33
265-49 basaltic andesite 52.81 1.78 14.00 11.09 0.21 5.98 10.03 3.23 0.27 0.20 0.04 0.21 99.84
265-60 basaltic andesite 53.08 1.74 14.06 10.93 0.21 5.93 10.12 3.26 0.27 0.19 99.80
264-21 basaltic andesite 53.15 1.82 14.00 11.09 0.21 5.82 9.69 3.41 0.29 0.22 99.71
265-122 basaltic andesite 53.28 2.06 13.40 12.27 0.23 5.36 8.84 3.52 0.30 0.23 99.50
265-106 basaltic andesite 53.33 1.94 13.26 12.35 0.24 5.11 8.71 3.67 0.32 0.39 0.19 99.31
264-13 basaltic andesite 53.41 1.68 14.44 10.43 0.20 5.76 9.53 3.42 0.33 0.20 99.40
265-118 basaltic andesite 53.42 1.86 13.57 11.90 0.21 5.16 8.85 3.54 0.31 0.35 99.16
133
Table 4-1. Continued
sample Rock Type SiO2 TiO2 Al2O3 FeO Mn MgO CaO Na2O K2O P2O5 Cl S Total
264-19 basaltic andesite 53.51 1.77 13.90 10.83 0.21 5.72 9.52 3.44 0.31 0.20 99.41
265-48 basaltic andesite 53.85 2.36 13.06 13.33 0.24 4.44 8.36 3.57 0.35 0.26 0.05 0.28 100.2
264-22 basaltic andesite 53.58 1.75 13.93 10.90 0.20 5.80 9.62 3.45 0.31 0.20 99.75
265-108
basaltic andesite 53.77 1.95 13.28 12.60 0.23 5.08 8.80 3.63 0.32 0.38 0.20 100.0
265-24 basaltic andesite 54.07 2.35 12.90 13.64 0.25 3.71 7.61 3.75 0.50 0.58 99.36
265-50 basaltic andesite 54.15 2.14 13.41 12.19 0.22 4.54 8.45 3.59 0.40 0.28 0.09 0.22 99.67
266-43 basaltic andesite 54.14 1.90 13.56 12.05 0.24 4.42 8.16 4.03 0.41 0.32 99.22
265-59 basaltic andesite 54.24 2.15 13.42 12.27 0.23 4.61 8.52 3.53 0.40 0.28 99.65
265-58 basaltic andesite 54.25 2.12 13.56 12.13 0.22 4.56 8.52 3.57 0.39 0.27 99.59
265-120
basaltic andesite 54.33 2.06 12.78 13.10 0.24 4.22 7.87 3.72 0.38 0.52 99.23
266-62 basaltic andesite 54.48 1.53 14.24 9.87 0.18 5.33 8.89 3.46 0.41 0.19 98.58
265-55 basaltic andesite 54.55 2.10 13.45 12.19 0.22 4.51 8.38 3.62 0.41 0.28 99.71
264-20 basaltic andesite 54.59 1.71 13.95 10.68 0.21 5.10 8.84 3.60 0.39 0.21 99.28
265-123
basaltic andesite 54.62 1.91 13.46 11.67 0.22 4.93 8.40 3.61 0.36 0.22 99.37
265-56 basaltic andesite 54.98 2.13 13.68 12.13 0.23 4.42 8.36 3.63 0.42 0.28 100.2
265-124
basaltic andesite 55.06 1.69 13.55 10.78 0.19 5.03 8.44 3.54 0.40 0.20 98.88
265-54 basaltic andesite 55.14 2.04 13.65 11.75 0.22 4.20 8.11 3.64 0.45 0.27 99.47
264-10 basaltic andesite 55.00 1.62 13.34 10.23 0.19 5.27 8.76 3.56 0.43 0.21 0.12 0.20 98.93
134
Table 4-1. Continued
sample Rock Type SiO2 TiO2 Al2O3 FeO Mn MgO CaO Na2O K2O P2O5 Cl S Total
265-102 basaltic andesite 55.31 2.22 12.51 14.05 0.27 3.34 7.14 3.85 0.44 0.65 99.79
WC-03 basaltic andesite 55.46 2.02 12.02 13.80 0.27 2.50 6.33 4.10 0.56 0.74 97.80
265-126 basaltic andesite 55.48 1.66 13.97 10.46 0.18 4.78 8.37 3.67 0.41 0.19 99.18
264-12 basaltic andesite 55.51 1.64 13.35 10.18 0.20 4.98 8.47 3.65 0.47 0.21 98.66
265-125 basaltic andesite 55.59 1.66 13.63 10.73 0.19 4.89 8.30 3.64 0.41 0.20 99.22
265-109 basaltic andesite 55.66 1.81 12.83 12.31 0.23 3.90 7.47 3.76 0.45 0.43 98.85
265-77 basaltic andesite 55.74 2.13 13.52 11.98 0.21 3.95 7.80 3.61 0.48 0.31 99.73
265-117 basaltic andesite 55.77 1.81 12.90 12.26 0.23 3.90 7.45 3.75 0.43 0.44 98.95
265-101 basaltic andesite 55.90 2.14 12.29 13.95 0.26 3.02 6.76 4.03 0.49 0.65 99.50
264-16 basaltic andesite 56.02 2.06 13.21 12.23 0.25 4.37 8.10 2.40 0.45 0.37 99.46
265-57 basaltic andesite 56.29 2.12 12.59 12.75 0.25 3.69 7.46 3.59 0.52 0.40 99.64
265-103 basaltic andesite 56.50 2.01 12.33 13.74 0.26 2.74 6.45 3.77 0.52 0.65 0.30 98.96
265-91 basaltic andesite 56.83 2.10 12.31 14.23 0.28 2.09 6.24 3.45 0.60 0.78 0.31 98.91
266-61 andesite 57.47 1.70 14.07 10.35 0.19 3.13 6.57 3.76 0.63 0.24 98.13 266-59 andesite 57.57 1.58 13.16 10.41 0.20 3.66 6.71 3.60 0.61 0.23 97.74 265-90 andesite 58.08 1.91 12.43 13.69 0.27 1.74 5.76 3.51 0.66 0.74 0.34 98.80 265-100 andesite 58.09 1.76 12.64 12.68 0.24 1.89 5.51 3.82 0.63 0.54 97.80 266-05 andesite 58.40 1.86 13.18 11.14 0.21 3.04 6.29 3.77 0.66 0.30 98.85 266-55 andesite 59.52 1.65 13.14 10.83 0.20 2.13 5.30 3.86 0.77 0.42 97.82 266-54 andesite 59.65 1.72 13.22 11.25 0.21 2.28 5.60 3.91 0.75 0.43 0.42 99.02 265-69 andesite 61.02 1.72 13.62 9.97 0.18 1.91 5.38 3.89 0.80 0.27 0.20 98.76 265-25 andesite 61.17 1.48 13.63 9.84 0.19 1.92 5.03 4.28 0.99 0.43 98.96 264-14 andesite 61.75 1.30 13.46 8.71 0.17 2.47 5.54 3.94 0.83 0.22 98.39
135
Table 4-1. Continued
sample Rock Type SiO2 TiO2 Al2O3 FeO Mn MgO CaO Na2O K2O P2O5 Cl S Total
266-49 andesite 62.34 1.27 13.15 8.64 0.15 2.19 4.86 4.18 0.91 0.21 97.90 266-57 andesite 62.47 1.30 13.16 9.20 0.16 1.59 4.37 4.11 0.98 0.29 0.51 97.64 266-56 andesite 62.47 1.35 13.25 9.24 0.17 1.58 4.45 3.95 0.96 0.33 97.74 265-66 andesite 62.81 1.43 13.13 9.05 0.18 1.99 4.98 3.86 0.89 0.24 0.23 98.55 266-58 dacite 63.01 1.10 13.06 8.43 0.16 1.75 4.34 3.63 0.96 0.26 96.69 265-65 dacite 63.79 1.26 13.25 8.14 0.15 1.34 4.21 3.84 0.97 0.22 97.17 265-64 dacite 64.04 1.28 13.12 8.27 0.16 1.60 4.45 3.46 0.97 0.20 0.24 97.55 265-67 dacite 64.10 1.34 13.33 8.49 0.16 1.49 4.41 3.93 0.95 0.23 98.42 266-50 dacite 64.26 1.07 13.17 8.08 0.14 1.27 3.78 4.23 1.10 0.24 97.35 266-53 dacite 64.28 1.06 13.31 8.06 0.14 1.12 3.73 4.16 1.09 0.25 0.65 97.20 265-84 dacite 64.39 1.13 13.17 8.18 0.15 1.23 3.92 3.41 1.19 0.22 96.98 265-63 dacite 64.43 1.29 13.26 8.22 0.15 1.29 4.21 3.71 0.99 0.21 97.76 266-47 dacite 64.53 0.99 13.18 7.74 0.14 1.02 3.53 4.94 1.22 0.23 97.54 265-85 dacite 65.01 1.06 13.13 7.99 0.16 1.18 3.78 3.67 1.22 0.20 0.64 97.41 266-46 dacite 65.03 0.94 12.90 7.17 0.15 1.41 3.71 4.76 1.19 0.17 97.43 265-94 dacite 65.22 0.97 13.04 7.90 0.14 1.13 3.54 4.29 1.14 0.23 97.61 264-09 dacite 65.76 0.89 13.15 7.03 0.13 1.06 3.48 4.24 1.21 0.20 0.58 0.06 97.78 265-70 dacite 66.26 0.87 13.20 7.17 0.14 0.80 3.23 4.08 1.33 0.19 0.70 97.27 265-42 dacite 66.46 0.94 13.04 7.92 0.16 0.89 3.50 3.99 1.20 0.21 0.51 0.07 98.91 265-83 dacite 67.46 0.76 13.27 6.68 0.13 0.67 2.98 3.88 1.37 0.16 0.67 97.37 265-95 dacite 67.46 0.77 13.10 6.47 0.12 0.94 3.01 4.43 1.21 0.15 97.67 265-40 dacite
136
Table 4-2. Trace element data 9°N OSC sample 265-43 264-08 266-33 267-09 265-88 265-72 266-22 265-18 264-04 266-28 266-07 265-78 267-15 266-01 Li 7.78 9.99 7.52 8.23 7.97 7.74 8.24 7.48 8.47 9.21 7.70 7.77 7.94 5.99 Sc 42 42 41 42 41 42 43 43 53 40 36 43 41 33 V 347 450 321 354 351 350 359 338 406 357 337 353 330 273 Cr 108 17.97 155 87.16 68.56 96.21 50.29 166 119 40.69 9.33 84.45 138 40.62 Co 41 43 40 41 42 42 43 42 51 41 38 43 40 34 Ni 54 34 48 46 47 54 38 53 72 32 47 52 51 38 Cu 59 60 61 60 60 60 61 64 76 55 50 64 58 51 Zn 94 116 88 96 97 96 99 95 103 100 93 95 93 79 Ga 18 21 18 19 18 18 19 18 24 20 17 18 18 14 Rb 1.14 2.12 1.20 1.50 1.25 1.09 1.05 1.00 1.64 1.50 1.55 1.22 1.37 0.98 Sr 120 126 114 113 124 117 120 126 146 111 98 125 108 89 Y 44 59 39 44 44 44 45 42 45 43 42 43 42 32 Zr 126 180 125 140 133 126 130 122 142 157 139 124 136 95 Nb 3.08 5.46 2.91 3.65 3.46 3.07 3.03 2.76 4.01 3.87 3.84 3.19 3.42 2.20 Cs 0.01 0.03 0.02 0.02 0.01 0.01 0.01 0.01 0.02 0.02 0.02 0.02 0.02 0.02 Ba 8.37 17.07 8.25 11.12 10.31 8.44 7.85 6.87 13.20 10.35 11.73 9.49 9.99 6.46 La 4.54 6.81 4.13 4.79 4.94 4.47 4.61 4.27 4.98 5.40 4.84 4.64 4.50 3.10 Ce 14.16 20.51 13.18 14.97 14.97 13.90 14.38 13.40 15.49 17.22 14.95 14.06 14.13 9.93 Pr 2.34 3.32 2.29 2.56 2.46 2.31 2.38 2.21 2.61 2.95 2.47 2.31 2.44 1.76 Nd 12.6 17.3 12.1 13.5 13.2 12.5 12.8 11.9 13.9 15.5 13.1 12.5 13.0 9.5 Sm 4.42 5.99 4.16 4.69 4.48 4.37 4.45 4.15 4.78 5.48 4.49 4.35 4.47 3.32 Eu 1.48 1.94 1.45 1.58 1.52 1.46 1.52 1.40 1.67 1.77 1.49 1.47 1.49 1.18 Gd 5.78 7.74 5.53 6.10 5.92 5.75 5.87 5.46 6.18 7.03 5.78 5.71 5.82 4.42 Tb 1.08 1.45 1.02 1.14 1.11 1.07 1.11 1.02 1.15 1.30 1.07 1.07 1.07 0.82 Dy 7.22 9.69 6.62 7.36 7.32 7.09 7.31 6.74 7.46 8.41 6.92 7.04 7.04 5.34 Ho 1.52 2.05 1.43 1.56 1.53 1.51 1.55 1.44 1.57 1.82 1.46 1.49 1.50 1.14 Er 4.39 5.94 4.04 4.47 4.48 4.34 4.51 4.15 4.55 5.18 4.27 4.33 4.27 3.26 Tm 0.67 0.91 0.62 0.68 0.68 0.66 0.69 0.63 0.68 0.80 0.64 0.66 0.66 0.50 Yb 4.43 6.01 3.97 4.35 4.47 4.33 4.55 4.16 4.41 5.15 4.17 4.37 4.19 3.19 Lu 0.68 0.93 0.60 0.66 0.68 0.67 0.70 0.64 0.68 0.76 0.63 0.66 0.64 0.48 Hf 3.44 4.83 3.20 3.58 3.59 3.40 3.50 3.23 3.59 4.04 3.49 3.37 3.46 2.50 Ta 0.21 0.37 0.20 0.25 0.23 0.21 0.21 0.19 0.28 0.26 0.27 0.22 0.22 0.16 Pb 0.39 0.61 0.60 0.59 0.45 0.40 0.41 0.38 0.59 0.87 0.58 0.42 0.77 0.42 Th 0.18 0.34 0.18 0.23 0.20 0.18 0.18 0.16 0.22 0.25 0.23 0.20 0.21 0.13 U 0.08 0.13 0.08 0.09 0.08 0.07 0.07 0.07 0.09 0.10 0.09 0.08 0.08 0.06
