Melt Generation in the Earth's Mantle
at Convergent Plate Margins
by
Christy Berna Till
MASSACHUSETTS INSTITUJTE
OF TECHNOLOGY
B.S. Geological Sciences, 2004
M.S. Geological Sciences, 2005
University of California, Santa Barbara
NOV 2 2 2011
L
Submitted to
the Department of Earth, Atmospheric, and Planetary Sciences
in partial fulfillment of the requirements for the degree of
RARIES
ARCHIVES
DOCTOR OF PHILOSOPHY
in Geology
at the
MASSACHUSETTS INSTITUTE OF TECHNOLOGY
September 2011
0 Massachusetts Institute of Technology
A uth or. .....................
......
.................................................
Dpartment of Earth, Atmospheric, and Planetary Sciences
August 5, 2011
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Certified by................
.......
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4
..............
............
Timothy L. Grove
Professor of Geology
Thesis Supervisor
C ertified by ....................................................
.......................................................................
Maria T. Zuber
E.A. Griswold Professor of Geophysics
Department Head
Melt Generation in the Earth's Mantle
at Convergent Plate Margins
by
Christy Berna Till
Submitted to the Department of Earth, Atmospheric and Planetary Sciences
on August 5th2011,
in partial fulfillment of the requirements for the degree of
Doctor of Philosophy in Geology
Abstract
The five geologic studies presented in this thesis document how the recycling of tectonic
plates at subduction zones has a profound effect on the melting behavior of the Earth's
mantle. Two experimental studies (Chapters 1 and 2) of the melting behavior of mantle
peridotite demonstrate that the forefathers of arc magmas are formed at extremely low
temperatures in the mantle wedge at convergent plate margins in the presence of excess
H2 0 or following the breakdown of the hydrous mineral chlorite. A new petrologic model
that simulates anhydrous melting of variably metasomatized upper mantle is developed to
investigate the petrogenesis of primitive basaltic lavas erupted in continental back-arc
and ocean island settings (Chapter 3). This model suggests that <10.5 Ma anhydrous
basaltic lavas erupted east of the Cascades arc were formed by mantle upwelling near the
lithosphere-asthenosphere boundary caused by plate subduction, not by a mantle plume
(Chapter 4). A geodynamic and petrologic study (Chapter 5) of asthenospheric flow at
the margins of thick continental lithosphere reveals small degree melts of the mantle may
be responsible for the large negative gradient in seismic wave speed observed at the
lithosphere-asthenosphere boundary below eastern North America. These studies
together advance our knowledge of how the recycling of tectonics plates on Earth affects
melt generation in the Earth's mantle and subsequently the unique differentiation of our
planet.
Thesis Supervisor: Timothy L. Gove
Title: Professor of Geology
3
Acknowledgements
"Teiecestat
Y am, [tiey/gatnertiem and7give tiem Aac 1 to me in affrthe rk/it order."
-
Toni Morrison
There are a great many individuals who have supported me so that I can be sitting
here now writing this. First and foremost I would like to acknowledge my advisor and
friend, Professor Tim Grove, who has given me a world of opportunities and knowledge
since we first met. I have had the good fortune to travel around the world with Tim, as
well as to spend countless hours together in the experimental petrology lab in our pursuit
to understand hydrous magma generation at subduction zones. In addition to the
tremendous body of geologic and petrologic knowledge he has passed on to me, I cannot
thank him enough for his unwavering enthusiasm, generosity, and dedication to my
growth as a scientist and citizen of the world. I would also like to thank my thesis
committee, Sam Bowring, Lindy Elkins-Tanton, Oli Jagoutz and Mark Behn who have
all provided insight, scientific and otherwise, throughout my time at MIT. A number of
other collaborators have made important contributions to my research while at MIT
including, Rick Carlson, Othmar Mtntener, Julie Donnelly-Nolan, Marc Hirschmann,
Tony Wither, Tom Sisson, and Karen Fischer. In addition, I wish to thank the EAPS
faculty, administration, and staff, and financial support from the National Science
Foundation Graduate Fellowship program, MIT and EAPS.
I owe a great deal of thanks to the individuals that comprised the Grove lab during
my time at MIT - Neel Chatterjee, Carol Zayotti, Etienne Mddard, Jay Barr, Mike
Krawczynski, Ben Mandler, Anne Pommier, and Bernard Charlier. Any difficult journey
is more rewarding with good friends at your side. In particular, Mike and Jay made the
innumerable experiences we shared as labmates infinitely richer and taught me a
significant amount. One could not ask for a better friend and officemate than Mike, who
was always there no matter the hour, to share the ups and downs and humor of graduate
school and I look forward to having him as a friend and colleague for the rest of my life.
In addition, I am thankful for the friendship of other EAPS graduate students, postdocs
and affiliates including, Scott Burdick, Krystle Catalli, Stacia Gordon, Noah McLean,
Terry Blackburn, Einat Lev, Lindsay Hays, Brent Grocholski, Pierre Bouilhol, and Matt
Rioux.
I would not be here without my family. My parents, Stephen Till and Linda
Butterworth-Till, who starting with their Jimmy Cliff lullabies, taught me to follow my
dreams, no matter how impossible they may seem. My brothers, Henry and Jeremy Till,
inspire me with their creativity and tenacity. And my aunt, Alison Till, an amazing friend
and mentor, opened my eyes to plate tectonics at the age of eleven, and later a career in
geology. A number of other mentors have played important roles in getting me here,
especially Frank Spera, Tanya Atwater, Paul Mejia, Wendy Davis and Neta Barker.
Last and certainly not least, I could not have done this without my partner in life,
Duane DeVecchio. From the roses he gave me in celebration of my acceptance to MIT to
those he gave me after my Ph.D. defense, he has been unwavering in his support and love
for me, including through four long years of an across-country long distance relationship.
I love you.
Table of Contents
A b stract................................................................................2
Acknowledgments....................................................................4
T able of C ontents.....................................................................6
Introduction ..........................................................................
10
Chapter 1. The Beginnings of Hydrous Mantle Wedge Melting...............16
A b stract.........................................................................
.... 17
1. Introduction .......................................................................
18
2. Experimental and Analytical Methods......................................
19
2.1 Starting Materials.................................................
19
2.2 Piston Cylinder Experimental Methods......................... 20
2.3 Multi-Anvil Experimental Methods..............................
21
2.4 Analytical Methods....................................................22
3. Experimental Results..........................................................23
3.1 Mineral Phase Relations & Textures.......
...........23
3.2 Chlorite Peridotite Melting Behavior............................ 26
4. Discussion...........................................
28
4.1 Comparison with Previous Experimental Studies................28
4.2 Chlorite Stability in the Mantle Wedge...........................35
4.3 Implications for Generation of Hydrous Melts
in the Mantle Wedge.................................................
37
R eferences C ited.................................................................
42
Figure & Table Captions...................................
51
F ig ures ...............................................................................
56
Chapter 2. Compositions of Low Extent Melts of H 20-Saturated Peridotite.70
Ab stract....... ............
. ..............................................
. 71
1. Introduction .......................................................................
72
2. Experimental and Analytical Methods......................................
74
3. Determination of Quench Compositions....................................
75
4 . D iscussion ........................................................................
78
4.1 Phase Proportions & Inferred Melting Reactions............... 78
4.2 Comparison to Previous Experimental Studies................. 79
4.3 Melt Productivity....................................................
80
4.4 Comparison of the Experimental Melts to Arc Lavas ........... 80
4.5 Where Do We Go From Here? ............ . . . . . . . . . . . . . . . . . . . . . . . . 81
5. Concluding Remarks............................................................
83
R eferences C ited....................................................................
84
Figure & Table Captions............................................................
88
F igures..........................................................................
... . 90
Chapter 3. Primary Basaltic Magmas From Variably Metasomatized
Mantle: The Effects of Pressure, Temperature, and
Composition on Melting Plagioclase and Spinel Lherzolite......100
101
Abstract................................................
102
I. Introduction ........................................................................
103
II. Experimental Data .........................................................
103
the
Literature............................
From
a. Experimental Data
104
b. New Experiments......................................................
.......... 105
c. Starting Material....................................
106
d. Experimental Methods................................................
107
e. Analytical Methods....................................................
f. Composition of Experimentally Produced Melts
107
Saturated with a Mantle Residue....................................
III. Parameterization of Plagioclase and Spinel Lherzolite
M elting Equilibria..............................................................108
IV. Plagioclase Lherzolite to Spinel Lherzolite Boundary....................112
V. Implementation of the Plagioclase and Spinel Lherzolite Models........ 114
a. Fractional Crystallization Corrections..............................114
b. Calculating Melts of Mantle Lherzolite.... .................... 117
VI. Comparison of Parameterization to Other Thermometers and
118
............ .........................................
B arometers .........
VII. Case Study of Basaltic Lavas from Oregon's HLP...................... 122
127
R eferences C ited...................................................................
133
Figure & Table Captions..........................................................
138
A ppendix 1..........................................................................
14 0
F ig ures ...............................................................................
Chapter 4. Constraints on the Depths and Temperatures of Anhydrous
Asthenospheric Melting and the Location of the
Lithosphere-Asthenosphere Boundary Beneath
Southern Oregon and Northern California..........................
158
A b stract..............................................................................159
16 1
..........................................
1. Introduction .... .....................
2. Tectonics and Volcanic History................................................163
3. M ethods............................................................................168
3.1 Lava Sample Selection...............................................168
3.2 Basalt Thermometry and Barometry...............................169
4. Pressure and Temperature of Mantle Melt Extraction for Primitive
HLP and Cascadian Basalts....................................................171
........... 172
5. Constraints on Crustal and Lithospheric Structure..........
Oregon
6. What is Driving Asthenospheric Melting Below Southern
. . . . . . . . . . . . . . . . . ..
.. ... ... ... .
174
and Northern California?............
179
.......................................
Acknowledgements...
R eferences C ited....................................................................179
187
Figure & Table Captions..........................................................
.......................... 189
F igures................................................
Chapter 5. A Mechanism For Low Extent Melts at the LithosphereAsthenosphere Boundary...............................................194
A bstract...............................................................................195
.196
1. Introduction..........................................
1
2 . M ethods............................................................................20
2.1 Solidi Used in Melting Calculations................................201
2.2 Numerical Models of Mantle Flow and Temperature............208
2.3 Melt Calculations: Criteria for Melt Production.................. 211
13
3. R esu lts..............................................................................2
2 15
4 . D iscussion ..........................................................................
4.1 The LAB beneath eastern North America..........................215
4.2 The Global LAB.......................................................219
4.3 Comparison to Other Melting Models..............................221
4.4 Implications of the Similar Expression for Solidus and
Liquidus Depression with H20......................................223
5. C onclusions........................................................................224
A cknow ledgm ents..................................................................
226
Figure Captions..........................................226
23 1
R eferences C ited .....................................................................
24 0
F ig u re ..................................................................................
9
Introduction
The plate tectonic processes unique to Earth have played an important role in the
chemical differentiation of the planet since early in its history, such that it is the only
terrestrial planet with a highly differentiated continental crust. Consequently, knowledge
of how plate tectonics affects melt generation in the Earth's interior is fundamental to
understanding the differentiation history of Earth. The link between convergent plate
margins and the formation of highly differentiated magmas at volcanic arcs was
recognized as early as 1892, when the "Volcanic Girdle of the Pacific" was first
identified, which later became known as the "Andesite Line" or the "Ring of Fire." But it
was not until 1949, when it was demonstrated that earthquakes are located at
progressively greater depths beneath the volcanic arc, that the presence of a subducting
plate below arc volcanoes was clearly established. Since then, petrologic, geochemical,
geophysical and geodynamic studies reveal thermal and chemical exchange between the
subducting plate and the overlying mantle wedge are responsible for forming a physically
and chemically complex melting regime. Volatiles such as H2 0 and CO 2 liberated from
the subducting plate and the accompanying soluble elements alter the composition and
melting behavior of the overlying mantle, a process known as "metasomatism." Viscous
coupling between the hot mantle and the cool subducting plate induces downward flow at
the base of the wedge and draws mantle into the top half of the wedge, a process known
as "corner flow." Despite these advances in our knowledge, many uncertainties remain
regarding the specifics of melt generation in subduction-modified mantle. This thesis
focuses in advancing our understanding of melt generation in mantle modified by both
modern and historical subduction. It is comprised of five distinct studies that employ
experimental, petrologic, geochemical, geodynamic, and geophysical tools to accomplish
this goal.