137
Table 4-2. Continued sample 266-10 265-35 265-31 265-113 265-24 265-50 265-125 265-109 264-16 265-57 265-103 265-91 265-90 Li 8.51 7.85 7.91 11.01 15.32 28.86 14.31 19.24 18.52 20.73 22.27 23.57 28.20 Sc 46 42 42 38 27 13 31 29 33 31 26 25 24 V 347 342 342 323 196 73 247 176 230 217 129 93 70 Cr 150 127 112. 32.12 12.25 4.07 50.08 29.19 44.46 31.96 6.40 2.31 4.82 Co 44 41 41 39 27 13 31 27 30 29 23 21 20 Ni 57 57 54 34 16 7 29 23 25 18 10 6 6 Cu 65 62 59 51 31 20 46 35 34 27 27 23 22 Zn 99 96 94 103 118 100 97 126 129 137 144 147 174 Ga 21 18 18 21 23 29 23 27 24 26 30 31 30 Rb 1.51 1.18 1.25 2.23 4.70 12.5 4.22 4.42 4.60 5.34 5.30 5.74 6.34 Sr 121 119 118 111 99 69 96 104 113 109 103 106 114 Y 47 45 44 61 95 138 72 110 101 116 135 144 170 Zr 156 130 131 229 468 967 380 557 371 425 671 724 680 Nb 3.87 3.15 3.21 5.67 10.33 14.82 6.33 10.95 8.89 10.06 13.51 14.52 15.81 Cs 0.02 0.02 0.02 0.03 0.06 0.12 0.05 0.05 0.05 0.07 0.06 0.06 0.08 Ba 11.08 8.69 8.75 15.60 29.41 58.22 24.31 28.72 29.81 32.73 33.81 36.06 42.35 La 5.17 4.66 4.71 7.63 14.86 28.00 11.10 15.75 13.61 15.76 19.57 21.43 23.72 Ce 16.46 14.43 14.56 23.69 45.03 79.99 32.73 48.54 41.17 47.56 60.43 66.14 73.19 Pr 2.77 2.38 2.40 3.92 7.01 11.64 5.08 7.76 6.42 7.34 9.64 10.54 11.55 Nd 14.7 12.7 12.7 20.1 35.1 52.4 24.6 38.9 30.9 35.3 48.6 52.9 55.8 Sm 5.00 4.47 4.44 6.68 10.74 15.64 7.54 12.07 10.37 11.90 14.99 16.18 18.61 Eu 1.66 1.48 1.48 1.99 2.75 2.96 1.98 3.15 2.68 2.98 3.78 4.10 4.57 Gd 6.40 5.82 5.83 8.52 12.94 17.82 9.21 14.75 12.89 14.65 18.33 19.83 22.78 Tb 1.19 1.10 1.09 1.58 2.36 3.41 1.72 2.70 2.44 2.78 3.34 3.57 4.26 Dy 7.71 7.24 7.24 10.13 15.14 22.37 11.31 17.61 16.19 18.52 21.76 23.38 27.74 Ho 1.61 1.54 1.54 2.20 3.19 4.79 2.39 3.71 3.44 3.93 4.52 4.89 5.88 Er 4.68 4.44 4.45 6.26 9.36 14.29 7.09 10.97 10.16 11.60 13.34 14.24 17.37 Tm 0.70 0.68 0.68 0.96 1.44 2.30 1.12 1.69 1.58 1.81 2.05 2.20 2.68 Yb 4.55 4.45 4.45 6.17 9.27 15.29 7.23 10.88 10.43 11.91 13.19 14.01 17.70 Lu 0.70 0.69 0.68 0.95 1.42 2.29 1.10 1.66 1.62 1.84 2.07 2.14 2.73 Hf 3.90 3.54 3.52 5.72 11.27 23.87 9.42 13.27 10.10 11.55 15.95 17.25 17.78 Ta 0.27 0.21 0.22 0.36 1.50 0.89 0.45 0.76 0.61 0.68 0.95 1.03 1.03 Pb 0.68 0.46 0.42 0.95 1.63 3.54 1.73 2.55 1.65 1.89 3.28 3.76 2.41 Th 0.22 0.19 0.20 0.40 0.84 2.53 0.74 0.78 0.76 0.88 0.96 1.05 1.15 U 0.09 0.08 0.08 0.15 0.31 0.84 0.28 0.30 0.30 0.35 0.36 0.39 0.46
138
Table 4-2. Continued sample 265-100 266-05 265-69 265-25 264-14 266-56 265-66 265-65 265-64 265-67 266-53 265-84 Li 27.14 20.40 26.34 28.11 26.38 27.40 36.88 31.68 33.61 31.78 30.17 28.59 Sc 24 26 22 19 24 17 26 17 20 18 15 15 V 100 216 180 87 160 93 184 122 140 121 61 102 Cr 9.51 9.77 16.05 9.15 40.69 5.34 15.33 12.94 12.27 12.40 3.81 1.75 Co 21 26 21 17 21 15 23 15 17 16 13 14 Ni 10 15 11 9 22 7 15 9 10 9 6 7 Cu 26 33 22 26 29 22 24 17 19 18 18 21 Zn 150 107 112 124 114 106 143 110 124 122 109 100 Ga 33 26 28 28 26 29 42 30 35 28 29 29 Rb 6.51 7.01 7.35 10.8 7.79 10.7 11.6 9.1 10.5 9.6 12.8 12.7 Sr 112 92 88 97 97 81 114 76 90 89 86 70 Y 153 105 114 154 124 133 164 132 148 142 157 133 Zr 721 645 605 881 542 901 945 735 842 622 856 968 Nb 16.24 10.90 11.39 17.36 11.40 15.17 16.79 12.98 14.81 12.97 15.52 14.61 Cs 0.07 0.08 0.08 0.14 0.10 0.11 0.12 0.11 0.12 0.12 0.14 0.13 Ba 41.77 38.30 40.22 58.57 43.51 53.60 62.57 50.53 57.07 52.86 64.61 59.79 La 23.67 17.91 18.93 28.81 19.39 25.36 28.95 23.25 26.31 23.43 28.53 27.27 Ce 72.67 52.40 55.75 84.42 56.99 73.59 84.22 67.48 76.53 67.97 82.17 77.81 Pr 11.30 7.89 8.40 12.25 8.47 10.86 12.53 10.06 11.37 10.05 11.87 11.21 Nd 56.3 37.2 39.7 54.9 39.1 49.6 58.3 46.4 52.8 45.9 53.7 50.9 Sm 17.27 11.43 12.26 16.73 12.38 14.91 17.07 14.09 15.37 14.62 16.25 14.32 Eu 4.20 2.60 2.81 3.31 2.75 3.07 3.78 3.00 3.35 3.12 3.18 2.80 Gd 20.71 13.66 14.69 19.55 14.82 17.58 20.38 16.83 18.13 17.43 19.00 16.58 Tb 3.79 2.53 2.76 3.68 2.86 3.25 3.83 3.19 3.43 3.36 3.65 3.09 Dy 24.73 16.76 18.27 24.22 19.28 21.34 25.32 21.08 22.81 22.85 24.21 20.31 Ho 5.24 3.60 3.98 5.17 4.14 4.60 5.42 4.58 4.88 4.92 5.19 4.33 Er 15.41 10.63 11.66 15.49 12.44 13.59 16.36 13.55 14.65 14.86 15.79 13.16 Tm 2.38 1.68 1.86 2.49 1.98 2.18 2.58 2.19 2.33 2.37 2.50 2.10 Yb 15.34 10.90 12.12 16.50 13.15 14.05 16.83 14.25 15.20 15.66 16.74 13.66 Lu 2.36 1.67 1.85 2.64 2.04 2.16 2.55 2.15 2.31 2.44 2.56 2.10 Hf 17.09 15.73 15.48 22.31 14.75 21.69 23.08 18.69 20.95 17.54 22.28 23.14 Ta 1.14 0.66 0.69 1.13 0.80 0.91 1.23 0.81 1.10 0.92 1.03 1.39 Pb 3.08 2.02 2.96 3.07 2.67 3.21 6.24 3.33 5.05 3.41 3.15 4.89 Th 1.17 1.31 1.35 2.31 1.28 2.08 2.00 1.74 1.82 1.64 2.29 2.28 U 0.45 0.46 0.46 0.89 0.51 0.70 0.71 0.59 0.65 0.64 0.86 0.82
139
Table 4-2. Continued sample 265-63 265-85 265-94 264-09 265-70 265-42 265-83 265-95 265-40 Li 30.23 31.42 30.77 26.68 33.56 31.62 32.45 31.37 30.55 Sc 17 14 13 12 12 14 11 10 12 V 121 73 63 46 45 58 32 52 51 Cr 9.79 4.65 1.49 3.82 3.70 3.42 3.40 3.01 4.17 Co 15 12 11 10 10 11 8 8 10 Ni 8 7 5 6 5 6 5 5 5 Cu 17 19 17 16 16 15 14 15 17 Zn 113 108 105 89 106 119 103 98 103 Ga 31 28 30 28 29 30 29 30 30 Rb 9.5 13.8 12.4 12.9 15.0 13.7 15.5 12.5 12.4 Sr 76 81 68 78 78 83 76 61 73 Y 132 154 146 151 160 160 159 145 146 Zr 745 872 1050 824 934 816 922 985 1013 Nb 13.18 15.27 16.15 14.89 15.90 16.36 15.57 16.53 16.60 Cs 0.10 0.15 0.13 0.13 0.17 0.16 0.17 0.13 0.13 Ba 49.78 68.07 60.04 65.65 72.69 70.01 76.40 62.17 59.70 La 23.53 29.07 29.10 28.98 30.89 29.03 30.69 29.16 29.47 Ce 68.11 82.46 83.93 83.89 88.15 82.95 87.16 83.65 85.00 Pr 10.16 11.78 12.15 11.97 12.49 11.98 12.32 12.02 12.44 Nd 47.3 52.1 55.2 52.7 55.0 53.5 54.0 54.6 56.6 Sm 13.77 16.01 15.74 15.92 16.67 16.69 16.46 15.64 16.89 Eu 2.99 3.01 3.02 2.95 3.09 3.39 3.05 2.89 3.32 Gd 16.37 18.51 18.17 18.54 19.47 19.69 18.90 17.66 19.68 Tb 3.10 3.55 3.39 3.55 3.69 3.76 3.64 3.35 3.67 Dy 20.53 23.80 22.36 23.80 25.06 25.37 24.51 22.13 23.89 Ho 4.40 5.12 4.77 5.12 5.38 5.47 5.27 4.74 5.20 Er 13.19 15.50 14.37 15.66 16.42 16.48 16.18 14.61 15.35 Tm 2.10 2.49 2.29 2.52 2.63 2.64 2.62 2.32 2.49 Yb 13.54 16.69 14.81 16.84 17.54 17.56 17.54 15.12 16.19 Lu 2.06 2.58 2.30 2.61 2.73 2.75 2.72 2.32 2.42 Hf 18.62 22.96 24.87 22.68 24.80 22.24 24.75 24.51 24.97 Ta 0.99 1.05 1.18 1.04 1.10 1.12 1.08 1.23 1.02 Pb 5.82 3.59 5.75 2.76 3.84 3.47 3.80 4.12 3.64 Th 1.65 2.43 2.35 2.59 2.68 2.36 2.80 2.36 2.49 U 0.59 0.92 0.84 0.98 1.04 0.91 1.05 0.86 0.84
140