The first two chapters are experimental investigations of the high pressure melting
behavior of mantle peridotite in the presence of excess H20. These chapters together
establish that the forefathers of highly differentiated arc magmas are born at the base of
the mantle wedge as a result of the high flux of H20 off the subducting plate. In Chapter
1, we resolve a long-standing controversy regarding the melting temperature, or solidus,
of H20-saturated mantle peridotite and demonstrate that the H20-rich mantle melts at the
extremely low temperature of ~800'C over the pressure range relevant for subarc mantle
melting. This study also suggests the hydrous mineral chlorite is stable in the mantle
wedge, as well the subducting plate. Furthermore, chlorite contains sufficient H2 0 to
trigger H20-undersaturated melting or perpetuate ongoing H20-saturated mantle melting
when it breaks down. When these findings are coupled with thermal and geodynamic
models of subduction zones, we observe that the beginning of hydrous melting in the
mantle wedge directly below arc volcanoes is primarily the result of chlorite breakdown
and the low melting temperature of H20-saturated peridotite. Chapter 2 examines the
composition of the low extent melts of the H20-saturated mantle peridotite generated in a
subset of the experiments in Chapter 1. The compositions of hydrous experimental melts
are notoriously difficult to analyze because of H20-exsolution and the growth of
"quench" minerals when the experiments are depressurized. In this chapter we explore a
variety of techniques to filter chemical analyses of the quenched material and we are able
to determine the compositions of both a high-pressure melt and vapor phase in these
experiments. These compositions are used to constrain the melting reactions and the
rates of melt production applicable to H20-saturated mantle at the base of the mantle
wedge. As these melts ascend through the mantle wedge and crust, modification by
crystallization and assimilation will result in the formation of the magma types erupted at
arc volcanoes.
A significant number of primitive anhydrous basaltic lavas erupted in continental
back-arc settings and at ocean islands exhibit geochemical characteristics that suggest
they are the product of melting previously metasomatized mantle. Chapters 3 and 4 shift
the focus from hydrous to anhydrous melting to explore the compositions and conditions
in the mantle required to generate these lavas. Although quantitative models exist for the
type of anhydrous mantle melting that occurs at mid-ocean ridges, they do not accurately
predict the melting behavior of mantle previously modified in subduction zones. In
Chapter 3, we develop a new petrologic model to simulate anhydrous melting of
variably metasomatized upper mantle and explore the effects of mantle pressure,
temperature, and composition on melt composition. The model is calibrated with a
compilation of anhydrous partial melting experiments on a variety of compositions
relevant to the mantle. This model can predict the temperatures and compositions of
melts formed from variably metasomatized mantle, or calculate the temperature and
pressure at which a primary basaltic magma of interest was produced. For example,
Chapter 4 applies our model to anhydrous melt generation in southern Oregon and
northern California, where the mantle has been metasomatized by subduction of the
Farallon plate for approximately 50 Ma. A notable number of volcanic centers have
formed east of the Cascades volcanic arc within the last 25 Ma and scientists have
ascribed their formation to a variety of tectono-magmatic processes, such as crustal
extension, lithospheric delamination, or mantle flow related to the ongoing subduction to
the west. We calculate the temperatures and pressures at which primitive basalts from
these volcanic centers originated in the mantle in order to draw conclusions about the
tectonic and magmatic driving forces responsible for their formation. The calculated
depths of melt extraction are also compared to the geophysical estimates of the location
of the crust-mantle boundary in this region to place new constraints on the thickness of
the continental lithosphere.
In the broadest sense, the lithosphere is the rigid material that makes up the
Earth's tectonic plates and it is underlain by the asthenosphere, which is the part of the
Earth's mantle that undergoes solid-state creep such that it exhibits fluid-like behavior on
geologic time scales. Yet the thermal, mechanical and chemical variations of the mantle
that create the lithosphere-asthenosphere boundary (LAB) are poorly understood.
Scientists have previously attributed the negative gradient in seismic velocity observed at
the LAB to variations in chemical composition, grain size, volatile content and/or
temperature between the lithosphere and the asthenosphere. However, none of these
mechanisms are capable of producing the large negative seismic velocity gradient
observed at the LAB below eastern North America. Thick continental lithosphere, such
as found below the North American continent, is constructed through many generations
of plate subduction. In Chapter 5, we conduct finite element models of asthenospheric
flow at an abrupt lateral decrease in lithospheric thickness, similar to the geometry of the
eastern edge of the North American continent. These models suggest this geometry,
together with lateral plate motions, produces edge-driven convection and asthenospheric
upwelling capable of triggering adiabatic decompression melting. Whether the upwelling
asthenospheric mantle will melt depends on its chemical composition and volatile
content, as demonstrated in Chapters 1 and 3. Using observations from these studies, we
develop a parameterization of the H20-undersaturated peridotite solidus that can be
applied to our finite element models. This study illustrates that mantle with average
temperatures and average to slightly elevated volatile contents will produce small extent
melts below eastern North America, suitable for creating the seismic wave speeds
observed at the LAB.
In summary, the recycling of tectonic plates at subduction zones has a profound
effect on the composition and melting behavior of the Earth's mantle. Volatiles fluxed
from a subducting plate trigger melting in the mantle wedge at temperatures ~600'C
lower than would otherwise occur, as demonstrated in Chapter 1. This melting process
produces H20-rich basaltic magmas that evolve to form the highly differentiated
continental crust unique to Earth, as examined in Chapter 2. The chemical alteration of
the overlying mantle in subduction zones has lasting effects on its melting behavior, as
illustrated in Chapter 3. This melting behavior is evident in the basalts from southern
Oregon and northern California studied in Chapter 4. Finally, Chapter 5 investigates how
an abrupt change in lithospheric thickness combined with the chemical and volatile
budget of the mantle affect the rheologic structure of the Earth. Together, the work
presented in thesis advances our knowledge of how plate tectonics drives magma genesis
and consequently, the chemical differentiation of our planet.
The chapters in this dissertation are intended to be five free-standing manuscripts,
which are published (Chapter 5), accepted for publication pending revision (Chapter 1) or
in preparation for submission (Chapters 2, 3, 4). The corresponding publication reference
information is listed below in the order of the chapters presented here.
Till, C.B., Grove, T.L., Withers, T., The Beginnings of Hydrous Mantle Wedge Melting,
accepted pending revision at Contributions to Mineralogy and Petrology.
Till, C.B., and Grove, T.L., Compositions of Low Extent Melts of H20-Saturated
Peridotite, in preparation.
Till, C.B., Grove, T.L., Krawczynski, M.J., Primary basaltic magmas from variably
metasomatized mantle: The effects of pressure, temperature, and composition on
melting plagioclase and spinel lherzolite, in preparation for Journal of Geophysical
Research.
Till, C.B., Grove, T.L., Carlson, R.W., Donnelly-Nolan, J.M., Fouch, M., Hart, W.K.,
Wagner, L., in preparation for Geochemistry, Geophysics, Geosystems, Special
Theme on Volcanism in the Pacific Northwest.
Till, C.B., Elkins-Tanton, L.T. and Fisher, K.M., 2010, Low Extent Melts at the
Lithosphere-Asthenosphere Boundary, G-cubed, Vol. 11, No. 10,
doi: 10. 1029/2010GC003234.
Chapter 1:
The Beginnings of Hydrous Mantle Wedge Melting
Abstract
This study presents new phase equilibrium data on primitive mantle peridotite
(0.33 wt.% Na 2O, 0.03 wt.% K20) in the presence of excess H20 (14.5 wt.% H2 0) from
760-1200'C at 3.2-6 GPa. Based on textural and chemical evidence, we find that the
H20-saturated peridotite solidus remains isothermal between 800 and 820'C at 3 - 6
GPa. We identify both quenched solute from the vapor phase and quenched silicate melt
in supersolidus experiments. Chlorite is stable on and above the H 20-saturated solidus
from 2 to 3.6 GPa and chlorite peridotite melting experiments (containing ~6 wt.%
chlorite) melt at the chlorite-out boundary over this pressure range, which is within 200 C
of the H20-saturated melting curve. Chlorite can therefore provide sufficient H20 upon
breakdown to trigger melting in the mantle wedge or perpetuate ongoing H 20-saturated
melting. Constraints from recent geodynamic models suggest in a hot subduction zone
like Cascadia, there is significantly more H20 fluxed from the subducting slab near 100
km depth than can be bound in a layer of chloritized peridotite - km thick at the base of
the mantle wedge. Consequently in hot subduction zones, H20 is available to trigger
melting at the H20-saturated solidus in the lowermost mantle wedge. Alternately in cool
subduction zones like the Northern Marianas, a layer of chloritized peridotite up to 1.5
km thick could contain all the H20 fluxed from the slab every million years near 100 km
depth, which suggests the dominant form of melting below arcs in cool subduction zones
is dehydration melting. Slab P-Tpaths from recent geodynamic models also allow for
melts of subducted sediment, oceanic crust and/or sediment diapirs to interact with
hydrous mantle melts within the mantle wedge at intermediate to hot subduction zones.
1. Introduction
Central in our pursuit to understand the processes that give rise to arc magmas,
are the observations of the petrologic, chemical and volatile content of the erupted rocks.
Work over the past 30 years has resulted in a preponderance of evidence to suggest arc
parental magmas commonly contain up to 4 to 6 wt.% H20 (e.g., Sisson and Grove,
1993a; Carmichael, 2002; Grove et al., 2003; Pichavant and MacDonald, 2007) and some
arc andesites contain up to 8 to 10 wt. % H20 (e.g., Sisson and Grove, 1993b;
Carmichael, 2002; Grove et al., 2005; Wallace, 2005). However, considerable
uncertainty remains about the physically and compositionally complex processes that
lead to the generation of hydrous arc magmas with these observed water contents.
Furthermore, a comprehensive understanding of the temperatures of hydrous melting in
the mantle and the properties of these melts is a fundamental constraint for geodynamic
models as they develop an integrated, three-dimensional picture of the flow of fluid and
melt through the mantle wedge at subduction zones (e.g., Cagnioncle et al., 2007; Billen,
2008).
Previous work on garnet peridotite melting in the presence of H2 0 has established
that H2 0 dramatically lowers the peridotite solidus (Kushiro et al., 1968b; Green, 1973;
Millhollen et al., 1974; Mysen and Boettcher, 1975; Kawamoto and Holloway, 1997;
Grove et al., 2006; Green et al., 2010). One challenge facing experimental studies of
hydrous peridotite melting is the exsolution of H20 from hydrous melts during
depressurization, which transforms the melt and/or vapor solute into a froth of micronthick bubbles and/or quench phases. Other challenges include, attaining chemical
equilibrium, maintaining bulk composition, and distinguishing between quench of a
silicate-rich vapor phase and a melt phase to determine if a near solidus experiment is
melted. Here we address these challenges and document the systematics of melting of
H20-oversaturated and chlorite-bearing undepleted peridotite from 3 to 6 GPa. These
experiments are used to understand the temperatures and chemical reactions of mantle
wedge melting that constitute the primary controls on 1) the location of arc volcanoes, 2)
the width of the volcanic arc, and 3) the abundance of wet vs. anhydrous primary
magmas at arcs.
2. Experimental and Analytic Methods
2.1 StartingMaterials
The primitive upper mantle composition of Hart and Zindler (1986) was
synthesized by two different methods as experimental starting materials for this study
(Table 1). The starting material for the H20-saturated experiments ("H & Z + H20" in
Table 1) was created from a mixture of high-purity synthetic oxides or metasilicates.
Brucite, Mg(OH) 2, was added to the synthetic oxide mixture instead of MgO, which
provided the starting composition with ~14.5 wt% H20. The chlorite peridotite
experimental starting material ("H & Z natural" in Table 1) is a mixture of natural
olivine, orthopyroxene, clinopyroxene, and garnet grains separated from a garnet
lherzolite xenolith found in Pali-Aike, Patagonia (Stern et al., 1989; Koga et al., 1999)
and chlorite from New Idria, California (Van Baalen, 2004). The chlorite (Mg#=0.94) is
near the Mg-end member clinochlore (MgsA1)((AlSi 3)O0o(OH) 8 ) which contains ~12 wt.
% H20 (Deer et al., 1962). Thus the chlorite peridotite starting material for these
experiments contains ~0.7 wt.% H2 0. All starting materials were weighed out and
ground for 3 hours under ethanol, using an agate mortar and pestle.
2.2 Piston Cylinder Experimental Methods
The H20-oversaturated experiments from 3.2 to 4 GPa and the chlorite peridotite
experiments were performed in a 0.5" end-loaded solid-medium piston cylinder apparatus
(Boyd and England, 1960) in the MIT Experimental Petrology Laboratory (Table 2). All
piston cylinder experiments were conducted in Au capsules to minimize the loss of Fe
from the starting material to the capsule and because Au is impermeable to hydrogen at
these experimental conditions. The experimental samples were prepared by packing 30
mg of powdered starting material into an Au capsule fabricated from 0.375" long Au
tubes with a 0.01" wall thickness. The Au tube was triple-crimped and welded at both
ends and flattened to make a uniformly thick capsule top and bottom. The welded
capsule is ~0. 15" in length and able to retain the H2 0 added to the starting material as
brucite or chlorite. The sealed capsule was then placed in a nonporous A12 0 3 sleeve and
positioned in the center of a graphite furnace using MgO spacers. The pressure medium
consisted of sintered BaCO 3 and the pressure was calibrated using the breakdown
reaction of Ca-tschermak pyroxene to anorthite + gehlenite + corundum and the spinel to
garnet transition in CMAS (Hays, 1966; Longhi, 2005). The run pressures for these
experiments are thought to be accurate to ± 0.05 GPa. The temperature was monitored
and controlled using W97Re3/W75Re 25 thermocouples with no correction for the effect of
pressure on thermocouple EMF. The thermal gradient in our piston cylinder assembly
has been determined by two independent methods: direct measurement with offset
thermocouples and temperature mapping using the kinetics of the MgO + A12 0 3 =
MgAl 20 4 reaction (Watson et al., 2002; Medard et al., 2008). Both methods indicate that
the hot spot in the graphite furnace containing the experimental capsule is -20'C (18 ± 6
'C: Medard et al., 2008) hotter than the thennocouple location, 1.5 mm above the
capsule. The temperature difference across the capsule is ~10 0 C (9 ± 5 'C: Medard et al.,
2008), with the hottest temperatures at the top of the capsule. The run temperatures for
our experiments are thus accurate to ± 100 C. To raise the piston cylinder to pressure,
experiments were pressurized to 1 GPa at room temperature, then the temperature was
raised at 100 0 C/min to 865'C or the final run temperature if below 865'C. The
experiment was held at these conditions for 6 minutes before raising the pressure to the
final experimental pressure (reported in Table 2) and the temperature at 50'C/min to the
final temperature if >865'C. The sample was held at isothermal and isobaric conditions
for the duration of the experiment (24-338 hours; Table 2) and then quenched by turning
off the power. Following each experiment, a hole was drilled in the capsule with a small
diameter hand drill and the presence or absence of liquid water was noted (Table 2). The
capsule was then sliced in half, dried, vacuum-impregnated with epoxy and polished for
electron microprobe analysis.
2.3 Multi-anvil ExperimentalMethods
Six H20-saturated experiments at 5 and 6 GPa (Table 2) were conducted in the
1000-ton Walker-type multi-anvil apparatus at the University of Minnesota Experimental
Petrology Lab. Au tubing 2 mm in diameter and 2.5 mm in length was quadruplecrimped, welded and flattened on one end to form the experimental capsule. The capsule
was then packed with 10 mg of powdered starting material and the open end was singlecrimped and welded in a water cooled fixture to form a closed capsule -2 mm in length.
For these experiments, WC anvils with 12 mm truncations, cast MgO-A 2 0 3 -SiO 2 -Cr 2 O3
octahedra with integrated gasket fins and straight-walled graphite heaters (comprising the
"12-TEL assembly") were used. Further details of the experimental design are given by
Withers et al. (in press). The pressure calibration for the 12-TEL assembly was
established by bracketing multiple high temperature fixed points (Dasgupta et al., 2004).
On the basis of 6 brackets and 2 half-brackets at temperatures that range from 700 to
1350'C, the pressure uncertainty is estimated to be ± 0.3 GPa. The force-pressure
relationship shows no observable dependence on experimental temperature within the
700-1350'C temperature range. The temperature uncertainty, based on orthopyroxeneclinopyroxene thermometric measurements and brackets of the congruent sphene melting
reaction, is estimated to be ± 15'C. Experiments were quenched by switching off the
heating circuit and subsequently were depressurized over a period of -10 hours.
2.4 Analytical Methods
The major element concentrations of all experimental products were analyzed by
wavelength dispersive spectrometry on the 4- or 5-spectrometer JEOL 733 and 8200
JEOL electron microprobes (EMP) at the Massachusetts Institute of Technology. All
analyses of minerals presented in Table 3 were conducted with a 15kV accelerating
voltage, a 10 nA beam current and a 1 pm spot size. Online data reduction utilized the
CITZAF correction package (Armstrong, 1995) and the atomic number correction, the
absorption coefficients, and the fluorescence correction available in CITZAF.
3. Experimental Results
3.1 Mineral PhaseRelations & Textures
Extensive EMP analyses of all minerals present in our experiments were
conducted to assess the homogeneity of each phase, and the extent of any systematic
changes in composition with temperature or pressure (Table 3). All experiments in this
study run below 1 100'C contain the minerals olivine, orthopyroxene, clinopyroxene and
the anhydrous aluminous phase, garnet (Figure 1; Table 2). At temperatures greater than
1 100 C at 3.2 GPa olivine, orthopyroxene and spinel are the dominant solid phases.
Textures in our experiments are identical to those described in Grove et al.
(2006). Olivines tend to be euhedral to subhedral (10-150 [m) and orthopyroxenes are
small (<20 [tm long) needle-like laths in the lower temperature experiments and more
prismatic at higher temperatures (<40 [tm). Clinopyroxenes also tend to be elongate and
prismatic and increase in size with temperature (20-100
stm).
Garnets tend to be larger
(100-200 [tm) and poikilitic in the experiments where chlorite is absent (Figure 2).
Where garnet and chlorite coexist, garnet tends to be smaller (<40 [tm), likely limited in
size by Al-partitioning between these two Al-bearing phases.
We observe a marked change in texture in our experiments at ~8 10 'C from 3 - 6
GPa. Experiments conducted at or below 800'C tend to be fine grained and
homogeneous in texture with no segregation of the phases over the height of the capsule,
whereas those at temperatures above 800'C are coarse grained, exhibit significantly more
void space and are mineralogically-zoned, with orthopyroxene, clinopyroxene and garnet
segregated toward the cool region (bottom) of the capsule and large olivine grains and
melt segregated toward the hot end (top) (Figure 2). These textures are similar to some
textures documented in other fluid-present melting experiments (e.g., Stalder and Ulmer,
2001; Schmidt and Ulmer, 2004; Kessel et al., 2005; Grove et al., 2006; Luth, 2006;
Medard and Grove, 2006). Our experiments also exhibit changes in mineral composition
that occur with increasing temperature above 800'C at 3.2 and 4 GPa. Figure 3 illustrates
a persistent increase in the Mg# of the solid phases with increasing temperature at
constant pressure. These experiments at constant pressure also reveal trends in the minor
element concentrations of the clinopyroxene, orthopyroxene and garnet with increasing
temperature as shown in Figure 4. Additional abrupt changes in texture or composition
of the experimental phases are not observed with increasing temperature, including the
temperature range from 1000 to 1300'C at 3.2 and 4 GPa. In experiments at or above
880 0 C, small wisps of inter-granular glassy quenched material are visible (Figure 2).
These observations lead us to interpret the abrupt change in texture between 800
and 820'C and the progressive change in mineral composition above this temperature, as
evidence of the presence of the H20-saturated solidus at ~8 100 C from 3 to 6 GPa.
Experiments bracket this critical behavior within 20'C at 3.2, 3.6, 4, and 5 GPa and 30'C
at 6 GPa and reveal the steep slope of the solidus in P-T space (Figure 1). We attribute
the modal zoning in experiments at or above 820'C as the result of thermal compaction
facilitated by the presence of melt, a small thermal gradient and the long experimental
durations. The modal zoning does not occur in experiments at temperatures less than
820 0 C, likely because only crystals and aqueous fluid are present and the solubility of the
silicates in solution is not sufficient for dissolution and recrystallization to occur.
Mineral phases in subsolidus experiments with durations longer than 168 hours
are compositionally uniform throughout the experiment and homogeneous (Figure 2;
Table 3). Supersolidus experiments, with durations of 168 hours or longer near 800'C
and 24 hours or longer for T >1050'C, also exhibit homogeneous mineral compositions
with orthopyroxene crystals displaying a more Ca-rich rim, which is likely the result of
interaction with the melt during quenching.
Chlorite is stable to temperatures just above the H20-saturated solidus between 2
and 3.6 GPa (Figure 1). Above 3.6 GPa, the H20-saturated solidus and the chlorite
stability fields diverge. The composition of the chlorite in the experimental products is
closest to the Mg-rich chlorite end-member, clinochlore (Table 3). Chlorite is distributed
throughout the experimental charges, and not restricted to the edges or cracks where
hydrous fluids segregate. A similar high-pressure chlorite phase has been observed in
experiments in the MgO-A12 0 3 -SiO 2-H 2 0 (MASH) (Pawley, 2003) and Na 20-CaO-FeOMgO-Al 2 0 3 -SiO 2 -H 20 (NCFMASH) systems (Fumagalli and Poli, 2005; Dvir et al.,
2011). This work contributes 10 new experiments to further define chlorite stability
between 3.2 and 4 GPa and better constrains this phase boundary at 3.6 and 4 GPa
relative to previously published experiments (Table 2).
A small amount of carbonate appears as a quench phase in experiments conducted
more than a year and a half after the preparation of our H & Z + H 2 0 starting material,
although no carbon was used in the synthetic oxide mixture and no melt traps were used
in the experiments. We interpret the appearance of carbonate as the result of progressive
carbonation of the synthetic oxide mixture over time from atmospheric CO 2 . This has
been observed in other experimental studies when the starting mix is not kept under
vacuum (M. Schmidt, personal communication, 2010). These experiments are still
valuable, as a small amount of CO 2 added into our starting material to create a mixed
C0 2-H 20 vapor is more realistic for simulating mantle melting processes than a pure H2 0
volatile component, as CO 2 concentrations of the mantle are estimated to be 50-2000 ppm
below mid-ocean ridges (Saal et al., 2002) and 1000-4000 ppm below oceanic islands
(Trull et al., 1993; Dixon et al., 1997; Pineau et al., 2004).
3.2 Chlorite Peridotite(H & Z Natural)Melting Behavior
Six chlorite peridotite melting experiments were conducted using the natural
mineral starting material, which contains no H2 0 except that bound in the hydrous phase
chlorite (chlorite contains 12 wt.% H20, ~6 wt.% chlorite in starting material =-0.7
wt.% H20 in starting material). In these experiments, the solidus is displaced to slightly
higher temperatures, relative to the H20-saturated case, and follows the chlorite-out
boundary between 2 and 3.6 GPa (Figure 5). This is significant because it suggests there
is sufficient H2 0 released from the breakdown of chlorite to allow H20-saturated
melting. The chlorite-out boundary, and therefore the solidus for chlorite peridotite, is 2030'C higher in temperature than the H20-saturated solidus between 2 and 3.6 GPa.
Above 3.6 GPa and below 2 GPa, chlorite breaks down at temperatures below the H 02
saturated solidus and the first melts occur at the H20-saturated solidus.
3.3 Approach to Equilibrium
The H20-saturated piston-cylinder and multi-anvil experiments reported in this
paper are all synthesis experiments (i.e., minerals grew from a synthetic oxide mixture or
are reacted natural minerals selected from a garnet peridotite and natural chlorite sample)
with durations from 24 to 338 hours (median = 168 hours) (Table 2). The combination of
the H20-saturated experiments and experiments using natural mineral powders as starting
material, are analogous to reversal-type experiments. Both types of experiments yield
similar mineral compositions and solidus temperatures, with the exceptions of
experiments where chlorite is stable above the H20-saturated solidus. In addition, three
observations indicate that in both types of experiments an approach to equilibrium was
attained: (1) the growth of compositionally homogeneous chemically unzoned minerals
with triple-junction contacts to one another, (2) maintenance of constant bulk
composition, and (3) temperatures calculated from coexisting orthopyroxene and high-Ca
clinopyroxene in 3.2 GPa H20-saturated experiments yield temperatures similar to the
experimental conditions(Grove et al., 2006). Two experiments that bracket the H 20saturated solidus at 3.2 GPa were run for 24 hours (D169 and D167) for comparison to
longer experiments at these conditions. Melting textures are detected in the higher
temperature of these short duration experiments, but mineral compositions are very
heterogeneous and individual grains are strongly zoned. An experiment run for only six
hours at 880'C and 3.2 GPa (D212; Table 2) remained synthetic powder starting material
with the exception of the growth of several small garnet porphyroblasts.
4. Discussion
4.1 Comparisonwith Previous ExperimentalStudies
Seven prior experimental studies have investigated the H20-saturated melting of
peridotite over a range of similar pressures (Kushiro et al., 1968b; Green, 1973;
Millhollen et al., 1974; Mysen and Boettcher, 1975; Kawamoto and Holloway, 1997;
Grove et al., 2006; Green et al., 2010) and are compared in Table 4. The P-T conditions
of H20-saturated melting in the 1960's and 1970's studies fall into two groups based on
the duration of the experiment. Shorter experimental durations (<24 hours) find solidus
temperatures of - 0000 C at 3 GPa and longer experimental durations (>24 hours) yield
solidus temperatures of -800'C at 3 GPa (Table 4). Modem experimental techniques
allow us to avoid problems encountered in the earlier studies, for example short
experimental durations, which lead to problems attaining equilibrium, and non-ideal
capsule material (e.g., Pt or Mo) that lead to loss of H2 and Fe. Short experimental
duration peridotite melting experiments yield quantities of solidus depression similar to
those observed for H20-saturated diopside, while longer experimental duration
experiments yield quantities of solidus depression that are -200'C greater, similar to
those observed for H20-saturated forsterite (Kushiro, 1969; Hodges, 1974; Eggler and
Burnham, 1984; Luth, 1993) and confirms the likelihood that shorter experimental
durations record the solidus depression of the phase with the fastest melting kinetics
rather than equilibrium melting. The results presented here are most similar in the
temperature of the H20-saturated solidus to those of Mysen and Boettcher (1975), whose
experiments contained ~20 wt.% H2 0, lasted up to 64 hours, and employed AgPd
capsules, which are inert to exchange of Fe with the starting material (Table 4).