Table 4-3. West limb major element data 9°N OSC sample Rock Type SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total 267-16 ferrobasalt 50.76 1.18 14.81 9.62 0.19 8.72 12.51 2.06 0.07 0.13 100.1 267-17 FeTi 51.07 2.12 13.54 12.18 0.22 6.74 10.70 2.98 0.11 0.23 99.89 267-18 ferrobasalt 51.15 1.71 14.29 10.63 0.18 7.70 11.18 2.70 0.10 0.17 99.81 267-19 ferrobasalt 50.97 1.83 13.97 11.13 0.20 7.40 10.95 2.78 0.10 0.20 99.53 267-20 FeTi 50.78 2.16 13.45 12.43 0.22 6.66 10.47 2.93 0.12 0.23 99.44 267-21 ferrobasalt 51.23 1.72 14.26 10.64 0.20 7.77 11.14 2.74 0.10 0.18 99.97 267-22 ferrobasalt 51.02 1.80 14.15 10.87 0.20 7.52 10.85 2.84 0.11 0.21 99.57 267-23 basaltic andesite 55.95 2.04 13.02 11.98 0.21 2.84 6.59 3.70 0.67 0.73 97.73 267-24 ferrobasalt 50.74 1.81 13.96 10.78 0.21 7.64 10.85 3.00 0.12 0.21 99.32 267-25 FeTi 50.90 2.06 13.49 12.42 0.23 7.29 10.84 2.59 0.15 0.24 100.2 267-26 ferrobasalt 50.70 1.86 13.73 11.57 0.22 7.21 10.77 2.87 0.16 0.22 99.31 267-27 ferrobasalt 50.69 1.87 13.44 11.44 0.22 7.12 10.60 2.82 0.17 0.25 98.61 267-29 ferrobasalt 51.68 1.61 13.62 11.20 0.21 7.00 10.45 2.80 0.22 0.19 98.97 267-30 basaltic andesite 52.33 1.92 13.28 12.09 0.23 5.43 9.06 3.25 0.37 0.48 98.45 267-32 ferrobasalt 50.57 1.75 14.05 10.61 0.20 7.76 10.89 2.99 0.12 0.20 99.14 267-33 ferrobasalt 50.72 1.78 14.04 10.72 0.20 7.71 10.89 2.99 0.12 0.21 99.39 267-34 ferrobasalt 50.67 1.80 14.00 10.72 0.21 7.73 10.85 2.98 0.12 0.18 99.25 267-35 FeTi 50.49 2.10 13.72 12.00 0.23 6.88 10.83 3.19 0.14 0.22 99.81 267-37 ferrobasalt 50.79 1.88 13.76 11.61 0.21 7.26 11.03 2.90 0.14 0.21 99.78 267-38 ferrobasalt 50.83 1.58 14.28 10.23 0.19 7.67 11.76 2.98 0.08 0.14 99.73 267-39 ferrobasalt 50.78 1.78 14.05 10.80 0.20 7.75 11.05 3.00 0.12 0.20 99.72 267-40 ferrobasalt 50.88 1.76 14.18 10.76 0.21 7.82 10.88 3.00 0.12 0.21 99.83 267-41 basaltic andesite 52.40 3.06 12.67 14.13 0.24 4.21 7.83 3.77 0.64 0.46 99.41 267-42 ferrobasalt 50.69 1.76 14.11 10.74 0.20 7.81 10.99 3.02 0.12 0.20 99.64 267-43 ferrobasalt 50.74 1.76 14.16 10.78 0.20 7.77 10.97 3.03 0.12 0.17 99.70 267-44 ferrobasalt 50.86 1.76 14.06 10.90 0.21 7.54 11.11 2.99 0.12 0.19 99.74 267-45 ferrobasalt 50.77 1.76 13.88 10.84 0.20 7.71 11.01 2.95 0.12 0.19 99.44 267-46 ferrobasalt 50.78 1.78 14.01 10.88 0.20 7.63 11.04 2.98 0.12 0.21 99.63 267-47 ferrobasalt 50.33 1.83 14.12 10.24 0.20 7.28 11.38 2.89 0.32 0.27 98.87 267-48 ferrobasalt 49.95 1.65 14.43 10.66 0.21 8.01 11.73 2.62 0.09 0.21 99.54 267-50 ferrobasalt 49.76 1.58 14.83 10.57 0.20 8.31 11.69 2.63 0.09 0.17 99.82 267-51 ferrobasalt 50.64 1.80 14.11 10.86 0.20 7.55 11.08 2.96 0.12 0.21 99.54 267-52 ferrobasalt 50.20 1.95 13.70 11.64 0.21 7.08 10.97 3.12 0.12 0.22 99.23 267-53 ferrobasalt 50.67 1.57 14.28 10.24 0.20 7.84 11.78 2.73 0.11 0.17 99.58
141
Table 4-3. Continued sample Rock Type SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total 267-54 ferrobasalt 49.98 1.71 14.19 10.68 0.20 7.65 11.52 2.91 0.11 0.18 99.13 267-55 ferrobasalt 50.30 1.54 14.48 10.02 0.20 7.98 11.83 2.71 0.11 0.17 99.33 267-56 FeTi 50.18 2.22 13.05 13.11 0.24 6.42 10.06 3.09 0.15 0.26 98.78 267-57 ferrobasalt 49.82 1.36 14.93 9.35 0.18 8.35 12.37 2.55 0.14 0.16 99.20 267-58 ferrobasalt 49.98 1.82 14.25 10.24 0.21 7.22 11.52 2.88 0.32 0.27 98.70 267-59 FeTi 49.67 2.00 14.29 10.64 0.20 6.99 11.16 3.10 0.36 0.26 98.67 267-60 FeTi 50.14 2.12 13.13 12.54 0.23 6.73 10.35 2.97 0.14 0.24 98.61 267-61 ferrobasalt 49.46 1.56 14.79 10.50 0.20 8.17 11.62 2.62 0.09 0.18 99.20 267-62 ferrobasalt 50.70 1.35 14.64 9.44 0.18 8.34 11.98 2.46 0.08 0.14 99.30 267-63 ferrobasalt 50.13 1.43 13.96 9.89 0.20 8.10 12.11 2.41 0.09 0.16 98.47 267-64 ferrobasalt 49.72 2.40 14.67 10.76 0.19 6.46 10.21 3.19 0.59 0.41 98.59 267-65 ferrobasalt 50.97 1.72 14.00 10.91 0.21 7.50 11.05 2.66 0.12 0.18 99.33 267-66 ferrobasalt 49.91 2.11 13.27 12.34 0.22 6.76 10.15 2.87 0.13 0.23 97.99 267-67 FeTi 50.70 2.05 13.44 12.03 0.23 6.87 10.27 2.91 0.12 0.21 98.84 267-68 ferrobasalt 50.66 2.02 13.58 11.94 0.21 6.95 10.33 2.86 0.12 0.22 98.89 267-69 FeTi 50.80 2.04 13.55 12.01 0.21 6.97 10.35 2.88 0.12 0.21 99.14
142
Table 4-4. West limb trace element data 9°N OSC
sample 267-16
267-18
267-23
267-62
267-63
267-64
267-68
267-69
Li 5.3 6.9 19.4 5.4 5.6 7.9 8.2 7.7 Sc 40 41 26 40 39 41 42 41 V 275 312 148 278 266 345 368 350 Cr 368 147 24 359 339 240 80 71 Co 43 40 21 40 38 43 43 42 Ni 117 64 15 72 66 113 48 47 Cu 146 64 27 78 73 66 56 56 Zn 73 86 117 75 73 96 100 96 Ga 15 17 25 15 15 25 19 19 Rb 0.52 1.02 8.91 0.89 0.98 11.13 1.01 1.15 Sr 87 109 102 96 96 327 116 106 Y 27 37 127 29 29 43 47 43 Zr 62 112 669 82 82 208 133 134 Nb 1.55 2.48 19.92 2.00 2.16 21.36 3.22 3.13 Cs 0.01 0.02 0.11 0.01 0.02 0.12 0.01 0.02 Ba 3.39 7.24 70.75 7.07 7.31 131.32 8.56 9.24 La 2.38 3.65 23.42 2.86 2.90 15.43 4.80 4.52 Ce 7.21 11.68 66.80 8.95 9.02 36.58 14.93 14.46 Pr 1.21 2.08 10.13 1.54 1.59 4.96 2.47 2.47 Nd 6.98 11.05 48.44 8.39 8.63 22.53 13.42 13.23 Sm 2.44 3.86 14.98 2.95 3.03 5.96 4.72 4.58 Eu 0.89 1.34 3.55 1.07 1.11 2.00 1.60 1.55 Gd 3.37 5.08 18.03 3.93 4.07 6.87 6.15 5.89 Tb 0.65 0.93 3.28 0.73 0.75 1.17 1.16 1.10 Dy 4.26 6.08 21.11 4.74 4.90 7.27 7.67 7.14 Ho 0.93 1.30 4.50 1.01 1.06 1.48 1.61 1.50 Er 2.68 3.69 12.96 2.90 2.99 4.18 4.68 4.36 Tm 0.41 0.56 2.02 0.44 0.46 0.62 0.71 0.66 Yb 2.68 3.60 12.92 2.83 2.93 3.93 4.68 4.23 Lu 0.41 0.55 1.98 0.43 0.44 0.60 0.72 0.65 Hf 1.74 2.86 16.00 2.09 2.17 4.70 3.56 3.41 Ta 0.10 0.17 1.11 0.14 0.15 1.37 0.22 0.22 Pb 0.13 0.52 2.02 0.24 0.41 1.86 0.44 0.39 Th 0.09 0.15 1.85 0.12 0.14 1.42 0.18 0.18 U 0.04 0.06 0.54 0.05 0.06 0.42 0.07 0.07
143
Table 4-5. West limb isotopic data 9°N OSC sample 267-18 267-23 267-62 267-63 267-64 267-69 208Pb/204Pb 37.736 37.846 37.748 37.737 38.022 37.738 2 sigma error 0.0018 0.0019 0.0020 0.0022 0.0017 0.0015 207Pb/204Pb 15.478 15.495 15.480 15.476 15.530 15.476 2 sigma error 0.0007 0.0007 0.0008 0.0008 0.0006 0.0006 206Pb/204Pb 18.253 18.368 18.258 18.256 18.590 18.263 2 sigma error 0.0008 0.0008 0.0010 0.0009 0.0006 0.0007 208Pb/206Pb 2.0674 2.0603 2.0674 2.0671 2.04529 2.0664 2 sigma error 0.00004 0.00003 0.00003 0.00003 0.00003 0.00003 207Pb/206Pb 0.8480 0.8435 0.8478 0.847 0.8354 0.8474 2 sigma error 0.00001 0.00001 0.00001 0.00001 0.00001 0.00001 87Sr/86Sr 0.70249 0.70265 0.70254 0.70259 0.70282 0.70254 2 sigma error 0.00001 0.00001 0.00001 0.00002 0.00002 0.00002 143Nd/144Nd 0.513152 0.513139 0.513149 0.51315 0.513049 0.513144 2 sigma error 0.000007 0.000006 0.000004 0.000007 0.000004 0.000004 Eps Nd 10.0 9.8 10.0 10.0 8.0 9.8
144
Table 4-6. Major and trace element data from 8°37'N EPR 3925-0 3925-2 3925-3 3926-1 3926-3 3926-4 3926-5 3927-1 3927-2 SiO2 51.97 50.7 50.64 50.59 55 57.14 50.73 50.47 50.74 TiO2 1.72 1.63 1.63 1.72 1.87 1.64 1.6 1.82 1.73 Al2O3 14.53 14.7 14.63 14.68 14.17 14.14 14.56 14.37 14.49 FeO 10.43 10.49 10.31 10.31 10.95 10.15 10.38 10.94 10.63 MnO 0.20 0.21 0.18 0.19 0.18 0.17 0.17 0.19 0.19 MgO 6.25 7.29 7.24 7.21 3.85 3.08 7.26 6.83 7.07 CaO 10.42 11.59 11.69 11.59 7.36 6.5 11.56 11.1 11.41 Na2O 3.19 2.84 2.83 2.94 4.05 4.25 2.86 3 2.9 K2O 0.33 0.16 0.16 0.25 0.69 0.87 0.17 0.2 0.18 P2O5 0.31 0.29 0.29 0.33 0.46 0.48 0.29 0.32 0.31 SO2 0.24 0.26 0.29 0.25 0.22 0.18 0.26 0.28 0.26 Total 99.61 100.2 99.93 100.1 98.83 98.61 99.86 99.55 99.93 Li 9.08 7.64 6.03 17.72 6.24 6.57 6.53 Sc 40.38 39.55 42.32 25.02 42.47 41.93 41.19 V 309. 288 315 152 314 331 315 Cr 114 179 185 43.99 168 89.65 138 Co 39.47 38.39 41.32 23.96 41.38 40.96 39.94 Ni 48.26 49.41 53.77 17.92 51.88 46.03 48.29 Cu 65.95 71.34 76.86 41.13 76.57 71.19 70.79 Zn 90.55 82.35 81.56 110.08 112.00 86.75 82.43 Ga 17.75 16.36 15.78 21.32 15.84 16.52 15.81 Rb 4.92 3.67 2.08 14.11 1.95 3.00 2.40 Sr 140 127 131 119 129 148. 134 Y 53.42 45.73 35.37 117.23 35.66 38.69 36.61 Zr 211 171 110 550 125 122 112 Nb 7.47 5.72 4.19 19.41 3.94 5.85 4.73 Cs 0.06 0.05 0.03 0.16 0.03 0.04 0.03 Ba 38.11 28.50 20.10 99.52 18.91 30.14 23.83 La 8.91 7.22 4.66 23.07 4.47 5.91 5.02 Ce 24.28 19.91 13.20 61.72 12.77 16.24 14.11 Pr 3.69 3.04 2.12 8.97 2.05 2.51 2.25 Nd 17.93 15.03 10.88 41.80 10.72 12.67 11.54 Sm 5.66 4.85 3.69 12.61 3.63 4.14 3.85 Eu 1.71 1.50 1.33 3.01 1.29 1.45 1.38 Gd 6.97 6.09 4.81 14.87 4.69 5.25 4.96 Tb 1.31 1.13 0.89 2.74 0.87 0.96 0.92 Dy 8.44 7.38 5.75 17.82 5.68 6.29 5.98 Ho 1.79 1.57 1.21 3.79 1.21 1.32 1.25 Er 5.22 4.58 3.49 11.08 3.48 3.81 3.61 Tm 0.80 0.70 0.53 1.75 0.52 0.57 0.54 Yb 5.23 4.51 3.43 11.48 3.37 3.69 3.48 Lu 0.80 0.69 0.52 1.76 0.52 0.56 0.53 Hf 5.37 4.48 2.91 13.78 3.23 3.15 2.95 Ta 0.46 0.36 0.26 1.16 0.25 0.37 0.30 Pb 0.87 0.77 0.59 2.16 1.50 0.68 0.61 Th 0.75 0.59 0.30 2.15 0.29 0.42 0.34 U 0.26 0.21 0.10 0.74 0.10 0.14 0.12