Scientists have long debated whether aqueous fluids in equilibrium with silicate
minerals at high pressures (> 3 GPa) contain high concentrations of silicates in solution
(Bowen and Tuttle, 1949; Kushiro, 1972a; Nakamura and Kushiro, 1974; Tatsumi et al.,
1986; Zhang and Frantz, 2000; Newton and Manning, 2002) and some recent work
suggests fluid solute concentrations may be on the order of 15-40 wt. % when in
equilibrium with model mantle compositions for the pressure range considered here
(Stalder et al., 2001; Mibe et al., 2002; Kessel et al., 2005; Dvir et al., 2011). Also
relevant are the many studies that examine the location of the second critical end point
(SCEP) in simple compositional systems (e.g., 1.5 GPa in albite-H 20: Bureau and
Keppler, 1999; Stalder et al., 2000); ~12 GPa in MSH-H 20: Stalder et al., 2001).
However the location of the SCEP is not well understood in the peridotite-H 20 system
due to the difficulty of these experiments. Direct observations of the SCEP in the
peridotite-H 20 system were made using in situ X-ray radiography techniques in a
multianvil apparatus and locate the SCEP by the disappearance of distinctive round
shapes, thought to represent spheres of either aqueous fluid or silicate melt in the other
phase formed from interfacial tension below the SCEP (Mibe et al., 2007). Using this
technique the SCEP was found to be -3.8 GPa and 1000 0 C. Based on the presence of
both quench spheres and wisps in our supersolidus experiments with distinctive
compositions (Figure 6, Table 3) that we interpret to be quenched fluid and melt,
respectively, we conclude all of our supersolidus experiments are located below the
SCEP. Therefore the question of how to determine quenched melt from quenched vapor
phase in our experiments logically arises.
Using recent observations of element solubility in high-pressure fluids in concert
with our experimental data, we can examine the quenched vapor versus melt question for
our experiments. Na, Al, are found to be the most soluble major elements in a high
pressure fluid phase in equilibrium with peridotite over the pressure range 2-6 GPa,
followed by Ca and Si, while Mg and Fe are found to be relatively insoluble (Brenan et
al., 1995; Adam et al., 1997; Ayers et al., 1997; Stalder et al., 1998; Dvir et al., 2011).
These fluids are also rich in large ion lithophile elements relative to the high field
strength elements. Figure 3 reveals a steady increase in the Mg# of the solid phases in our
experiments and Figure 4 illustrates the change in minor element concentrations in the
two pyroxene phases and garnet, both with increasing temperature at 3.2 GPa. As
demonstrated by the predicted batch melting curves for these elements, the changes in the
composition of the solid phases are not consistent with simple partitioning of elements
into a vapor phase. Similar changes in the clinopyroxene composition with increasing
degree of melting are observed in Tenner et al. (2009) and Kawamoto and Holloway
(1997) used an increase in the Mg# of olivine crystals with increasing temperature as a
criterion for determining when experiments were above the solidus (Table 4). Therefore
we conclude the experiments above 81 00 C exhibit evidence for quenched melt, in
addition to a quenched vapor phase.
In addition to the persistent increase in Mg# of the solid phases, two abrupt
changes in the Mg# of several of the solid phases can also be identified: a jump in the
Mg# of all solid phases between 840 and 880'C corresponds to the breakdown of chlorite
and a drop in the Mg# of olivine and clinopyroxene between 940 and 980'C corresponds
to the breakdown of clinohumite (Figure 3). The breakdown of chlorite can be easily
observed in our experiments as discussed in section 4.4. The abrupt change in Mg#
between 940 and 980'C is consistent with the location of the clinohumite breakdown
reaction in the MASH system (Pawley, 2000; Stalder and Ulmer, 2001). Small amounts
of titanoclinohumite are observed in our experiments at temperatures below 900'C at
both 3.2 and 3.6 GPa (Table 2 & 3). Breakdown of clinohumite may occur in a reaction
proposed by Stalder and Ulmer (2001) where,
0.5 enstatite + clinohumite = 5forsterite + H2O.
(1)
Thus the decrease in the modal abundance of orthopyroxene and the simultaneous
increase in the abundance of olivine observed in our experiments between 880 and 980'C
also suggests clinohumite breakdown is occurring in our experiments between these two
temperatures (Table 2). Because our experiments are not conducted in the simplified
MASH system, the iron in the orthopyroxene may then be partitioned into olivine and
clinopyroxene and drive their Mg# down at clinohumite breakdown. A similar "leveling"
of the Mg# of the solid phases between phase boundaries is observed in hydrous
peridotite melting experiments by Mysen and Kushiro (1977) at constant pressure and
increasing temperature.
An alternative interpretation of the change in Mg# of two solid phases at either
temperature is that they represent the solidus boundary. For example, if there is an
increase in the Mg/Si ratio of the solute, either at the solidus or with increasing pressure
(Stalder et al., 2001; Dvir et al., 2011), this could produce a corresponding drop in the
Mg-content of solids. However, these compositional shifts do not correspond to any
determinations of the temperature H20-saturated peridotite solidus, at either -800'C
(Mysen and Boettcher, 1975; Grove et al., 2006) or >1000'C at 3 to 6 GPa (Kushiro et
al., 1968b; Green, 1973; Millhollen et al., 1974; Kawamoto and Holloway, 1997; Green
et al., 2010). Furthermore, no abrupt change in Mg# is observed in garnet or
orthopyroxene between 940 and 980'C. Our starting material has sufficient H20 to have
an H20 vapor phase present in addition to the H20 partitioned into a melt to more than 55
wt.% melting. Therefore changes in the Mg# of the solid phases above 840'C are either
the result of phase transitions, such as chlorite and clinohumite breakdown, or require a
better understanding of element partitioning between solid, melt, and fluid to resolve.
In the low-K 20, -Na 20, -TiO 2 and high H 20 (14.5 wt.%) bulk mantle composition
used in our experiments, amphibole is stable to -2 GPa and 10000 C as presented by
Grove et al. (2006). Other experimental studies find that alkali and TiO 2 contents, as well
as oxygen fugacity, have a significant effect on amphibole stability and can stabilize
amphibole to pressures up to 3 GPa in lherzolite compositions (Wyllie, 1978; Mengel and
Green, 1989; Niida and Green, 1999; Fumagalli and Poli, 2005). Green et al. (2010)
suggest pargasite is stable to 3 GPa for a bulk composition similar to ours with
significantly lower H20 concentrations (< 1.45 wt.%), which they argue prevents
leaching of highly soluble elements such as Na 20. They present P-T conditions and the
phase assemblage for two experiments with pargasite stable at 2.5 GPa and 0.3 wt.% H2 0
and for an experiment at 3.0 GPa and 0.5 wt.% H20. Pargasite is also listed as a stable
phase in a 2.5 GPa experiment at anhydrous conditions. The high H20 content of our
experiments may affect amphibole stability, as the mole fraction of H20 in a mixed CO 2
+ H20 fluid phase has been observed to affect the thermal stability of pargasite at
constant pressure in the system pargasite-H20-CO2 (Holloway, 1973) and phlogopite in a
spinel lherzolite composition (Yoder and Kushiro, 1969; Wendlandt and Eggler, 1980).
However, changes in hydrous phase stability with bulk H20 content did not affect the
solidus temperature in the above cases and likewise would not affect the determination of
the H20-saturated peridotite solidus presented here.
Instead, the method of determining whether an experiment is sub- vs. supersolidus
may play a large role in the determination of solidus temperature. Green et al. (2010)
uses the presence of interstitial patches of amphibole and/or quench clinopyroxene, mica
and glass, all with Mg#'s of 70-85 within the lherzolite and melt trap layers as the
diagnostic criterion for the presence of quenched hydrous silicate melt in an experiment.
Conversely, the lack of the quench mafic silicates, as well as the presence of hydrous
phlogopite or pargasite with Mg#'s greater than or equal to olivine is used to diagnose
their subsolidus experiments. This choice of criterion for sub- and supersolidus
experiments may bias their identification of the solidus to high temperatures where melts
are mafic enough to grow amphibole ± clinopyroxene upon quench. Experimental work
thus far on the composition of hydrous spinel and garnet lherzolite melts suggests the
melt is andesitic at lower pressures and temperatures and becomes more mafic as melting
continues (Kushiro et al., 1968b; Kushiro, 1972b; Kawamoto and Holloway, 1997;
Gaetani and Grove, 1998; 2003; Grove et al., 2006). The coexistence of melt with
olivine in the hotter portion of hydrous melting capsules indicates peritectic melting is
occurring to produce a SiO 2-rich melt composition (Kushiro et al., 1968b). Furthermore,
Gaetani and Grove (1998; 2003) found that the SiO 2 increases by -1 wt.% and FeO and
MgO decrease by -2 wt.% in hydrous experimental melts with the addition of 3-6 wt.%
dissolved H20. Preliminary analyses of the melt in our experiments support these
findings (Table 3). Our textural and chemical diagnostic criterion as discussed above, do
not bias our detection of melt to any melt composition. Green et al. (2010) also observes
euhedral olivine or pyroxene crystals "coated with fragmented glass 'froth' or with films
of silicate glass" in high water content (> 5 wt.%) experiments, which they interpret to be
subsolidus. We observe similar features in experiments we interpret to be supersolidus
based on the Mg# and minor element composition of the solid phases and textures. The
lack of compositional data for the minerals, melts and fluids in the experiments presented
in the Green et al. (2010) study make it difficult to further resolve our conflicting
interpretations of the solidus temperature.
Other experimentally-determined hydrous solidi have similarly steep slopes to
that presented here for the H20-peridotite system. The H20-saturated solidus for a KLB1 model mantle of Kawamoto and Holloway (1997) as well as that for the system
forsterite+ H20 (Luth, 1993) both remain isothermal between 3 and 12 GPa. The critical
curves for both hydrous sediment and basalt also exhibit steep slopes over several GPa
near their inflection point (e.g., Hermann and Spandler, 2008; Hill and Boettcher, 1970b),
as does the MgSiO 3 + H20 system (Kushiro et al., 1968a; Hodges, 1974). Theoretical
analysis of wet melting relations in a MSH system suggests that such steep slopes are
common as P-T minima are reached in these types of critical curves and represent an
inflection in density where A V=0 for the phases involved (Boettcher and Wyllie, 1969;
Hack et al., 2007). Eventually at a pressure to be determined, the H20-saturated
peridotite solidus will turn over to have a positive slope as the change in volume of the
vapor phase and solids outpaces the change in volume of the melt phase for a given AP.
4.4 Chlorite Stability in the Mantle Wedge
Overall the chlorite-out boundary from our experiments agrees with that from the
MASH system (Pawley, 2003) and is -50'C higher at 3.2 to 4 GPa than experiments in
the NCFMASH system (Fumagalli and Poli, 2005; Dvir et al., 2011), which exhibit less
magnesian chlorite than our experiments. Figure 3 illustrates iron's affinity for garnet
relative to the other solid phases including chlorite in our experiments. The chlorite
breakdown reaction calculated at 4 GPa between 760 and 800'C is,
0.06 enstatite + 0.05 clinopyroxene + 0.19 chlorite = 0.19 olivine + 0.11 garnet
(2)
Thus in more iron-rich peridotite compositions chlorite stability is shifted to lower
temperatures in favor of olivine and garnet. This thermal depression in the stability of
chlorite with increasing iron content has been observed in other hydrous melting
experiments on Martian mantle compositions, which contain two to three times as much
wt.% iron as the Earth's mantle (Medard and Grove, 2006).
A variety of rocks exhumed from high pressures in the Earth support our
experimental observations of chlorite stability in peridotite at upper mantle pressures.
Chlorite has been found in relatively high-grade metamorphosed ultramafic rocks from
the lower crust, such as in the Western Gneiss Region, Norway (Medaris, 1984; Hacker
et al., 2010) and in peridotite inclusions in kimberlites from the Colorado Plateau, USA
(Smith, 1979). New evidence from exhumed exposures of the sub-arc mantle, like the
Higashi-akaishi peridotite of the Sanbagawa belt in SW Japan (Enami et al., 2004;
Hattori et al., 2009; Till et al., 2009), also support the conclusion that chlorite is stable in
the mantle wedge. High-Mg chlorite is found as a stable prograde mineral and occurs as
euhedral crystals intergrown with and as inclusions in garnet or pyroxene crystals in the
garnet peridotites from Higashi-akaishi, where the garnet rims represent the peak
metamorphic conditions of 2.9-3.8 GPa and 700-810 C ca. 110 Ma, similar to the P-T
conditions of experiments in this study containing chlorite.