145
Table 4-7. Radiogenic isotope ratios 8°37'N EPR Sample 3925-R0 3925-R2 3925-R3 3926_R4 3926-R5 3927-R1 3927-R2 208Pb/204Pb 37.957 37.938 37.976 37.911 37.965 37.992 37.983 2 sigma error 0.0020 0.0019 0.0036 0.0020 0.0017 0.0017 0.0023 207Pb/204Pb 15.510 15.508 15.513 15.505 15.514 15.515 15.517 2 sigma error 0.0008 0.0008 0.0015 0.0008 0.0007 0.0006 0.0008 206Pb/204Pb 18.468 18.445 18.484 18.437 18.463 18.4960 18.482 2 sigma error 0.0008 0.0008 0.0016 0.0008 0.0008 0.0007 0.0009 208Pb/206Pb 2.0553 2.0568 2.0546 2.0562 2.0562 2.0541 2.0551 2 sigma error 0.00004 0.00004 0.00005 0.00003 0.00003 0.00003 0.00005 207Pb/206Pb 0.83984 0.84077 0.83928 0.84093 0.84025 0.83881 0.83955 2 sigma error 0.00001 0.00001 0.00001 0.00001 0.00001 0.00001 0.00001 87Sr/86Sr 0.70262 0.70267 0.70257 0.70265 0.70258 0.70265 0.70260 2 sigma error 0.00001 0.00001 0.00002 0.00001 0.00002 0.00001 0.00001 143Nd/144Nd 0.51314 0.51315 0.51312 0.51313 0.51314 0.51314 0.51316 2 sigma error 0.00001 0.00001 0.00001 0.00001 0.00001 0.00001 0.00001 Eps Nd 9.7 9.9 9.5 9.6 9.8 9.9 10.2
146
Figure 4-1. Bathymetric map of the northern EPR, including the location of 9°N, 8°37, the Clipperton Transform and the Siqueiros Transform. (Data from GeoMapApp; Carbotte et al., 2004). View looking north.
147
Figure 4-2. Bathymetric map of the 9°N OSC with 50m contours. Circles show the location of rock samples collected using the Jason2 ROV during the MEDUSA2007 cruise, with warmer colors representing higher silica contents. Colored outlines delineate regions discussed in the text.
148
Figure 4-3. Bathymetric map of the 9°N OSC with 50m contours. The melt sills underlying the east and west limbs of the OSC are shaded in gray (Kent et al., 2000). Black bar represents the approximate extent of dacites.
149
Figure 4-4. Side scan sonar mosaic from data collected on the MEDUSA2007 cruise using DSL-120A (White et al., 2009). Circles show the location of rock samples collected using the Jason2 ROV during the MEDUSA2007 cruise, with warmer colors representing higher silica contents. Yellow bar represents the approximate extent of the axial summit trough (AST) in the region. South of the AST, the neo-volcanic zone is difficult to identify.
150
Figure 4-5. FeO versus MgO for glasses collected from the east limb of the 9°N OSC. Shaded regions discriminate between different rock types, which are dominated by different petrologic processes. Black lines with x’s represent various bulk mixes of high-silica dacites and lower silica end-members. Red and blue lines with crosses show liquid-lines of descent (calculated using MELTS; Ghiorso and Sack, 1995) of two different oxygen fugacities (blue = QFM-1 and red = QFM). The majority of lavas erupted at the OSC can be explained by fractional crystallization or mixing various proportions of a high-silica and basaltic end-member.
151
Figure 4-6. Major element variations versus MgO (wt%) for glasses collected from the east limb of the 9°N OSC. Compositions are compared to two fractional crystallization trends (calculated using MELTS; Ghiorso and Sack, 1995) with the same parental composition but different oxygen fugacities. Black lines with x’s represent bulk mixing trends. The majority of east limb lavas can be explained by either by fractional crystallization or by mixing of an evolved and dacitic end-members.
A. B.
C.
D.
E.
F.
152
Figure 4-7. P2O5/TiO2 versus MgO (wt%) for east limb glasses. The high-P2O5 andesites lie along calculated fractional crystallization trends, while the low-P2O5 andesites and dacites have lower P2O5/TiO2 ratios for a given MgO. Many of the basaltic andesites can be explained by mixing of various end-members.
153
Figure 4-8. Trace element concentrations versus Zr for glasses collected from the east limb of the OSC. Not all concentrations can be explained by fractional crystallization alone. Mixing of high-silica lavas with various ferrobasalts can explain a wide range of compositions erupting at the OSC.
A. B.
C. D.
E. F.
154
Figure 4-9. Incompatible trace element ratios versus Zr for glasses erupted at the OSC. Low U/Nb ratios in high-P2O5 andesites can be explained by fractional crystallization, however, Zr/Nb ratios indicate another process, such as mixing or assimilation, must be involved.
A. B.
C. D.
155
A
Figure 4-10. Radiogenic isotope ratios showing the variation in sources along the northern EPR, from 9°50 N to the Siquerous Transform Fault. A) Pb/Pb data of lavas collected south of the OSC, generally have more radiogenic signatures than lavas collected to the north of the OSC. N-MORB from both the east and west limb are similar to other N-MORB lavas erupted along the EPR but the west limb are slightly more radiogenic. The west limb has also erupted E-MORB lavas. B) Epsilon Nd versus Sr isotopes, showing the variation in sources erupting at the EPR.
156
B
Figure 4-10. Continued.
157
A
Figure 4-11. Major element concentrations and ratios versus MgO comparing the east and west limb of the OSC. Samples collected from the EPR north of the OSC (black crosses) and during the CHEPR cruise near the OSC and 8°37’(green symbols) are shown for comparison. A) Major element variations versus MgO. The east limb of the OSC has erupted a much wider compositional range compared to the west limb but the west limb is, on average, more primitive. Red line with crosses represents a calculated fractional crystallization trend using a primitive EPR lava as a parent. B) Major element ratios versus MgO. The west limb has erupted basaltic compositions with higher K2O/TiO2 than the east limb, consistent with E-MORB compositions.
158
B
Figure 4-11. Continued.
159
Figure 4-12. Trace element concentrations versus Zr comparing east and west limb basalts and basaltic andesites. The west limb lavas are more primitive relative to the east limb lavas.
A B
C D
E
F
160
Figure 4-13. Primitive mantle normalized diagram showing variations in andesites and basaltic andesites erupted at the OSC. Red line is a MORB from the east limb. Blue lines are the high-P2O5 andesites from the east limb. Green line is the andesite erupted on the west limb. Gray lines are andesites and basaltic andesites erupted on the east limb. The west limb andesite and the high-P2O5 andesite lack the distinct negative Nb, Ta anomaly and do not have as high U and Th compared to the other east limb lavas.
161
CHAPTER 5 CONCLUSIONS
The majority of eruptions at spreading centers produce basalts with relatively
limited chemical variability; however, compositions ranging from basalts to dacites have
been sampled at ridge segment ends. We have documented the eruption of high –silica
lavas on the propagating eastern limb of the 9°N overlapping spreading center (OSC)
on the East Pacific Rise. The dacites, which have erupted on several other ridges,
appear to represent an end-member composition that shows similar major element
trends and incompatible trace element enrichments, suggesting similar processes
controlled their petrogenesis.
The formation of highly evolved lavas on MOR requires a combination of partial
melting, assimilation and crystal fractionation. The highly enriched incompatible trace
element signatures cannot be produced through crystal fractionation alone and appears
to require partial melting of altered ocean crust. EC-AFC modeling suggests significant
amounts (>75%) crystallization of a MORB parent magma and modest amounts (5-
20%) of assimilation of hydrothermally altered ocean crust can produce geochemical
signatures consistent with dacite compositions. The AFC process explains trace
element abundances in high-silica lavas and accounts for several major and minor
element concentrations (i.e. Al2O3, K2O and Cl).
The formation of dacitic lavas on MOR appears to require a unique tectono-
magmatic setting, where episodic magma supply allows for extensive crystal
fractionation, partial melting and assimilation of altered crustal material. These
conditions are met in regions of ridge propagation, such as OSC and propagating ridge
tips, where down axis diking allows for episodic injection of magma into older, altered
162
ocean crust. Here, the magma undergoes extensive crystallization without repeated
replenishment, creating enough latent heat of crystallization to melt and assimilate
surrounding wall rock.
Variations in volatile concentrations and δ18O in 9°N OSC lavas also suggest that
the OSC magmas have experienced assimilation during their petrogenesis, with the
most extreme signatures observed in high-silica andesites and dacites and little
evidence in basaltic lavas. H2O concentrations are up to two times higher in dacitic
lavas compared to calculated fractional crystallization trends, whereas Cl has excesses
of seven to ten times predicted values. δ18O values are on average ~1‰ lower than
ratios expected from fractional crystallization of ferromagnesian silicates and Fe-Ti
oxide phases, consistent with assimilation of an additional component or components.
The source of the excess H2O and Cl and low δ18O values is partially melted,
hydrothermally altered oceanic crust. Vapor saturation pressures calculated from H2O-
CO2 data suggest that assimilation most likely occurs at the top of the melt lens, which
at the 9°N OSC, corresponds approximately to the base of the sheeted dikes.
The distribution of evolved lavas and E-MORB lavas across the OSC is not
symmetric, suggesting that the 2nd order discontinuity represents a division in the
magmatic plumbing system of the EPR. E-MORB lavas are only observed on the dying
western limb and overall the lavas are less evolved than the adjacent eastern limb. We
suggest that the lower magma supply at the west limb allows for the preservation and
eruption of E-MORB compositions, whereas the more robust magmatic system on the
east limb overwhelms this signature. N-MORB lavas on the west limb have more
radiogenic Pb and Sr and less radiogenic Nd compared to east limb N-MORB lavas.
163
Lavas erupted to the south of the OSC (8°37’N) also have more radiogenic Pb and Sr
isotope ratios. This suggests a slightly different mantle is feeding this section of the
EPR and that this large OSC provides a fundamental division between mantle sources
beneath the ridge axis.
164
LIST OF REFERENCES
Agrinier, P., Hekinian, R., Bideau, D. & Javoy, M. (1995). O and H stable isotope compositions of oceanic crust and upper mantle rocks exposed in the Hess Deep near the Galapagos Triple Junction. Earth and Planetary Science Letters 136, 183-196.
Alt, J., Honnorez, J., Laverne, C. & Emmerman, R. (1986). Hydrothermal Alteration of a
1 km Section Through the Upper Oceanic Crust, Deep Sea Drilling Project Hole 504B: Mineralogy, Chemistry, and Evolution of Seawater-Basalt Interations. Journal of Geophysical Research 91, 10309-10335.
Alt, J., Laverne, C., Vanko, D. A., Tartarotti, P., Teagle, D. A. H., Bach, W., Zuleger, E.,
Erzinger, J., Honnorez, J., Pezard, P., Becker, K., Salisbury, M. H. & Wilkens, R. H. (1996). Hydrothermal Alteration of a Section of Upper Oceanic Crust in the Eastern Equatorial Pacific: A synthesis of results from site 504 (DSDP Legs 69, 70, and 83, and ODP Legs 111, 137, 140, and 148). Proceedings of the Ocean Drilling Program, Scientific Results 148, 417-434.
Alt, J. & Teagle, D. A. H. (eds.) (2000). Hydrothermal alteration and fluid fluxes in
ophiolites and oceanic crust. Boulder, CO: Geological Society of America Alt, J. C. & Teagle, D. A. H. (2003). Hydrothermal alteration of upper oceanic crust
formed at a fast-spreading ridge: mineral, chemical, and isotopic evidence from ODP Site 801. Chemical Geology 201, 191-211.
Anderson, A. T., Clayton, R. N. & Mayeda, T. K. (1971). Oxygen isotope thermometry of
mafic igneous rocks. Journal of Geophysical Research 85, 715-729. Aumento, F. (1969). Diorites from the mid-Atlantic ridge at 45°N. Science 165, 1112-
1113. Bachmann, O. & Bergantz, G. W. (2004). On the Origin of Crystal-poor Rhyolites:
Extracted from Batholithic Crystal Mushes. Journal of Petrology 45, 1565-1582. Barker, A. K., Coogan, L. A. & Gillis, K. M. (2008). Strontium isotope constraints on fluid
flow in the sheeted dike complex of fast spreading crust: Pervasive fluid flow at Pito Deep. Geochemistry Geophysics Geosystems 9, 1019.
Batiza, R. & Niu, Y. (1992). Petrology and Magma Chamber Processes at the East
Pacific Rise ~ 9°30'N. Journal of Geophysical Research 97, 6779-6797. Bazin, S., Harding, A. J., Kent, G. M., Orcutt, J. A., Tong, C. H., Pye, J. W., Singh, S.
C., Barton, P. J., Sinha, M. C., White, R. S., Hobbs, R. W. & Van Avendonk, H. J. A. (2001). Three-dimensional shallow crustal emplacement at the 903' N overlapping spreading center on the East Pacific Rise: Correlations between
165
magnetization and tomographic images. Journal of Geophysical Research 106, 16,101-116,117.
Beard, J. S. & Lofgren, G. E. (1991). Dehydration melting and water-saturated melting
of basaltic and andesitic greenstones and amphibolites at 1, 3, and 6.9 kb. Journal of Petrology 32, 365-401.
Beccaluva, L., Chinchilla-Chaves, A. L., Coltorti, M., Giunta, G., Siena, F. & Vaccaro, C.
(1999). Petrological and structural significance of the Santa Elena-Nicoya ophiolitic complex in Costa Rica and geodynamic implications. Contributions to Mineralogy and Petrology 11, 1091-1107.
Bédard, J. H., Hébert, R. & Berclaz Alain, V., Veronika (eds.) (2000). Syntexis and the
genesis of lower oceanic crust. Boulder, CO: Geological Society of America Berndt, M. E. & Seyfreid, W. E. J. (1990). Boron, bromine, and other trace elements as
clues to the fate of chlorine in mid-ocean ridge vent fluids. Geochemica et Cosmochimica Acta 54, 2235-2245.
Bohnenstiehl, D. R., Tolstoy, M., Fox, C. G., Dziak, R. P., Chapp, E., Fowler, M., Haxel,
J., Fisher, C., Van Hilst, C., Laird, R., Collier, R., Cowen, J. P., Lilley, M., Simons, K., Carbotte, S. M., Reynolds, J., R. & Langmuir, C. H. (2003). Anomalous Seismic Activity at 8°37-42'N on the East Pacific Rise: Hydroacoustic Detection and Site Investigation. Ridge 2000 Newsletter 1, 18-20.
Bohrson, W. A. & Reid, M. R. (1997). Genesis of Silicic Peralkaline Volcanic Rocks in
an Ocean Island Setting by Crustal Melting and Open-system Processes: Socorro Island, Mexico. Journal of Petrology 38, 1137-1166.
Bohrson, W. A. & Reid, M. R. (1998). Genesis of Evolved Ocean Island Magmas by
Deep- and Shallow-Level Basement Recycling, Socorro Island, Mexico: Constraints from Th and other Isotope Signatures. Journal of Petrology 39, 995-1008.
Bohrson, W. A. & Spera, F. (2001). Energy-Constrained Open-System Magmatic
Processes II: Application of Energy Constrained Assimilation-Fractional Crystallization (EC-AFC) Model to Magmatic Systems. Journal of Petrology 42, 1019-1041.
Bowen, N. L. (1928). The evolution of igneous rocks: Princeton University Press. Brophy, J. G. (2009). La-SiO2 and Yb-SiO2 systematics in mid-ocean ridge magmas:
implications for the origin of oceanic plagiogranite. Contributions to Mineralogy and Petrology 158, 99-111.
166
Bryan, W. B. & Moore, J. G. (1977). Compositional variations of young basalts in the Mid-Atlantic Ridge rift valley near 36°49' N. Geological Society of America Bulletin 88, 556-570.
Byerly, G. R. (1980). The nature of differentiation trends in some volcanic rocks from the
Galapagos Spreading Center. Journal of Geophysics 85, 3797-3810. Byerly, G. R. & Melson, W. G. (1976). Rhyodacites, andesites, ferro-basalts and ocean
tholeiites from the Galapagos spreading center. Earth and Planetary Science Letters 30, 215-221.
Canales, P. J., Detrick, R. S., Carbotte, S. M., Kent, G. M., Diebold, J. B., Harding, A. J.,
Babcock, J., Nedimovie, M. R. & van Ark, E. (2005). Upper crustal structure and axial topography at intermediate spreading ridges: Seismic constraints from the southern Juan de Fuca Ridge. Journal of Geophysical Research 110, 1-27.
Carbotte, S. M., Arko, R., Chayes, D. N., Haxby, W., Lehnert, K., O'Hara, S., Ryan, W.
B. F., Weissel, R. A., Shipley, T., Gahagan, L., Johnson, K. T. M. & Shank, T. M. (2004a). New Integrated Data Management System for Ridge2000 and MARGINS Research. EOS, Transactions American Geophysical Union 85(51), 553.
Carbotte, S. M. & Macdonald, K. C. (1992). East Pacific Rise 8°-10°30'N: Evolution of
ridge segments and discontinuities from SeaMARC II and three-dimensional magnetic studies. Journal of Geophysical Research 97, 6959-6982.
Carbotte, S. M., Small, C. & Donnelly, K. (2004b). The influence of ridge migration on
the magmatic segmentation of mid-ocean ridges. Nature 429, 743-746. Casey, J. F. (ed.) (1997). Comparison of major- and trace-element geochemistry of
abyssal peridotites and mafic plutonic rocks with basalts from the MARK region of the mid-Atlantic ridge. College Station: Ocean Drilling Program.
Chadwick, J., Perfit, M. R., Ridley, W. I., Jonasson, I. R., Kamenov, G., Chadwick, W.
W. J., Embley, R. W., le Roux, P. J. & Smith, M. C. (2005). Magmatic effects of the Cobb hot spot on the Juan de Fuca Ridge. Journal of Geophysical Research 110, 1-16.
Christie, D. M. & Sinton, J. M. (1981). Evolution of abyssal lavas along propagating
segments of the Galapagos Spreading Center. Earth and Planetary Science Letters 56, 321.
Clague, D. A. & Bunch. (1976). Formation of ferrobasalt at East Pacific mid ocean
spreading centers. Journal of Geophysical Research 81, 4247-4256.
167
Coleman, R. G. & Donato, M. M. (1979). Oceanic plagiogranite revisited: Elsevier, Amsterdam.
Coogan, L. A. (2003). Contaminating the lower crust in the Oman ophiolite. Geology 31,
1065-1068. Coogan, L. A., Gillis, K. M., MacLeod, C. J., Thompson, G. M. & Hekinian, R. (2002).
Petrology and geochemistry of the lower ocean crust formed at the East Pacific Rise and exposed at Hess Deep: A synthesis and new results. Geochemistry Geophysics Geosystems 3, 1-30.
Coogan, L. A., Banks, G. J., Gillis, K. M., MacLeod, C. J. & Pearce, J. A. (2003a).