The experiments in this study support the hypothesis of Schmidt and Poli (1998)
and Grove et al. (2006) that chlorite is a stable mineral in the P-T-X range appropriate for
the base of the upper mantle wedge. Chlorite, as well as serpentine, likely form with
fluids released during the dehydration of serpentine and amphibole in the subducting
oceanic crust and lithosphere (Schmidt and Poli, 1998; Iwamori, 1998; Pawley, 2003;
Grove et al., 2006; Iwamori, 2007) and/or from the dehydration of subducted sediments
in the mantle wedge nose at <80 km depth. Seismic images of the Honshu subduction
zone reveal a low velocity layer that is interpreted to result from the presence of hydrous
phases in the shallow wedge (Kawakatsu and Watada, 2007). The chlorite and serpentine
can then be advected downward in the mantle wedge by corner flow. Provided all the
A12 0 3 present in the fertile mantle composition used here is available to form chlorite, the
mantle above the slab can contain up to 6.5 wt.% chlorite, which equates to 2 wt.% H2 0
in the mantle. The degree of serpentinization in an upper mantle with forsteritic olivine
depends on the extent to which the system is open and fluids can transport elements such
as magnesium away from the reacting volume. In an open system where complete
serpentization can occur, the maximum H20 content of the upper mantle will be that of
serpentine; 13 wt.% bulk H20. Therefore, 1 km3 of serpentinized mantle (up to 13 wt.%
H2 0) with a density of 3.3 g/cm 3 can hold up to can hold up to 390 Tg of H2 0 and 1 km 3
of chloritized mantle (2 wt.% H20) can hold up to ~66 Tg of H20. In a more depleted or
Fe-rich bulk composition, the addition of water can completely metamorphose olivine to
serpentine and magnetite in a closed system and chlorite will breaks down at slightly
lower temperatures as discussed above. Our experiments illustrate chlorite is stable at the
undepleted peridotite solidus between 2 and 3.6 GPa and thus has the ability to supply
water during its breakdown to trigger H 20-saturated melting of the mantle, like in the
experiments presented here with natural chlorite peridotite starting material, or to
perpetuate H20-saturated melting previously triggered by H20 fluxed from the hydrous
phases in the subducting mantle lithosphere.
4.5 Implicationsfor the Generation of Hydrous Melts in the Mantle Wedge
When combined with numerical models, our observations on chlorite stability and
the H20-saturated peridotite solidus can provide constraints on melt generation in the
mantle below the more than 40,000 km of arcs on Earth today. Forward numerical
models of subduction zones have provided insights into how plate tectonic variables,
such as convergence rate, slab age, dip, and structural setting (i.e., trench advance or
rollback), affect the temperature distribution, dynamics and structure of the mantle wedge
at subduction zones (e.g., Davies and Stevenson, 1992; Peacock et al., 2005; Kincaid and
Griffiths, 2003; van Keken et al., 2002; Cagnioncle et al., 2007; van Keken et al., 2008).
Here we focus on recent geodynamic models that explore slab surface temperature
distributions calculated for a global suite of arc systems and therefore cover the global
variability in plate tectonic variables. These models pay particular attention to the
location and nature of the transition from decoupling to coupling between the slab and the
overlying mantle wedge near the cold corner (e.g., Wada and Wang, 2009; Syracuse et
al., 2010). For all of the different sets of assumptions regarding the location of the slabmantle coupling, global slab surface P-T paths cross the serpentine-out curve (Ulmer and
Trommsdorff, 1995), the chlorite-out curve (this study), and the H20-saturated peridotite
solidus (this study) in the range of pressures and temperatures that are present at the slab
surface directly below the positions of volcanoes in the global arc data set (2.5-6 GPa,
650-900'C: the region of overlap of the gray rectangles in Fig. 7). Therefore according
to the numerical models, the thermal conditions just above the slab and directly below all
global arc volcanoes are ideal for triggering the breakdown of hydrous minerals in the
mantle wedge and subsequently H20-saturated mantle wedge melting.
Using these thermal models and recent estimates of H20 flux from subducting
slabs at depth (van Keken et al., 2011), we can estimate the percentage of H20 fluxed
from the slab that can be bound in chlorite in the base of the mantle wedge at subarc
depths. Our constraints on the P-T conditions of chlorite breakdown and the slab surface
temperatures from Syracuse et al. (2010) where slab-mantle coupling occurs at 80 km
depth predict chlorite would be stable in a region -1 km thick above a slab surface at 100
km depth in the hottest subduction zones and a region up to 8 km thick in the coldest
subduction zones, if the thermal gradient at the slab surface varies between 250 C/km 50'C/km (Cagnioncle et al., 2007). The P-T paths for the crust-mantle boundary at 7 km
beneath the slab surface in the Syracuse et al. (2010) models predict that the subducting
lithospheric mantle in all except the coolest subduction zones cross the serpentine
dehydration curve but remain below the chlorite dehydration curve at subarc depths.
Therefore only H20 from serpentine is released from the subducting mantle lithosphere at
subarc depths. The calculations of van Keken et al. (2011) indicate that for the warm
Cascadian subduction zone, 11 Tg H20/Myr/m (along strike) is fluxed from the slab
between 55 and 102 km depth if the slab has a serpentinized upper mantle that is 2 wt.%
H20 to 2 km depth. At the other end of the spectrum, they determine 3 Tg H2 0/Myr/m is
fluxed from the cool North Marianas slab between 79 and 114 km depth given the same
set of assumptions regarding slab hydration. Therefore near 100 km depth, - 1000 Tg
H 2 0 are released over an area of 47 km 2 on the Cascadian slab surface due to serpentine
dehydration, or -3000 Tg H20 over 35 km 2 on the Northern Marianas slab, every million
years. Recall 1 km 3 of chloritized mantle peridotite can hold up to 66 Tg of H2 0, so a
layer of chlorite peridotite ~1 km thick over 47 km 2 could hold -3102 Tg H2 0.
Therefore in a hot subduction zone like Cascadia, there is significantly more H20 fluxed
from the subducting slab near 100 km depth every 1 million years than can be bound in a
layer of chloritized mantle wedge peridotite -1 km thick. This suggests, free H2 0 is
available to trigger melting at the H20-saturated solidus in the lowermost mantle wedge
(Fig 8a & b) and the subsequent dehydration of chlorite formed at the base of the mantle
wedge can perpetuate ongoing H20-saturated melting or ensure that H20 is present above
the H20-saturated solidus at subarc depths. In the Northern Marianas, a layer of
chloritized peridotite up to 1.5 km thick could contain all the H20 fluxed from the slab
every million years near 100 km depth, which is well within the 8 km region above the
slab where we predict chlorite to be stable in cool subduction zones. Therefore in a cool
subduction zone like the Northern Marianas, no free H20 is available to trigger melting
until chlorite breakdown in the lowermost mantle wedge and the dominant form of
melting at subarc depths is dehydration melting (Fig. 8c & d). This highlights the
significance of chlorite as a stable phase in the mantle wedge and as a provider of the
H2 0 necessary to trigger melting below arcs in cool subduction zones. Only the coldest
outlier P-T path (i.e., Tonga) in Figure 7, which is slightly too cold at subarc depths to
cross the H 20-saturated peridotite solidus, does not fit a model where the inception of
hydrous melting below arcs is caused by some combination of the low temperature of the
H 20-saturated solidus and chlorite dehydration. We thus envisage a continuum in the
way in which melting is initiated in the base of the mantle wedge at global arcs between a
Cascadia-type melting scenario and a Marianas-type melting scenario.
Many arc lava trace element and isotope compositions suggest a subducted
sediment and/or altered oceanic crust component is also added to hydrous melts of the
mantle wedge at subduction zones. However, the trace elements such as Ba, Th, Be and
Pb that are indicative of a subducted sediment component are relatively insoluble in the
aqueous fluids and thus cannot be carried by the fluids fluxed from the subducting slab
alone. A number of the P-T paths for the intermediate (e.g, Sunda and Solomon) to hot
(e.g., Chile and Cascadia) global arcs cross both the water-saturated sediment (Hermann
and Spandler, 2008; Poli and Schmidt, 2002; Nichols et al., 1994; Vielzeuf and Schmidt,
2001) and oceanic basalt solidus (Lambert and Wyllie, 1970; Hill and Boettcher, 1970a;
Liu et al., 1996) between 1.5 and 4 GPa, followed by the H20-saturated peridotite solidus
(Figure 7). Thus above 1.5 GPa at these arcs, melts of both sediment and basalt may
form at mantle wedge conditions and then ascend into the mantle wedge and interact with
early melts of the H20-saturated mantle, just above the slab surface to form a composite
melt with the trace element characteristics of many arc lavas. An alternate observation is
that global slab P-T paths when combined with estimates of the maximum water contents
of hydrated MORB at subarc P-T conditions (Hacker, 2008) suggest most slab crust has
fully dehydrated when it reaches subarc depths (Syracuse et al., 2010). Recent
numerical models find that at temperatures less than -850'C, subducted sediment layers
form buoyant instabilities (i.e., diapirs) that rise into the overlying mantle wedge within
40 km of the slab depth below the arc at 13 or 17 global subduction zones (Behn et al., in
press). Therefore alternatively, sediment diapirs may advect relatively fluid-poor
sedimentary material into the mantle wedge where they contribute their chemical
signature to hydrous mantle wedge melts in the hotter core of the mantle wedge, which
reaches temperatures in excess of -1050'C, the temperature required to deplete
metasediments of Th, Sr, Pb and Nd (Bebout et al., 1999; Becker et al., 2000; Busigny et
al., 2003; Behn et al., in press). In either case, the H20-saturated peridotite solidus
presented here provides a mechanism for hydrous mantle melts to be generated at
relatively low temperatures in the base of the mantle wedge and then interact with either
sediment melts or diapirs and perhaps in some cases melts of the oceanic crust, to form
the chemical composition of arc volcanics worldwide.
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Wallace, P.J., 2005, Volatiles in subduction zone magmas: concentrations and fluxes
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Watson, E.B., Wark, D.A., Price, J.D., and Van Orman, J.A., 2002, Mapping the thermal
structure of solid-media pressure assemblies: Contributions to Mineralogy and
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Wendlandt, R.F., and Eggler, David H, 1980, The Origins of Potassic Magmas: 2.
Stability of Phlogopite in Natural Spinel Lherzolite and in the System KAlSiO 4
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Withers, A.C., Hirschmann, M.M., and Tenner, T.J. The effect of Fe on olivine H 0
2
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Figure Captions
Figure 1. Phase equilibria for Hart and Zindler (1986) primitive mantle composition
with the P-T conditions of experiments from this study (black) and Grove et al. 2006
(gray). Olivine, clinopyroxene and orthopyroxene are stable everywhere in the diagram
with the exception of the highest temperature experiments at 3.2 GPa where
clinopyroxene is absent. The aluminous and hydrous phase stability with the exception
of chlorite breakdown between 2 - 4 GPa are from Ulmer and Trommsdorff (1995),
Schmidt and Poli (1998), and Fumagalli and Poli (2005). The H20-saturated solidus
from this study (3.2-6 GPa) and Grove et al. (2006) (<3.2 GPa) is indicated by the bold
black line.
Figure 2. Electron back-scatter images of experiments bracketing the change in
texture we interpret as the H 20-saturated peridotite solidus. Experiments at <800'C
are homogeneous and fine-grained in texture, whereas experiments at >800'C are more
coarse-grained, porous, and exhibit mineralogical-zoning with olivine and melt
segregated toward the hot (top) and opx and cpx toward the cold (bottom) end of the
capsules (orientation as shown for piston cylinder experiments), garnet is distributed
throughout. Thermal gradients are ~1 0 C across the gold capsule, which is the white area
in each picture. Photos for experiments D179, D183, M338, M337 and D214 (listed CW
from lower left) are all shown at the same scale with the scale bar in M338 equivalent to
900 jim, and photos of experiments C237, D184, D180, D194, D165, and D164 (CW
from lower left) are shown with the scale bar in D164 equivalent to 100 [tm.
Figure 3. Mg# of solid phases in experiments above 8401C at 3.2 GPa. The
increasing Mg# of all solid phases is interpreted to be the result of solid-melt partitioning.
Changes in Mg#s between 840 and 880'C, and 940 and 980'C likely result from the
breakdown of chlorite and clinohumite, respectively. Two sigma errors for the Mg#
based on the electron microprobe compositions and number of analyses are shown for all
phases possible.
Figure 4. Minor element concentrations in garnet, clinopyroxene and
orthopyroxene from experiments at 840-1200*C at 3.2 GPa. Trends in oxide
concentrations with increasing temperature are consistent with partitioning between solid
and melt, rather than solid and a vapor phase. The modeled behavior of an element
during batch melting is illustrated by the bold lines and were calculated using solid-melt
partition coefficients determined from the hydrous melting experiments of Gaetani et al.
(1998; 2003). The larger of the standard deviation or the error determined from the EMP
counting statistics is plotted as the error. Oxide concentrations and their standard
deviations determined from EMP analyses are also given in Tables 2 and 3.
Figure 5. Phase diagram for chlorite peridotite. The illustrated experiments were
conducted using a Hart and Zindler (1986) primitive mantle composition synthesized
from natural peridotite minerals + chlorite for a starting material with a total of 0.7 wt.%
H20 (H & Z natural; Table 1).
Figure 6. Scanning electron backscatter images of experiments D206 and D214.
Both images show the two quench phases with distinct chemical compositions (Table 3)
observed in our experiments above the H20-saturated solidus. We interpret the spherical
features to be quenched of the vapor phase and the wisps to be quench of silicate melt.
Figure 7. Comparison of Subduction-Relevant Solidi, Dehydration Reactions and
Global Slab P-T paths. Slab P-T paths for global subduction zones are plotted from the
numerical models of Syracuse et al. (2010) where full coupling between the slab and
mantle occurs at 80 km depth. The majority of slab surface P-T paths are illustrated as a
gray field and cross the serpentine dehydration curve from Ulmer and Trommsdorf
(1995) (black curve labeled "serp-out"), the chlorite dehydration curve (black curve
labeled "chl-out"), and the H20-saturated peridotite solidus (bold black curve labeled
"peridotite solidus") as determined from our experiments within the range of
temperatures and pressures of the slab surfaces directly below global arcs, as shown by
the overlapping gray boxes (Syracuse et al., 2010). The colored dashed lines illustrate
the P-T paths for a couple of the hottest and coolest subducting slabs. The approximate
location of the solidus for subducted oceanic basalt (Lambert and Wyllie, 1970; Hill and
Boettcher, 1970a; Liu et al., 1996; Vielzeuf and Schmidt, 2001) is shown by the black
dashed line labeled "OC," and the approximate location of the subducted sediment
solidus (Nichols et al., 1994; Poli and Schmidt, 2002; Hermann and Spandler, 2008) is
shown by the dotted black curve labeled "sed." The P-T paths for warm subduction
zones cross the sediment and oceanic basalt solidi in the coolest and shallowest region of
the subarc mantle.