Hidden melting signatures recorded in the Troodos ophiolite plutonic suite: evidence for widespread generation of depleted melts and intra-crustal melt aggregation. Contributions to Mineralogy and Petrology 144, 484-505.
Coogan, L. A., Mitchell, N. C. & O'Hara, M. J. (2003b). Roof assimilation at fast
spreading ridges: An investigation combining geochemical, and field evidence. Journal of Geophysical Research 108, 1-14.
Cotsonika, L. (2006). Petrogenesis of Andesites and Dacites from the Southern Juan
De Fuca Ridge. Department of Geological Sciences. Gainesville: University of Florida, 188.
Danyushevsky, L. V. (2001). The effect of small amounts of H2O on crystallization of
mid-ocean ridge and backarc basin magmas. Journal of Volcanology and Geothermal Research 110, 265-280.
De Paolo, D. J. (1981). Trace element and isotopic effects of combined wallrock
assimilation and fractional crystallization. Earth and Planetary Science Letters 53, 189-202.
Delacour, A., Fruh-Green, G. L., Frank, M., Gutjahr, M. & Kelley, D. S. (2008). Sr- and
Nd-isotope geochemistry of the Atlantis Massif (30°N, MAR): Implications for fluid fluxes and lithospheric heterogeneity. Chemical Geology 254, 19-35.
Detrick, R. S., Buhl, P., Vera, E., Mutter, J., Orcutt, J. A., Madsen, J. & Brocher, T.
(1987). Multichannel seismic imaging of a crustal magma chamber along the East Pacific Rise. Nature 326, 35-41.
Dick, J. B., Natland, J. H., Alt, J., Bach, W., Bideau, D., Gee, J. S., Haggas, S.,
Hertogen, J. G. H., Hirth, G., Holm, P. M., Ildefonse, B., Iturrino, G. J., John, B. E., Kelley, D. S., Kikawa, E., Kingdon, A., le Roux, P. J., Maeda, J., Meyer, P. S., Miller, D. J., Naslund, H. R., Niu, Y., Robinson, P. T., Snow, J., Stephen, R. A., Trimby, P. W., Worm, H. U. & Yoshinobu, A. (2000). A long in situ section of the
168
lower ocean crust: results of ODP Leg 176 drilling at the Southwest Indian Ridge. Earth and Planetary Science Letters 179, 31-51.
Dixon, J. E. & Pan, V. (1995). Determination of the molar absorptivity of dissolved
carbonate in basanitic glass. American Mineralogist 80, 1339-1342. Dixon, J. E., Stolper, E. M. & Delaney, J. R. (1988). Infrared spectroscopic
measurements of CO2 and H2O in Juan de Fuca Ridge basaltic glasses. Earth and Planetary Science Letters 90, 87-104.
Dixon, J. E., Stolper, E. M. & Holloway, J. R. (1995a). An Experimental Study of Water
and Carbon Dioxide Solubilities in Mid-Ocean Ridge Basaltic Liquids. Part I: Calibration and Solubility Models. Journal of Petrology 36, 1607-1631.
Dixon, J. E., Stolper, E. M. & Holloway, J. R. (1995b). An Experimental Study of Water
and Carbon Dioxide Solubilities in Mid-Ocean Ridge Basaltic Liquids. Part II: Applications to Degassing. Journal of Petrology 36, 1633-1646.
Donelly, K. E., Goldstein, S. J., Langmuir, C. H. & Spiegelman, M. (2004). Origin of
enriched ocean ridge basalt and implications for mantle dynamics. Earth and Planetary Science Letters 226, 347-366.
Dunn, R. A., Toomey, D. R., Detrick, R. S. & Wilcock, W. S. D. (2001). Continuous
Mantle Melt Supply beneath an Overlapping Spreading Center on the East Pacific Rise. Science 291, 1955-1958.
Eiler, J. M. (2001). Oxygen Isotope Variations of Basaltic Lavas and Upper Mantle
Rocks. Reviews in Mineraology and Geochemistry 43, 319-364. Eiler, J. M., Farley, K. A., Valley, J., Hofmann, A. W. & Stolper, E. M. (1996). Oxygen
isotope constraints on the sources of Hawaiian volcanism. Earth and Planetary Science Letters 144, 453-468.
Embley, R. W., Chadwick, W. W. J., Clague, D. A. & Stakes, D. S. (1999). 1998
eruption of Axial Volcano: Multibeam anomalies and seafloor observations. Geophysical Research Letters 26, 3425-3428.
Embley, R. W., Chadwick, W. W. J., Perfit, M. R. & Baker, E. T. (1991). Geology of the
northern Cleft segment, Juan de Fuca Ridge: Recent lava flows, sea-floor spreading, and the formation of megaplumes. Geology 19, 771-775.
Embley, R. W., Jonasson, I. R., Perfit, M. R., Franklin, J. M., Tivey, M. A., Malahoff, A.,
Smith, M. F. & Francis, T. J. G. (1988). Submersible investigations of an extinct hydrothermal system on the Galapagos Ridge: Sulfide mound, stockwork zone, and differentiated lavas. Canadian Mineralogy 26, 1-35.
169
Embley, R. W. & Wilson, D. (1992). Morphology of the Blanco Transform Fault zone-NE Pacific: Implications for its tectonic evolution. Marine Geophysical Research 14, 25-45.
Fornari, D. J. (2003). A new deep-sea towed digital camera and multi-rock coring
system. EOS, Transactions American Geophysical Union 84, 69-76. Fornari, D. J., Perfit, M. R., Malahoff, A. & Embley, R. W. (1983). Geochemical studies
of abyssal lavas recovered by DSRV Alvin from the Eastern Galapagos Rift, Inca Transform, and Ecuador Rift 1. Major element variation in natural glasses and spatial distribution of lavas. Journal of Geophysical Research 88, 10159-10529.
Fundis, A., Soule, A. S., Fornari, D. J. & Perfit, M. R. (2010). Paving the seafloor:
volcanic emplacement processes during the 2005-06 eruption at teh fast-spreading East Pacific Rise, 9°50'N. Geochemistry Geophysics Geosystems in press.
Garcia, M. O., Muenow, D. W., Aggrey, K. E. & O'Neil, J. R. (1998). Major element,
volatile and stable isotope geochemistry of Hawaiian submarine tholeiitic glasses. Journal of Geophysical Research 94, 10,525-510,538.
Gautason, B. & Muehlenbach, K. (1998). Oxygen isotopic fluxes associated with high-
temperature processes in the rift zones of Iceland. Chemical Geology 145, 275-286.
Ghiorso, M. S. & Sack, R. O. (1995). Chemical Mass Transfer in Magmatic Processes.
IV. A Revised and Internally Consistent Thermodynamic Model for the Interpolation and Extrapolation of Liquid-Solid Equilibria in Magmatic Systems at Elevated Temperatures and Pressures. Contributions to Mineralogy and Petrology 119, 197-212.
Gillis, K. M. (2008). The roof of an axial magma chamber: A hornfelsic heat exchanger.
Geology 36, 299-302. Gillis, K. M. & Coogan, L. A. (2002). Anatectic Migmatites from the Roof of an Ocean
Ridge Magma Chamber. Journal of Petrology 43, 2075-2095. Gillis, K. M., Coogan, L. A. & Chaussidon, M. (2003). Volatile element (B, Cl, F)
behaviour in the roof of an axial magma chamber from the East Pacific Rise. Earth and Planetary Science Letters 213, 447-462.
Gillis, K. M., Muehlenbach, K., Stewart, M., Gleeson, T. & Karson, J. A. (2001). Fluid
flow patterns in fast spreading East Pacific Rise crust exposed at Hess Deep. Journal of Geophysical Research 106, 26311-26329.
170
Gillis, K. M. & Roberts, M. D. (1999). Cracking at the magma-hydrothermal transition: Evidence from the Troodos ophiolite, Cyprus. Earth and Planetary Science Letters 169, 227-244.
Goss, A., Perfit, M., Ridley, W. I., Rubin, K. H., Kamenov, G., Soule, A. S., Fundis, A. &
Fornari, D. J. (2010). Geochemistry of lavas from the 2005-2006 eruption at the East Pacific Rise, 9°46'N-9°56'N: Implications for ridge crest plumbing and decadal changes in magma chamber compositions. Geochemistry Geophysics Geosystems 11, 1-35.
Grimes, C. B., Ushikubo, T., John, B. E. & Valley, J. (2010a). Uniformly mantle-like
δ18O in zircons from oceanic plagioranites and gabbros. Contributions to Mineralogy and Petrology in press, 1-21.
Grimes, C. B., Ushikubo, T. & Valley, J. W. (2010b). Low δ18O (Zrc) in plagiogranites at
Oman: evidence for remelting. Geochemica et Cosmochimica Acta 74, A355. Grove, T. L., Kinzler, R. J. & Bryan, W. B. (1992). Fractionation of mid-ocean ridge
basalts. Geophysical Monograph 71, 281-310. Haase, K. M., Stroncik, N. A., Hékinian, R. & Stoffers, P. (2005). Nb-depleted andesites
from the Pacific-Antarctic Rise as analogs for early continental crust. Geology 33, 921-924.
Hacker, B. R. (1990). Amphibolite-facies to granulite-facies reactions in experimentally
deformed, unpowered amphibolite. American Mineralogist 75, 1349-1361. Hanson, G. N. (1977). Geochemical evolution of the suboceanic mantle. Journal of
Geological Society of London 134, 235-260. Harding, A. J., Kent, A. J. R. & Orcutt, J. A. (1993). A multichannel seismic investigation
of upper crustal structure at 9°N on the East Pacific Rise: Implications for crustal accretion. Journal of Geophysical Research 98, 13925-13944.
Hart, S. R., Schilling, J. G. & Powell, J. L. (1973). Basalts from Iceland and along
Reykjanes Ridge -- Sr isotope geochemistry. Nature 246, 104-107. Hey, R. N., Duennebier, F. K. & Morgan, J. W. (1980). Propagating rifts on mid ocean
ridges. Journal of Geophysical Research 85, 3647-3658. Hey, R. N., Johnson, G. L. & Lowrie, A. (1977). Recent tectonic evolution of the
Galapagos area and plate motions in the East Pacific. Geological Society of America Bulletin 88, 1385-1403.
171
Johannes, W. & Koepke, J. (2001). Incomplete reaction of plagioclase in experimental dehydration melting of amphibolite. Australian Journal of Earth Science 48, 581-590.
Johnson, E. R., Wallace, P., Delgado Granados, H., Manea, V. C., Kent, A. J. R.,
Bindeman, I. N. & Donegan, C. S. (2009). Subduction-rlated Volatile Recycling and Magma Generation beneath Central Mexico: Insights from Melt Inclusions, Oxygen Isotopes and Geodynamic Models. Journal of Petrology 50, 1729-1764.
Juster, T. C., Grove, T. L. & Perfit, M. R. (1989). Experimental Constraints on the
Generation of FeTi Basalts, Andesites, and Rhyodacites at the Galapagos Spreading Center, 85°W and 95°W. Journal of Geophysical Research 94, 9251-9274.
Kamenov, G., Saunders, A. D. & Hames, W. E. (2007). Mafic Magmas as Sources for
Gold in Middle-Miocene Epithermal Deposits of Northern Great Basin, USA: Evidence from Pb Isotopic Compositions of Native Gold. Economic Geology 102, 1191-1195.
Karsten & Delaney, J. R. (1989). Hot spot-ridge crest convergence in the northeast
Pacific. Journal of Geophysical Research 94, 700-712. Kelley, D. S. & Delaney, J. R. (1987). Two-phase separation and fracturing in mid-
ocean ridge gabbros at temperatures greater than 700°C. Earth and Planetary Science Letters 83, 53-66.
Kent, A. J. R., Harding, A. J. & Orcutt, J. A. (1993). Distribution of magma beneath the
East Pacific Rise near 9°03'N overlapping spreading center from forward modeling of CDP data. Journal of Geophysical Research 98, 13971-13995.