Figure 8. Diagrams of the locations of H20 release and the initiation of melting in
the mantle wedge below arc volcanoes. a) and c) cartoons of the mantle wedge
illustrating chlorite and serpentine stability and the locations where the breakdown of
these hydrous phases produces free H2 0 available to trigger H20-saturated melting in the
mantle wedge. The locations of the isotherms are taken from the geodynamic models of
Grove et al. (2009) using a dip of 30' and a convergence rate of 40 mm yr-'. b) and d)
hydrous undepleted peridotite phase diagrams illustrating the locations where the
breakdown of these hydrous phases produces free H2 0 available that triggers H20saturated melting in the mantle wedge. At warm subduction zones such as Cascadia, the
quantity of H20 fluxed from the slab at subarc depths is greater than that which can be
bound in a -1 km thick layer of chloritized peridotite at the base of the mantle wedge,
and will therefore trigger melting at the temperature of the H20-saturated solidus. The
dehydration of chlorite in the mantle wedge will perpetuate ongoing H20-saturated
melting. In cool subduction zones such as the Northern Marianas, the lesser quantity of
H20 fluxed from the slab at subarc depths can be bound in a -1.5 km thick layer of
chloritized peridotite and melting therefore will not occur until the P-T conditions of
chlorite dehydration.
Table 1. Composition of experimental starting materials.
Table 2. Experimental run conditions and phases present. Mass balance calculations are
included for selected experiments.
Table 3. Normalized electron microprobe analyses of run products from H20-saturated
experiments. "Analyses" indicates number of individual electron microprobe analyses
included in average.
Table 4. Compilation of H20-rich peridotite melting experiments.
|KEY
6-
180
al
5 =-
150
4-
120
0=
90
cl
CL
2
60
1 -
30
0500
0
600
Figure 1.
700
800
1000
900
T (*C)
1100
1200
1300
.
180
120
90
.
60.
30
500
600
700
D164
Figure 2.
0
1200
1300
DU
0.95
--
0.94-
0.93-
0.92-
0.910.90-
0.89-
0.820.81-
2
0.80-
0.79-
0.780.77-
0.76-
0.76-
0.74-
0.73-
0.72 -
700
800
900
1000
1100
Temperature (0C)
Figure 3.
1200
1300
25.00
0
25.00
LI
0
~
0
20.00
20.00
15.00
15.00
0
10.00
10.00
5.00
5.00
I
~u
0.00
800
850
900
950
1000
1050
Temperature(*C)
-
3
1100
__I
1150
1200
0.00
800
850
0.4
-
0.3
2
0.25
1.5
A
p
900
p
P
A
A
A
950
1000
1050
Temperature (*C)
1100
1150
1200
1100
1150
1200
I
0.35
2.5
p
I
0.2
0
0.15
1
{
0.1
0.5
0.05
04
800
&
850
900
1050
950
1000
Temperature("C)
1100
1150
1200
800
850
900
950
a
1000
1050
Temperature (*C)
0.9
0.8
0.7
KEY
0.6
0.5
S0.4
*
Garnet
o
Clinopyroxene
A Orthopyroxene
0.3
-predicted
0.2
0
00
800
Figure 4.
850
-predicted
900
1000
1050
950
Temperature(*C)
Opx-Liq
predicted Cpx-liq
1100
1150
1200
Garnet-liq
180
150
120
go
90~
a.
60
30
0
500
Figure 5.
600
700
800
T (*C)
900
1000
1100
] D206:1060C, 3.2 GPa
Figure 6
I
|D214: 1100*'C, 3.2 GPa
7-200
654100
0
200
400
T (*C)
Figure 7.
600
800
0
1000
Warm Subuction Zones (e.g., Cascadia)
f%
50
100
150
Ii
600
800
Distance (km)
1000
1200
1000
1200
T(*C)
Cool Subuction Zones (e.g., N. Marianas)
50
100
150
600
800
Distance (km)
T(*C)
Key
Figure 8
/*
Free H20 released
from hydrous mineral
breakdown in the slab
Free H20 released
{*
from hydrous mineral
breakdown in the
mantle wedge
Initiation of Melting
Table 1. Composition of starting materials.
H&Z + H20 H&Z natural
46.20
46.03
0.18
0.18
4.06
4.06
0.40
0.36
7.56
7.56
0.10
0.12
37.82
37.81
3.22
3.22
0.33
0.34
0.03
0.28
100.17
99.68
0.90
0.90
SiO2
TiO2
A1203
Cr203
FeO
MnO
MgO
CaO
Na20
K20
NiO
Sum
Mg#
Proportions in H&Znatural
*Pali-Aike Mineral Analyses are from Stern et al., 1989
Pali-Aike xenolith Minerals*
olivine
opx
40.5
56.1
0
0.15
0
3.4
0.06
0.42
9.98
6.32
0.11
0.11
49.3
33.9
0
0.66
0
0.14
99.95
0.90
0.476
101.2
0.91
0.249
cpx
53.2
0.57
5.62
1.07
3.12
0.07
15.6
19.2
2.11
garnet
42.4
0.16
23.7
1.06
8.14
0.31
20.7
4.74
0
New Idria
chlorite
35.29
0.01
12.48
0
3.72
0.18
34.45
0.01
0
100.56
0.90
0.142
101.21
0.82
0.067
86.14
0.94
0.066
Table 2. Experimental run conditions and phases present. Mass balance calculations are included for selected experiments.
Expt.
P
(GPa)
T
("C)
Duration
(hrs)
H20
present?
Phases
ol
opx
cpx
gar
sp
chi
tchu
melt
H20-saturated Experiments (H&Z + H20 starting material)
B1206
D169
D167
D164
D212
D162
D163
D165
D206
D214
D205
D207
D200
D209
D194
D179
D178
D180
D183
D181
D191
D184
D189
D231
D230
M351
M338
M337
M356
M419
M424
1.6
3.2
3.2
3.2
3.2
3.2
3.2
3.2
3.2
3.2
3.2
3.2
3.2
3.2
3.2
3.6
3.6
3.6
4
4
4
4
4
4
4
5
5
5
5
6
6
940
780
840
880
880
940
980
1020
1060
1100
1100
1125
1150
1175
1200
800
820
840
760
800
820
840
920
1050
1100
740
780
840
880
800
830
96.7
48.4*
49*
215.6
6*
212.2
97
95.8
48.2
168.4
114.5
52.5
64.5
27.5
41.8
167.5
168.2
168.8
338.4
168.8
161
159.6
164.2
49
24.5
72
72
72
72
168
168
yes
0.52
0.06
0.21
x
x
x
x
x
x
x
x
0.44
0.33
0.08
0.14
starting material powder + garnet porphyroblasts
x
x
x
x
x
x
x
x
x
x
0.55
0.19
0.11
0.05
0.54
0.23
0.03
<0.01
x
x
0.51
0.26
0.48
0.26
0.48
0.25
0.5
0.16
x
x
x
x
x
x
0.39
0.3
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
x
0.52
0.13
0.18
0.17
x
x
x
x
x
x
x
x
0.46
0.26
0.07
0.19
Chlorite Peridotite Experiments (H&Z Natural starting material)
D193
D196
D168
D192
D166
D202
*intentionally
2.8
850
184
2.8
880
168.8
3.2
780
171.7
3.2
820
168
3.2
880
165.6
3.8
820
166.8
short run times to test equilibration
no
no
no
no
no
no
x
x
x
x
x
xx
0.02
0.14
<0.01
0.01
0.01
<0.01
x
0.09
0.19
x
0.23
0.24
0.25
0.33
Table 3. Normalized electron microprobe analyses of run products from H20-saturated experiments. Analyses indicate number of individual electron
micropro be analyses included in average.
Phase Analyses
SiO2
TiO2
A1203
Cr203
FeO
MnO
MgO
CaO
Na20
K20
NiO
Sum
1.6 Gpa
B1206
oliv
10
opx
11
cpx
12
sp
10
amph
7
glass*
12
40.18
0.31
54.64
0.81
51.97
0.47
0.26
0.13
46.26
1.31
59.34
14.21
0.04
0.04
0.17
0.02
0.56
0.07
0.4
0.05
0.84
0.05
0.44
0.32
0.03
0.04
4.82
0.89
4.37
0.32
45.63
1.72
12.12
0.59
9.99
4.5
40.97
1.24
53.11
1.99
54.09
1.23
14.3
0.53
38.47
1.17
0.04
0.01
0.07
0.02
0.15
0.05
0.33
0.16
4.06
0.16
0.13
0.13
2.53
1.22
1.63
0.51
21.77
1.7
0.17
0.29
40.72
0.59
56.69
0.6
54.39
0.27
41.73
0.06
0.04
0.02
0.12
0.02
0.15
0.02
0.19
0
40.56
0.2
57.48
0.01
0.01
0.3
0.11
0.32
0.05
15.76
2.05
0.47
0.08
9.52
0.28
6.72
0.17
3.01
0.18
18.78
0.73
4.69
0.15
3.15
1.04
0.14
0.02
0.13
0.03
0.09
0.01
0.14
0.03
0.06
0.02
0.08
0.09
50.14
0.23
33.18
0.51
16.55
0.28
18.21
0.62
19.63
0.47
18.2
4.37
0.09
0.02
0.68
0.22
22.62
0.41
0.1
0.06
11.37
0.39
7.29
5.08
0.24
0.08
0.45
0.15
1.74
0.38
0.11
0.02
8.89
0.28
6.22
0.7
2.37
0.31
9.35
0.34
8.15
1.01
0.05
0.03
0.01
0.01
0.06
0.04
0.33
0.04
0.08
0.02
49.17
1.63
36.13
2.29
18.75
1.66
18.22
0.99
48.12
1.54
0.27
0.12
0.72
0.61
22.19
1.04
6.95
1.43
0.27
0.45
0.02
0.03
1.68
0.44
1.16
0.29
22.04
0.54
0.01
0.01
0.37
0.07
0.36
0.14
2.32
0.01
9.05
0.19
6.11
0.37
2.3
0.28
8.92
0.07
0.03
0.03
0.05
0.03
0.02
0.02
0.38
0.05
49.77
0.71
34.5
0.57
18.35
0.66
17.45
0.32
0.03
0.03
0.42
0.05
22.99
0.48
6.03
0.18
0.05
0.05
0.27
0.06
0
0
0.05
0.04
0.05
0.03
0.05
1.3
0.03
0.01
0.22
9.45
0.98
6.01
0.05
0.03
0.07
48.38
1.12
34.09
0.14
0.3
0.44
0
0.21
0.02
54.2
0.34
41.43
0.17
0.01
0.01
0.33
0.15
1.34
0.12
21.85
1.98
0.5
0.04
1.76
0.92
2.55
0.01
8.6
0.11
0.08
0.01
0.29
0.03
17.84
0.17
18.98
1.89
23.13
0.27
6.79
0.88
0.18
0
0.01
0.01
40.81
0.21
56.7
0.72
54.78
0.25
41.46
0.38
0.05
0.03
0.14
0.02
0.2
0.03
0.35
0.07
0.14
0.19
2.4
0.92
1.51
0.38
22.76
0.93
0.05
0.02
0.4
0.23
0.39
0.13
1.43
1.04
9.01
0.23
5.88
0.15
2.49
0.2
8.37
0.29
0.1
0.03
0.13
0.03
0.08
0.02
0.28
0.04
48.63
0.34
32.98
0.44
17.63
0.4
19.05
0.4
0.04
0.02
0.54
0.05
23.04
0.57
6.3
0.37
0.02
0.01
0.27
0.04
0
0
41.89
0.71
56.67
1.17
54.17
0.54
41.27
0.26
59.68
5.69
81.23
7.79
0.07
0.01
0.14
0.03
0.21
0.02
0.39
0.14
0.92
0.3
0.2
0.09
0.07
0.09
1.87
1.22
1.33
0.22
21.76
0.22
14.8
2.44
11.95
3.52
0.07
0.05
0.32
0.18
0.34
0.05
2.3
0.18
0.04
0.03
0.11
0.05
8.83
0.3
5.82
0.3
2.42
0.18
7.77
0.1
4.91
1.14
0.43
0.21
0.07
0.01
0.1
0.01
0.09
0.02
0.33
0.02
0.04
0.02
0.07
0.03
49.83
0.97
34.74
1.02
18.94
0.29
17.57
0.42
16.72
4.29
1.02
0.74
0.1
0.1
0.75
0.08
22.47
0.51
7.16
0.51
1.92
1.29
3.51
1.08
0.01
0.01
0.08
0.02
0.01
0.01
0.67
0.29
0.64
0.31
40.43
0.26
56.5
0.58
54.1
0.5
41.55
1.03
0.02
0.01
0.11
0.01
0.17
0.02
0.39
0.12
0.03
0.02
1.98
1.09
1.26
0.17
21.44
1.22
0.05
0.01
0.36
0.1
0.31
0.07
2.28
0.88
8.53
0.4
5.58
0.23
2.43
0.18
7.62
0.37
0.09
0.02
0.1
0.02
0.07
0.02
0.29
0.04
50.68
0.27
34.72
0.71
18.94
0.52
19.28
1.04
0.04
0.01
0.7
0.05
22.67
0.58
7.15
0.98
3.2 GPa
0.44
0.07
0.04
0.03
0.15
0.02
100.68
99.62
0.57
0.04
1.71
0.08
1.34
1.13
100.59
0.19
0.04
0.18
0.11
99.93
97.35
73.75
D1 64
oliv
14
opx
8
cpx
5
gar
7
tchu
4
oliv
19
opx
8
cpx
16
gar
2
0.51
0.12
0.95
1.02
0.33
0.14
0.02
0.03
D1 62
0.01
0.01
99.65
99.75
100.14
100.27
0.47
0.03
97.45
0.43
0.11
100.07
100.23
100.15
99.06
D1 63
7
1
2
6
D1 65
5
5
9
7
D206
oliv
3
opx
6
cpx
7
gar
3
glass
19
vapor quench
10
99.3
99.67
99.83
100.03
99.49
99.17
100.38
100.01
0.15
0.15
101.08
100.43
100.04
98.57
61.62
72.4
D214
oliv
20
opx
16
cpx
14
gar
19
0.13
0.04
0.003
0.005
0.06
0.03
0.01
0.01
100.48
100.83
101.15
100.03
MnO
MgO
CaO
0.11
0.03
0.14
0.01
0.12
0.02
0.29
0.05
50.59
0.67
34.59
0.31
18.38
0.18
18.4
0.17
0.03
0.01
0.69
0.05
22.37
0.68
6.67
0.49
8.34
0.09
5.67
0.82
20.63
0.16
0.12
0.03
0.13
0.02
0.17
0.01
49.72
0.61
35.85
4.77
15.37
0.65
0.05
0.01
0.61
0.16
0.09
0.08
0.07
0.13
0.62
0.21
52.36
1.01
7.67
0.95
5.22
0.2
16.7
1.03
0.09
0.02
0.12
0.03
0.2
0.02
49.43
5.09
35.47
0.56
15.15
0.15
0.08
0.16
0.54
0.08
0.03
0.01
0.08
0.01
0.82
0.25
7.15
0.35
5.34
1.08
0.1
0.01
0.12
0.06
51.21
0.63
34.4
1.8
0.05
0.01
0.73
0.48
0.13
0.02
0.09
0.02
51.15
0.32
34.99
0.35
0.05
0.01
0.54
0.09
0.16
0.03
446.1
1.8
0.1
0.19
0.1
0.02
0.03
18.37
2.13
9.64
20.68
1.88
16.57
0.23
0.05
0.96
0.11
0.07
SiO2
TiO2
A1203
Cr203
41.39
0.22
57.9
0.42
55.55
0.25
41.92
0.02
0.03
0.01
0.11
0.02
0.18
0.01
0.35
0.11
0.03
0.02
1.75
0.2
1.45
0.2
22.27
0.18
0.02
0.01
0.33
0.02
0.37
0.04
2.5
0.76
41.32
0.2
54.9
4.7
0.4
0.19
0.01
0.01
0.03
0.01
0.33
0.01
0.01
0.01
2.14
0.72
19.24
0.01
0.05
0.02
0.64
0.18
43.48
1.4
42.58
4.87
56.85
0.69
0.42
0.42
0.02
0.01
0.03
0.01
0.26
0.01
0.13
0.4
1.62
0.32
15.49
0.58
41.22
0.41
56.86
0.91
0.02
0.01
0.03
0.03
0.03
0.06
1.7
0.69
41.2
0.22
57.22
0.73
0.02
0.01
0.01
0.01
0.01
0.01
1.62
0.12
12
40.95
1.38
0.07
0.04
0.14
0.15
0.02
0.02
5
54.45
0.88
44.69
0.05
0.01
0.07
1.58
0.79
24.94
0.13
0.06
0.1
Phase Analyses
FeO
Na20
K20
NiO
Sum
0.14
0.05
101.06
D205 (cont)
oliv
9
opxa
7
cpx
4
gar
2
oliv
13
opx
14
101.11
0.01
0.01
0.08
0.03
0.03
0.04
100.79
98.31
D207
sp
2
0.4
0.04
0.03
0.05
99.67
99.77
0.18
0.01
96.73
0.22
0.07
100.28
D200
oliv
11
opx
15
sp
5
100.47
0
0.01
0.01
0.01
100.67
0.14
0.05
100.35
D209
oliv
11
opx
14
sp
0.001
0.003
100.38
na
D194
oliv
36
opx
10
0.24
0.03
0.02
0.01
99.94
101.10
sp
3.6 Gpa
D179
oliv
11.92
2.21
0.53
0.05
100.6
opxa
1
chl
11
32.72
1.36
0.02
0.01
16.21
1.01
1.47
0.3
4.15
0.18
0.04
0.02
30.16
0.85
0.2
0.27
oliv
7
41.32
0.18
0.11
0.1
0.07
0.07
0.06
0.03
8.89
0.8
0.07
0.01
49.25
0.43
0.04
0
4
53.24
0.65
42.65
1.17
0.06
0.01
0.68
0.26
2.49
0.77
18.63
1.16
0.11
0.13
0.26
0.05
17.97
2.68
17.47
1.84
20.82
1.82
10.04
2
11
41.62
1.31
0.09
0.12
0.14
0.13
0.08
0.02
9.22
0.98
0.09
0.04
48.2
1.26
0.08
0.07
cpx
8
gar
5
54.13
0.79
44.53
1.39
0.06
0.02
0.35
0.07
1.73
0.58
17.28
2.22
0.15
0.05
1.74
0.53
4.06
1.14
9.26
0.31
0.13
0.05
0.27
0.06
20.2
4.63
18.52
1.85
19.37
5.26
8.03
1.41
oliv
11
opx
6
cpx
10
11
0.15
0.25
0.1
0.13
0.06
0.04
0.02
0.03
0.15
0.12
1.77
0.79
1.62
0.55
17.46
0.64
0.07
0.03
0.13
0.09
0.16
0.06
chl
40.6
0.7
54.29
1.91
53.96
0.68
40.03
1.46
0.16
0.03
0.11
0.03
0.15
0.08
0.03
0.01
Mg-carb
9
11.04
0.66
7.89
1.71
4.46
1.01
4.88
0.2
14.87
1.52
47.27
1
31.61
1.3
19.64
2.64
36.88
0.68
81.53
1.26
0.06
0.03
3.89
1.78
19.67
3.13
0.27
0.29
3.6
0.39
100.1
100
85.11
0.03
0.02
D180
0.48
0.07
100.28
opxa
6
0.28
0.02
0.11
0.05
99.82
100.5
4 GPa
D181
oliv
0.48
0.08
100.09
opxa
0.16
0.07
0.05
0.02
98.88
100.7
D183
0.49
0.09
0.2
0.05
0.29
0.07
0.43
0.35
100.62
99.88
99.99
85.82
43.3
Phase
Analyses
SiO2
TiO2
A1203
Cr203
FeO
MnO
MgO
CaO
oliv
9
6
cpx
5
gar
7
41.42
0.99
56.27
0.72
54.37
0.41
44.78
1.14
0.06
0.02
0.1
0.03
0.02
0
0.37
0.11
0.05
0.05
3.66
0.58
1.89
0.16
15.76
2.38
0.07
0.02
opx
8.85
0.72
5.68
0.31
3.6
0.71
8.74
0.5
0.06
0.02
0.1
0.03
0.06
0.02
0.22
0.05
49.42
1
33.55
0.72
22.28
1.69
20.11
3.52
0.07
0.05
0.56
0.16
17.19
2.47
8.4
1.92
oliv
17
41.15
0.85
57.74
0.87
41.8
0.27
1.53
0.67
0.05
0.07
0.07
0.01
0.3
0.09
0.04
0.03
0.04
0.15
1.51
0.72
21.55
0.71
0.64
0.36
0.05
0.07
0.23
0.02
2.55
0.45
0
8.73
0.44
5.51
0.15
7.77
0.22
7.9
1.98
0.09
0.03
0.09
0.03
0.29
0.04
0.54
0.08
49.72
1.57
34.1
0.9
19.03
0.96
8.24
1.58
0.06
0.05
0.74
0.13
6.71
0.54
81.06
6.06
Na20
K20
NiO
Sum
0.52
0.05
0.07
0.03
100.52
D1 91
0.16
0.03
1.38
0.35
0.1
0.03
0.03
0.02
99.6
99.69
99.78
D230
opx
5
gar
14
carb
5
oliv
8
41.1
0.52
0.1
0.12
0.01
0.02
0.09
0.02
8.92
0.35
0.06
0.02
48.1
0.9
0.15
0.3
0.11
0.09
0.01
0.01
0
0.03
0.02
100.59
100.99
100.69
0.03
0.01
55.76
5 GPa
M338
0.8
0.1
99.32
opxa
2
56.16
gar
10
0.6
41.11
1.07
0.03
0.02
0.45
0.08
0.67
0.22
21.09
1.93
0.1
0.02
1.36
0.26
5.35
1.79
8.73
0.85
0.17
0.13
0.3
0.02
23.54
1.56
15.72
1
12.49
0.01
10.85
1.13
oliv
13
opx
11
41.45
0.48
56.41
2.28
56.61
0.06
0.04
0.05
0.07
0.03
0.01
0.02
0.58
0.44
0.57
0.06
0.04
0.23
0.06
0.22
8.61
0.14
5.36
0.66
4.36
0.09
0.03
0.08
0.03
0.12
48.55
0.73
36.93
2.39
26.47
0.04
0.01
0.34
0.37
10.25
0.04
0.03
0.31
99.85
1.16
0.58
8.74
1.02
0.34
0.05
17.41
0.35
7.89
1.31
0.03
0.02
100.55
8.63
0.31
5.07
0.25
2.94
0.54
9.5
0.94
0.05
0.02
0.06
0.02
0.1
0.06
0.28
0.05
49.55
0.81
35.49
0.45
20.71
1.27
17.49
2.44
0.06
0.03
0.32
0.27
20.23
1.83
8.99
1.25
cpx"
98.79
99.62
M337
cpxb
1
gar
9
41.42
0.36
0.35
0.1
23.2
0.7
oliv
10
opx
7
cpx
6
gar
7
41.06
0.57
58.54
0.24
55.54
0.96
41.52
0.83
0.02
0.02
0.01
0.01
0
0
0.37
0.1
0.07
0.05
0.4
0.19
0.42
0.09
20.51
1.8
1.3
0.31
oliv
13
40.68
0.82
58.3
0.99
55.43
0.61
41.15
0.47
0.08
0.16
0.02
0.02
0
0
0.29
0.11
0.01
0.02
0.17
0.06
0.21
0.08
20.64
1.55
0.06
0.02
0.14
0.08
0.09
0.04
1.62
0.52
0.78
0.12
99.66
98.94
6 GPa
M419
0.55
0.05
0.1
0.04
0.07
0.03
0.04
0.05
99.46
99.45
99.4
100.92
M424
8.43
0.05
50.31
0.03
0.35
100.75
0.46
0.03
0.68
0.03
0.17
opx
13
5.39
0.09
35.67
0.02
0.18
101.39
0.34
0.06
0.51
0.03
0.09
cpx
4
2.41
0.08
19.33
22.36
0.09
100.17
0.34
0.02
1.57
2.92
0.02
gar
11
10.51
0.35
18.19
7.24
0.01
100.42
0.92
0.05
2.37
1.68
0.01
carb
9
9.74
89.73
0.54
43.47
*Quenched melt is small wisps of intergranualar material that are difficult to analyze and composition likely reflects the comoposition of quench crystals
around and below the wisp analyzed
aPyroxenes in these samples are very small and we were not able to analyze any points with confidence
bPyroxenes are small and difficult to analyze and therefore composition listed represents single points of pyroxene with lowest Al and Ti contents
69
Chapter 2:
Compositions of Low Extent Melts of H2 0-Saturated
Peridotite
Abstract
We present a study of the nature and composition of quench products in a subset of
the super-solidus H 20-saturated peridotite melting experiments of Till et al. (2011)'.
Electron microprobe analyses of the quench products substantiate the presence of both a
low alkali tholeiitic melt and a high-silica vapor-phase quench. When multiple filtering
techniques are applied to the tholeiitic quench analyses, we are able to infer equilibrium
melt compositions at 3.2-4 GPa. These melt compositions suggest a switch in melting
reaction at clinopyroxene- and garnet-out (~ 125'C, 3.2 GPa), from oliv + cpx+ garnet
4 opx + melt, to oliv+ opx 4 melt. The isobaric melt production rates in these
experiments are extremely low (dF/dTp = 0.05-0.13%/'C) and may represent a minimum
in peridotite melt production rates, which occur when a melt is H20-saturated. These
melting reactions and melt production rates compare favorably with previous studies of
hydrous and anhydrous peridotite melting behavior and are compatible with the solidus
temperature of -800'C at 3.2 GPa interpreted by Till et al. (2011). Further work is
required to refine these estimates of the equilibrium melt compositions and determine the
cause of anomalously high KDM-FeO1iv-Liq values present in a number of the experiments.