Kent, A. J. R., Norman, M. D., Hutcheon, I. D. & Stolper, E. M. (1999). Assimilation of
seawater-derived components in an oceanic volcano: evidence from matrix glasses and glass inclusions from Loihi seamount, Hawaii. Chemical Geology 156, 299-319.
Kent, A. J. R., Peate, D. W., Newman, S., Stolper, E. M. & Pearce, J. A. (2002).
Chlorine in submarine glasses from the Lau Basin: seawater contamination and constraints on the composition of slab-derived fluids. Earth and Planetary Science Letters 202, 361-377.
Kent, G. M., Singh, S. C., Harding, A. J., Sinha, M. C., Orcutt, J. A., Barton, P. J., White,
R. S., Bazin, S., Hobbs, R. W., Tong, C. H. & Pye, J. W. (2000). Evidence from three-dimensional seismic reflectivity images for enhanced melt supply beneath mid-ocean ridge discontinuities. Nature 406, 614-618.
172
Klein, E. M. (ed.) (2005). Geochemistry of the Igneous Oceanic Crust. Oxford: Elsevier–Pergamon.
Klein, E. M. & Langmuir, C. H. (1987). Global correlations of ocean ridge basalt
chemistry. Journal of Geophysical Research 92, 8089-8115. Koepke, J., Berndt, J., Feig, S. T. & Holtz, F. (2007). The formation of SiO2-rich melts
within the deep oceanic crust by hydrous partial melting of gabbros. Contributions to Mineralogy and Petrology 153, 67-84.
Koepke, J., Feig, S. T., Snow, J. & Freise, M. (2004). Petrogenesis of oceanic
plagiogranites by partial melting of gabbros: an experimental study. Contributions to Mineralogy and Petrology 146, 414-432.
Kurras, G. J., Fornari, D. J., Edwards, M. F., Perfit, M. R. & Smith, M. C. (2000).
Volcanic morphology of the East Pacific Rise crest 9 49'-52'N: Implications for volcanic emplacement processes at fast-spreading mid-ocean ridges. Marine Geophysical Research 21, 23-41.
Kvassnes, A. J. S. & Grove, T. L. (2008). How partial melts of mafic lower crust affect
ascending magmas. Contributions to Mineralogy and Petrology 156, 49-71. Langmuir, C. H. (1989). Geochemical consequences of in situ crystallization. Nature
340, 199-205. Langmuir, C. H., Bender, J. F. & Batiza, R. (1986). Petrological and tectonic
segmentation of the East Pacific Rise, 5°30'-14°30'N. Nature 322, 422-429. Langmuir, C. H., Klein, E. M. & Plank, T. (1992). Petrological Systematics of Mid-Ocean
Ridge Basalts: Constraints on Melt Generation Beneath Ocean Ridges. Geophysical Monograph 71 Mantle Flow and Melt Generation at Mid-Ocean Ridges, 183-280.
le Roux, P. J., Shirey, S. B., Hauri, E. H., Perfit, M. R. & Bender, J. F. (2006). The
effects of variable sources, processes and contaminants on the composition of northern EPR MORB (8-10N and 12-14N): Evidence from volatiles (H20, CO2, S) and Halogens (F, Cl). Earth and Planetary Science Letters 251, 209-231.
Macdonald, K. C. & Fox, P. J. (1983). Overlapping spreading centers: new accretion
geometry on the East Pacific Rise. Nature 302, 55-58. Macdonald, K. C. & Fox, P. J. (1988). Earth and Planetary Science Letters 88, 119.
173
Macdonald, K. C., Fox, P. J., Perram, L. J., Eisen, M. F., Haymon, R. M., Miller, S. P., Carbotte, S. M., Cormier, M.-H. & Shor, A. N. (1988). A new view of the mid-ocean ridge from the behaviour of ridge-axis discontinuities. Nature 335, 217-225.
Macdonald, K. C., Scheirer, D. S. & Carbotte, S. M. (1991). Mid-Ocean Ridges:
Discontinuities, Segments, and Giant Cracks. Science 253, 986-994. Maclennan, J. (2008). The Supply of Heat to Mid-Ocean Ridges by Crystallization and
Cooling of Mantle Melts. Geophysical Monograph 178, 45-73. Mandeville, C. W., Webster, J. D., Rutherford, M. J., Taylor, B. E., Timbal, A. & Faure,
K. (2002). Determination of molar absorptivities for infrared absorption bands of H2O in andesitic glass. American Mineralogist 87, 813-821.
Matsuhisa, Y., Matsubaya, O. & Sakai, H. (1973). Oxygen Isotope Variations in
Magmatic Differentiation Processes of the Volcanic Rocks in Japan. Contributions to Mineralogy and Petrology 39, 277-288.
McBirney, A. R. (ed.) (1993). Differentiated rocks of the Galapagos hotspot: Geological
Society of London Special Publications Michael, P. & Cornell, W. C. (1998). Influence of spreading rate and magma supply on
crystallization and assimilation beneath mid-ocean ridges: evidence from chlorine and major element chemistry of mid-ocean ridge basalts. Journal of Geophysical Research 103, 18325-18356.
Michael, P. & Schilling, J. G. (1989). Chlorine in mid-ocean ridge magmas: Evidence for
assimilation of seawater-influenced components. Geochimica Cosmochemica Acta 53, 3131-3143.
Michard, A. & Albarède, F. (1986). The REE content of some hydrothermal fluids.
Chemical Geology 55, 51-60. Muehlenbach, K. & Byerly, G. R. (1982). 18O-Enrichment of Silicic Magmas Caused by
Crystal Fractionation at the Galapagos Spreading Center. Contributions to Mineralogy and Petrology 79, 76-79.
Muehlenbach, K. & Clayton, R. N. (1972). Oxygen isotope studies of fresh and
weathered submarine basalts. Canadian Journal of Earth Science 2, 172-184. Nakamura, K., Kato, Y., Tamaki, K. & Ishii, T. (2007). Geochemistry of hydrothermally
altered basaltic rocks from the Southwest Indian Ridge near the Rodriquez Triple Junction. Marine Geology 239, 125-141.
174
Natland, J. H., Langmuir, C. H., Bender, J. F., Batiza, R. & Hopson, C. (1986). Petrologic systematics in the vicinity of the 9 degrees N nontransform offset, East Pacific Rise. EOS, Transactions American Geophysical Union 67, 1254.
Natland, J. H. & MacDougall, J. D. (1986). Parental abyssal tholeiites and alkali basalts
at the East Pacific Rise near 9 degrees N and the Siqueiros fracture zone. EOS, Transactions American Geophysical Union 67, 410-411.
Newman, S. & Lowenstern, J. B. (2002). VOLATILECALC: a silicate melt–H2O–CO2
solution model written in Visual Basic for excel. Computers and Geosciences 28, 597-604.
Nicholson, H., Condomines, M., Fitton, J. G., Fallick, A. E., Gronvold, K. & Rogers, G.
(1991). Geochemical and Isotopic Evidence for Crustal Assimilation Beneath Krafla, Iceland. Journal of Petrology 32, 1005-1020.
Niu, Y., Gilmore, T., Mackie, S., Greig, A. & Bach, W. (2002). Mineral chemistry, whole-
rock compositions, and petrogenesis of Leg 176 gabbros: data and discussion. College Station: Ocean Drilling Program.
Nunnery, A., Klein, E. M., Perfit, M., White, R. S., Mason, J. L. & Zaino, A. J. (2008).
Correlation of Seafloor Surface Features and Underlying Melt Bodies at the 9° N Overlapping Spreading Center, East Pacific Rise. EOS, Transactions American Geophysical Union 89(53), Fall Meet. Suppl., Abst V53D-03.
O'Hara, M. J. (1977). Geochemical evolution during fractional crystallization of a
periodically refilled magma chamber. Nature 266, 503-507. O'Nions, R. K. & Gronvold, K. (1973). Petrogenetic relationships of acid and basic rocks
in iceland: Sr-isotopes and reare-earth elements in late and postglacial volcanics. Earth and Planetary Science Letters 19, 397-409.
Pedersen, R. B. & Malpas, J. (1984). The Origin of oceanic plagiogranites from the
Karmoy ophiolite, western Norway. Contributions to Mineralogy and Petrology 88, 36-52.
Perfit, M., Ridley, W. I. & Cotsonika, L. (in prep). Juan de Fuca Dacites. Perfit, M. R. & Chadwick, W. W. J. (1998). Magmatism at Mid-Ocean Ridges:
Constraints from Volcanological and Geochemical Investigations. Geophysical Monograph 106 Faulting and Magmatism at Mid-Ocean Ridges, 59-115.
Perfit, M. R. & Fornari, D. J. (1983). Geochemical studies of abyssal lavas recovered by
DSRV Alvin from Eastern Galapagos Rift, Inca Transform, and Ecuador Rift 2. Phase chemistry and crystallization history. Journal of Geophysical Research 88, 10530-10550.
175
Perfit, M. R., Fornari, D. J., Malahoff, A. & Embley, R. W. (1983). Geochemical Studies of Abyssal Lavas Recovered by DSRV Alvin From Eastern Galapagos Rift, Inca Transform, and Ecuador Rift 3. Trace Element Abundances and Petrogenesis. Journal of Geophysical Research 88, 10,551-510,572.
Perfit, M. R., Fornari, D. J., Smith, M. C., Bender, J. F., Langmuir, C. H. & Hayman, N.
W. (1994). Small-scale spatial and temporal variations in mid-icean ridge crest magmatic processes. Geology 22, 375-379.
Perfit, M. R., Ridley, W. I. & Jonasson, I. R. (1999). Geologic, petrologic and
geochemical relationships between magmatism and massive sulfide mineralization along the eastern Galapagos Spreading Center. Reviews in economic geology 8, 75-99.
Perfit, M. R., Schmidt, A. K., Ridley, W. I., Rubin, K. H. & Valley, J. (2007). Petrogenesis
of dacites from the southern Juan de Fuca Ridge. Goldschmidt Conf. Abstracts, Geochimica et Cosmochimica Acta 72, A736.
Pollard, D. D. & Sydin, A. (1984). Propagation and Linkage of Oceanic Ridge
Segments. Journal of Geophysical Research 89, 10,017-010,028. Pollock, M. A., Klein, E. M., Karson, J. A. & Tivey, M. A. (2005). Temporal and spatial
variability in the composition of lavas exposed along the Western Blanco Transform Fault. Geochemistry Geophysics Geosystems 6, 1-17.
Rapp, R. P., Watson, E. B. & Miller, C. F. (1991). Partial melting of amphibolite/eclogite
and the origin of Archean trondhjemites and tonalites. Precambrian Research 51, 1-25.
Regelous, M., Niu, Y., Went, J. I., Batiza, R., Greig, A. & Collerson, K. D. (1999).
Variations in geochemistry of magmatism on the East Pacific Rise at 10º30' N since 800 ka. Earth and Planetary Science Letters 168, 45-63.
Reynolds, J., R. (1995). Segment-scale systematic of mid-ocean Ridge magmatism and
geochemistry. Palisades: Columbia University, 483. Reynolds, J., R. & Langmuir, C. H. (1997). Petrological systematics of the Mid-Atlantic
Ridge south of Kane: Implications for ocean crust formation. Journal of Geophysical Research 102, 14915-14946.
Reynolds, J., R., Langmuir, C. H., Bender, J. F., Kastens, K. A. & Ryan, W. B. F. (1992).
Spatial and temporal variability in the geochemistry of basalts from the East Pacific Rise. Nature 359, 493-499.
176
Rubin, K. H. & Sinton, J. M. (2007). Inferences on mid-ocean ridge thermal and magmatic structure from MORB compositions. Earth and Planetary Science Letters 260, 257-276.
Saal, A. E., Hauri, E. H., Langmuir, C. H. & Perfit, M. (2002). Vapour undersaturation in
primitive mid-ocean-ridge basalt and the volatile content of Earth's upper mantle. Nature 419, 451-455.
Schiffman, P., Zierenber, R., Chadwick, W. W. J., Clague, D. A. & Lowenstern, J. B.
(2010). Contanination of basaltic lava by seawater: Evidence found in a lava pillar from Axial Seamount, Juan de Fuca Ridge. Geochemistry Geophysics Geosystems 11, 1-12.
Schmitt, A. K., Perfit, M., Rubin, K. H., Stockli, D. F., Smith, M. C., Cotsonika, L.,
Zellmer, G. F., Ridley, W. I. & Lovera, O. M. (submitted). Zircon in Evolved Mid-Ocean Ridge Lavas Constrains Fast Cooling Rates of Oceanic Lower Crust. Nature.
Sempere, J.-C. & Macdonald, K. C. (1986). Deep-Tow Studies Of The Overlapping
Spreading Centers at 903' N On The East Pacific Rise. Tectonics 5, 881-900. Sempre, J.-C. & Macdonald, K. C. (1986). Deep-Tow Studies Of The Overlapping
Spreading Centers at 903' N On The East Pacific Rise. Tectonics 5, 881-900. Sigurdsson, H. (1977). Generation of Icelandic rhyolites by melting of plagiogranites in
the oceanic layer. Nature 269, 25-28. Sigurdsson, H. & Sparks, R. S. J. (1981). Petrology of rhyolitic and mixed magma ejecta
from the 1875 eruption of Askja, Iceland. Journal of Petrology 22, 41-84. Sims, K. W. W., Blichert-Toft, J., Fornari, D. J., Perfit, M. R., Goldstein, S. J., Johnson,
P., De Paolo, D. J. & Michael, P. (2003). Abberant Youth: Chemical and isotopic constraints on the young off-axis lavas of the East Pacific Rise. Geochemistry Geophysics Geosystems 4, 8621.
Sims, K. W. W., Goldstein, S. J., Blichert-Toft, J., Perfit, M. R., Kelemen, P. B., Fornari,
D. J., Michael, P., Murrell, M. T., Hart, S. R., DePaolo, D. J., Layne, G., Ball, L., Jull, M. & Bender, J. F. (2002). Chemical and isotopic constraints on the generation and transport of magma beneath the East Pacific Rise. Geochimica et Cosmochimica Acta 66, 3481-3504.
Singh, S. C., Harding, A. J., Kent, G. M., Sinha, M. C., Combier, V., Bazin, S., Tong, C.
H., Pye, J. W., Barton, P. J., Hobbs, R. W., White, R. S. & Orcutt, J. A. (2006). Seismic reflection images of the Moho underlying melt sills at the East Pacific Rise. Nature 442, 1-4.
177
Sinton, J. M., Ford, L. L., Chappell, B. & McCulloch, M. T. (2003). Magma Genesis and Mantle Heterogeneity in the Manus Back-Arc Basin, Paupa New Guinea. Journal of Petrology 44, 159-195.
Sinton, J. M., Smaglik, S. M. & Mahoney, J. J. (1991). Magmatic processes at superfast
spreading mid-ocean ridges: glass compositional variations along the East Pacific Rise 13º-23ºS. Journal of Geophysical Research 96, 6133-6155.
Sinton, J. M., Wilson, D., Christie, D. M., Hey, R. N. & Delaney, J. R. (1983).
Petrological consequences of rift propagation on oceanic spreading ridges. Earth and Planetary Science Letters 62, 193-207.
Smith, M. C., Perfit, M. R., Fornari, D. J., Ridley, W. I., Edwards, M. H., Kurras, G. J. &
Von Damm, K. L. (2001). Magmatic processes and segmentation at a fast spreading mid-ocean ridge: Detailed investigation of an axial discontinuity on the East Pacific Rise crest at 9°37'N. Geochemistry Geophysics Geosystems 2, 1-32.
Smith, M. C., Perfit, M. R. & Jonasson, I. R. (1994). Petrology and geochemsitry of
basalts from the Southern Juan de Fuca Ridge: Controls on the spatial and temporal evolution of mid-ocean ridge basalts. Journal of Geophysical Research 99, 4787-4812.
Soule, A. S., Fornari, D. J., Perfit, M. R., Tivey, M. A., Ridley, W. I. & Schouten, H.
(2005). Channelized lava flows at the East Pacific Rise crest 9-10 N: The importance of off-axis lava transport in developing the architecture of young oceanic crust. Geochemistry Geophysics Geosystems 6, 1-28.
Sparks, J. W. (1995). Geochemistry of the Lower Sheeted Dike Complex, Hole 504B,
Leg 140. Proceedings of the Ocean Drilling Program, Scientific Results 137/140, 81-97.
Stakes, D. S., Perfit, M. R., Tivey, M. A., Caress, D. W., Ramirez, T. M. & Maher, N.
(2006). The Cleft revealed: Geologic, magnetic, and morphologic evidence for construction of upper oceanic crust along the southern Juan de Fuca Ridge. Geochemistry Geophysics Geosystems 7, 1-32.
Sun, S. & McDonough, W. F. (eds.) (1989). Chemical and Isotopic Systematics of
oceanic basalts: implications for Mantle Composition and Processes: Geological Society of London, Special Publications
Sun, S. S., Tatsumoto, M. & Schilling, J. G. (1975). Mantle Plume Mixing Along the
Reykjanes Ridge Axis: Lead Isotopic Evidence. Science 190, 143-147.
178
Sun, W. D., Binns, R. A., Fan, A. C., Kamenetsky, V. S., Wysoczanski, R., Wei, G. J., Hu, Y. H. & Arculus, R. J. (2007). Chlorine in submarine volcanic glasses from the eastern Manus basin. Geochemica et Cosmochimica Acta 71, 1542-1552.
Taylor, H. P. J. (1968). The oxygen isotope geochemistry of igneous rocks.
Contributions to Mineralogy and Petrology 19, 1-71. Taylor, H. P. J. (1980). The effects of assimilation of country rocks by magmas on
18O/16O and 87Sr/86Sr systematics in igneous rocks. Earth and Planetary Science Letters 47, 243.
Thy, P., Lesher, C. E. & Mayfield, J. D. (1999). 9. Low-Pressure Melting Studies of
Basalts and Basaltic Andesites From The Southeast Greenland Continental Margin And The Origin Of Dacites At Site 917. Proceedings of the Ocean Drilling Program, Scientific Results 163, 95-112.
Tong, C. H., Pye, J. W., Barton, P. J., White, R. S., Sinha, M. C., Singh, S. C., Hobbs,
R. W., Bazin, S., Harding, A. J., Kent, G. M. & Orcutt, J. A. (2002). Asymmetric melt sills and upper crustal construction beneath overlapping ridge segments: Implications for the development of melt sills and ridge crests. Geology 30, 83-86.
Toomey, D. R., Jousselin, D., Dunn, R. A., Wilcock, W. S. D. & Detrick, R. S. (2007).
Skew of mantle upwelling beneath the East Pacific Rise governs segmentation. Nature 446, 409-414.
Unni, C. K. & Schilling, J. G. (1978). Cl and Br degassing by volcanism along the
Reykjanes Ridge and Iceland. Nature 272, 19-23. Wanless, D., Perfit, M., Ridley, W. I. & Klein, E. M. (2008). Origin of Dacites on Mid-
Ocean Ridges. EOS, Transactions American Geophysical Union 89(53), Fall Meeting Supplementary, Abstract V41B-2084.
Wanless, D., Perfit, M., Ridley, W. I., Klein, E. M., Wallace, P., Valley, J. & Grimes, C.
(submitted). Role of Assimilation in the Petrogenesis of Lavas on Mid-Ocean Ridges inferred from Cl, H2O, CO2 and Oxygen Isotope Analyses.
Wanless, D., Perfit, M., Ridley, W. I., Klein, E. M. & White, R. S. (in prep). 9°N
Overlapping Spreading Center: Magma Genesis and Ridge Segmentation. Geochemistry Geophysics Geosystems.
Wanless, D., Perfit, M. R., Ridley, W. I., Wallace, P., Valley, J., Grimes, C. & Klein, E.
M. (2009). Geochemical Evidence for Crustal Assimilation at Mid-Ocean Ridges Using Major and Trace Elements, Volatiles and Oxygen Isotopes. EOS, Transactions American Geophysical Union, Fall Meeting Supplementary, Abstract V51A-1665.
179
Wanless, V. D., Perfit, M., Ridley, W. I. & Klein, E. M. (accepted). Dacite Petrogenesis on Mid-Ocean Ridges: Evidence for Crustal Melting and Assimilation. Journal of Petrology Accepted.
Watson, E. B. (1979). Apatite saturation in basic to intermediate magmas. Geophysical
Research Letters 6, 937-940. Webster, J. D., Kinzler, R. J. & Mathez, E. A. (1999). Chloride and water solubility in
basalts and andesite melts and implications for magmatic degassing. Geochemica et Cosmochimica Acta 63, 729-738.
West, M., Menke, W., Tolstoy, M., Webb, S. & Sohn, R. (2001). Magma storage
beneath Axial volcano on the Juan de Fuca mid-ocean ridge. Nature 413, 833-836.
White, S. M., Haymon, R. M., Fornari, D. J., Perfit, M. R. & Macdonald, K. C. (2002).
Correlation between tectonic and volcanic segmentation at fast-spreading ridges: evidence from the distribution of volcanic structures and lava flow morphology along East Pacific Rise at 9° - 10°N. Journal of Geophysical Research 107.
White, S. M., Mason, J. L., Macdonald, K. C., Perfit, M. R., Wanless, V. D. & Klein, E.
M. (2009). Significance of widespread low effusion rate eruptions over the past two million years for delivery of magma to the overlapping spreading centers at 9°N East Pacific Rise. Earth and Planetary Science Letters 280, 175-184.
White, W. M. & Schilling, J. G. (1978). The nature and origin of geochemical variation in
Mid-Atlantic Ridge basalts from the Central North Atlantic. Geochemica et Cosmochimica Acta 42, 1501-1516.
Wilson, D. S., Teagle, D. A. H., Alt, J., Banerjee, N. R., Umino, S., Miyashita, S., Acton,
G. D., Anma, R., Barr, S. R., Belghoul, A., Carlut, J., Christie, D. M., Coggon, R. M., Cooper, K. M., Cordier, C., Crispini, L., Durand, S. R., Einaudi, F., Galli, L., Gao, Y., Geldmacher, J., Gilbert, L. A., Hayman, N. W., Herrero-Bervera, E., Hirano, N., Holter, S., Ingle, S., Jiang, S., Kalberkamp, U., Kerneklian, M., Koepke, J., Laverne, C., Vasquez, H. L. L., Maclennan, J., Morgan, S., Neo, N., Nichols, H. J., Park, S.-H., Reichow, M. K., Sakuyama, T., Sano, T., Sandwell, R., Scheibner, B., Smith-Duque, C. E., Swift, S. A., Tartarottie, P., Tikku, A. A., Tominaga, M., Veloso, E. A., Yamasaki, T., Yamazaki, S. & Ziegler, C. (2006). Drilling to Gabbro in Intact Ocean Crust. Science 312, 1016-1020.
Wolf, M. B. & Wyllie, P. J. (1994). Dehydration-melting of amphibolite at 10 kbar: the
effects of temperature and time. Contributions to Mineralogy and Petrology 115, 369-383.
180
Zaino, A. J. (2009). Petrology and Mineral Chemistry of the 9°03'N Overlapping Spreading Center, East Pacific Rise. Department of Earth and Ocean Sciences. Durham: Duke Univeristy, 60.
181
BIOGRAPHICAL SKETCH
Virginia Dorsey Wanless was born in Denver, Colorado but grew up in Topeka,
Kansas. She earned her Bachelor of Arts in geology from Colgate University in 2001
and her Master of Science from the Department of Geology and Geophysics at the
University of Hawai`i in 2005. Interspersed with her education she worked at the
Hawaiian Volcano Observatory (HVO), FUGRO Sea Floor Surveys, and as a side scan
sonar analyst for the Hawaiian Research Mapping Group (HMRG) and the REMUS
6000 team at Woods Hole Oceanographic Institution (WHOI). After completion of her
Doctor of Philosophy in the summer of 2010, she began a postdoctoral fellowship at
WHOI.