These preliminary results support previous interpretations that low extent melts of H2 0saturated peridotite at 3-4 GPa are the forefathers of magmas erupted at volcanic arcs.
1 Same as Chapter 1 in Till, C.B., 2011, Melt Generation in the Earth's Mantle at Convergent
Plate Margins, MIT PhD Thesis.
1. Introduction
Scientists have long sought to understand the link between convergent margins and
the formation of hydrous andesitic magmas erupted at volcanic arcs (e.g., Gill, 1981).
Geochemical and experimental studies of the erupted lavas, quenched magmatic
inclusions, and mineral-hosted melt inclusions found at volcanic arcs have been
successful in reconstructing the composition of parental arc magmas (e.g., Yogodzinski et
al., 1995; Jicha et al., 2009). Parental mantle-derived magmas are found to be primitive
magnesian andesites and to a lesser extent basaltic andesites that contain ~52-58 wt.%
Si0 2, 9-11 wt.% MgO, a minimum of ~4.5-6 wt.% H2 0 and trace element abundances
inherited from slab-derived fluids (e.g., Grove et al., 2002; 2005). These melts tend to
experience fractional crystallization in the shallow crust to produce the more evolved
andesites and dacites that often form the bulk of arc volcanoes (Grove et al., 2005;
Krawczynski and Grove, 2011 a). Although the crustal history of parental arc magmas
can be well constrained, considerable uncertainty remains about the processes that lead
their generation in the mantle prior to their emplacement in the base of the crust.
It is generally accepted that partial melting of peridotite in subduction zones is
initiated by an influx of volatiles from the subducted ocean lithosphere (e.g., Gill, 1981;
Tatsumi, 1986; Davies and Stevenson, 1992). However the specifics of the mantle
melting process, as well as the composition of the hydrous partial melts generated, have
been the subject of experimental investigations and scientific debate for more than 30
years. Much of our understanding about these melts comes from the early experimental
work of Kushiro and co-workers, which focuses on using simplified compositional
systems, such as Mg 2SiO 4 -SiO 2-H 20 (Kushiro et al., 1968; Kushiro 1969, 1972). This
work suggests hydrous peridotite melting produces Si0 2-rich melt; consistent with the
hypothesis that andesite magmas erupted at arc volcanoes are the descendants of partial
melts of hydrous mantle peridotite (Poldervaart, 1955; Kushiro, 1972). Later
experimental studies on natural peridotite compositions indeed produce hydrous glasses
high in Si0 2 and low in FeO and MgO (Kushiro et al., 1972; Mysen et al., 1974; Nehru
and Wyllie, 1975) but raised questions regarding the modification of melt composition
via growth of amphibole and pyroxene upon quench (e.g., Green, 1973; 1976). Careful
experimental studies by Gaetani and Grove (1998, 2003) at 1100-1345'C and 1.2-2 GPa
demonstrate that the addition of 3-6 wt.% H20 in spinel lherzolite melts increases SiO 2
by -1 wt.% and decreases FeO + MgO by -2 wt.% relative to analogous anhydrous
peridotite melts because of the lower temperature of melting in the presence of H20.
Their results also support the interpretation that arc parental magmas with high preeruptive H20 contents form when fluid-saturated peridotite partial melts percolate
upward in the mantle wedge and maintain equilibrium with the hotter overlying peridotite
(i.e., reactive porous flow).
Within the last five years, new petrologic evidence has arisen that resolves the long
standing discrepancy in the solidus temperature of the H20-saturated mantle (BVSP,
1981) and demonstrates that mantle peridotite melts at extremely low temperatures when
in the presence of excess water. New long-duration high-pressure mantle melting
experiments by Grove et al. (2006) and Till et al. (2011) indicate that the H 20-saturated
peridotite solidus is located between 8000 and 820'C at 3-6 GPa. Likewise, studies of a
peridotite massif exposed in the Japanese Sanbagawa belt (Hattori et al., 2009; Till et al.,
2009) find evidence for hydrous melting processes at -800'C and -3 GPa in a paleomantle wedge. Consequently, analysis of the composition of low extent melts produced
in the experiments of Till et al. (2011) can shed new light on the composition of hydrous
partial melts and the specifics of hydrous melting behavior in the mantle wedge at
subduction zones.
Quenching experimental melts with a high H20-content has long been recognized to
be problematic, as the exsolution of H2 0 during the rapid cooling and depressurization of
the experiment causes the melt to quench as a froth of bubble walls and/or as "quench"
mineral phases, rather than as a glass. Here we present a study of the nature and
composition of quench products in a subset of the Till et al. (2011) experiments where
they were of sufficient size and quantity to be chemically analyzed on an electron
microprobe. We find evidence for multiple quench phases and discuss methods for
filtering the chemical analyses of the quench products. The melting reactions and rates of
magma production implied by the quench compositions are discussed. We also explore
additional techniques to determine an equilibrium melt composition.
2. Experimental and Analytical Methods
Till et al. (2011) present a comprehensive study of piston cylinder and multi-anvil
experiments conducted between 3.2 - 6 GPa and 740-1200'C on the primitive upper
mantle composition of Hart and Zindler (1986) in the presence of excess H20 (Table 1).
The experimental and analytical methods for these experiments are discussed in detail in
Till et al. (2011). The starting material for these experiments was created from a mixture
of high-purity synthetic oxides or metasilicates. Brucite, Mg(OH) 2 was added to the
synthetic oxide mixture instead of MgO, which provided the starting material with ~14.5
wt.% H2 0.
The experiments considered here were conducted at the MIT Experimental Petrology
Laboratory in a 0.5" solid medium piston cylinder apparatus (Boyd and England, 1960)
using the hot piston-in technique (Johannes et al., 1971). The experiments were
conducted in Au capsules to minimize the iron loss and prevent hydrogen loss from the
starting material. The pressure medium was BaCO 3 and pressure was calibrated using the
anorthite + gehlenite + corundum = Ca-tschermak pyroxene reaction and the spinel to
garnet transition in the CMAS system (Hays, 1966; Longhi et al., 2005). Experimental
pressures are thought to be accurate to ± 0.05 GPa. Temperature was monitored and
controlled using W 97Re 3-W75Re 25 thermocouples with no correction applied for the effect
of pressure on the thermocouple EMF. The run temperatures for the experiments are
accurate to ± 10 C (Medard et al., 2008). To raise the piston cylinder to pressure,
experiments were pressurized to 1 GPa at room temperatures then the temperature was
raised at 100 C/min to 865'C. The experiment was held at these conditions for 6 minutes
before raising the temperature at 50'C/min to the final temperature and pressure. The
sample was held at isothermal and isobaric conditions for the duration of the experiment
and then quenched by turning off the power. Following each experiment a hole was
drilled in the capsule with a small diameter hand drill and the presence or absence of
liquid water was noted. The capsule was then sliced in half, dried, vacuum-impregnated
with epoxy and polished for electron microprobe analysis.
The major element concentrations of all experimental products were analyzed by
wavelength dispersive spectrometry on the 4- or 5-spectrometer JEOL 733 and 8200
electron microprobes (EMPs) at the Massachusetts Institute of Technology. All analyses
of minerals were conducted with a 15 kV accelerating voltage, a 10 nA beam current and
a 1 pm spot size. Quenched melt analyses were conducted with a range of beam currents
(3 to 10 nA), and spot sizes (1-30 pm) and measurement times (5-40 sec), depending on
the size and fragility of the quench product. Online data reduction utilized the CITZAF
correction package (Armstrong, 1995) and the atomic number correction, the absorption
coefficients, and the fluorescence correction available in CITZAF.
3. Determination of Quench Compositions
Experimental determinations of the H 20 solubility in silicate melt indicate at 30 kbar,
27.0 ± 1.0 wt.% H 20 is required to saturate forsterite melt and 21.5 ± 1.0 wt.% H2 0 to
saturate diopside melt (Hodges, 1974). This data suggests the olivine(±diopside)saturated melts in the Till et al. (2011) experiments are H20-saturated until >55% melted
and will suffer from extensive H20-exsolution during depressurization for the entire
range of melting extents relevant to the lower half of the mantle wedge (<50%). In near
solidus experiments, the melt wisps tend to coat the crystalline mineral phases making
them especially difficult to chemically analyze on the EMP. However, in the higher
temperature experiments (>250-300'C above solidus) of Till et al. (2011), there is a
sufficient fraction of melt such that it segregates toward the hot end of the capsule
making EMP analyses of the quench products possible (Figure 1). Therefore we focus on
this subset of the Till et al. (2011) experiments conducted between 1 100 C and 1200'C at
3.2 and 4 GPa. The compositions of the quench products in these experiments are
illustrated in Figure 2.
In experiments between 1100-1200'C, EMP major element totals for the quench vary
between 48-100 wt.%. Totals of 65-100 wt.% may reflect the volatile content of the
quench products. We dismiss analyses with totals <65 wt.% from consideration here, as
they are likely poor analyses resulting from the high aspect ratio of the individual melt
wisps and the size of the EMP beam (1-30 pm) or the EMP beam puncturing a thin layer
of melt and penetrating a crystalline mineral phase. The quench with totals >65 wt.% are
normalized on an anhydrous basis and the resulting average quench composition for each
experiment is presented as "Filter 1 Results" in Table 1. These quench compositions
exhibit a significant amount of intra-experiment variability, as well as inter-experiment
variability (Figure 2, Table 1). Several compositional filters are applied to these data in
an attempt to hone in on the compositions of the melts in equilibrium with the crystalline
mineral phases.
The exchange coefficient for iron and magnesium between olivine and silicate liquid
(KDMg-FeO 1iv-Liq)
is widely used to assess equilibrium between phases in experimental
studies, as well as for natural minerals and melts. KD Mg-Fe Oliv-Liq is commonly assumed to
have a constant value of 0.3 over a wide range of temperatures, pressures and
compositions, as originally proposed by Roder and Emslie (1970).
The KDMg-FeOliv-Liq for
quench analyses with totals >65 wt.% varies between 0.10-0.80 (Figure 3). Published
values for KDMg-FeOJiv-Liq in experiments are between -0. 17-0.45 (Toplis, 2005) and the
extent to which KDMg-FeO1 iv-Liq is actually constant has been a subject of debate (e.g.,
Mysen 1975; Toplis, 2005). Pressure, temperature, olivine composition, melt
composition, and water content are potential sources of variation in KDMg-FeOliv-Lig (e.g.,
Ulmer, 1989; Kushiro and Walter, 1998; Toplis, 2005). Toplis (2005) illustrated that the
effects of increasing temperature and changing olivine composition tend to counteract
one another as melting proceeds. The effect of pressure on KDMg-FeOliv-Liq is fairly
minimal (0.01/GPa), which equates to an acceptable range of 0.33-0.34 in the
experiments considered here. Melt composition is a likely source of variation in KDMgFeOliv-Liq
in the Till et al. (2011) experiments. Melt composition can account for variations
in KDMg-FeOliv-Li between -0.2-0.42 (Toplis, 2005). As such, we use these as maximum
and minimum values for KDMg-FeOliv-Liq tO filter the analyses of the quench with totals >65
wt.%, since the correct value for a given experiment is unknown. The remaining quench
compositions are normalized to 100%, averaged for each experiment, and presented as
"Filter 2 Results" in Table 1. Experiments by Gaetani and Grove (1998) and Ulmer
(1989) suggest that H20 likely increases the value of the KDMg-FeOliv-Liq relative to
anhydrous melting experiments (Toplis, 2005). The H20-content of the experiments
examined here is constant and would therefore affect all of the experiments uniformly.
But the magnitude of this H20-effect is uncertain and we choose not to interpret quench
analyses with KDMg-Fe0
liv-Liq>
0.42 as equilibrium melts.
In addition, the composition of the quench products was compared to those of the
crystalline mineral phases in a given experiment (Figure 2). Humayun et al. (2010)
demonstrate that laser ablation ICP-MS (LA-ICP-MS) analyses can yield major and
minor element compositions of similar or better accuracy and precision than EMP
analyses of USGS glass reference materials. They also collect transects of major element
LA-ICP-MS analyses across large pools of the quench products in anhydrous partial
melting experiments of KLB-1 peridotite at 3 GPa and 1450'C. The LA-ICP-MS results
illustrate that analyses dominated by crystalline minerals form excellent correlations in
MgO vs. SiO 2 plots, whereas analyses dominated by quench do not exhibit simple mixing
lines between the compositions of olivine and pyroxene, as determined by EMP analysis.
When our EMP analyses of the quench products are compared on MgO vs. SiO 2 plots,
they too reveal that a number of the analyses appear to plot on mixing lines between a
crystalline mineral and what is presumably the equilibrium melt composition. We adopt
the technique of Humayun et al. (2011) and remove EMP analyses we interpret as
analysis sampling crystalline phases from the dataset and the resulting average quench
compositions are presented under "Filter 3 Results" in Table 1.
There is a significant variability in the average quench composition produced by the
three filters, especially for the 1100 and 1200'C experiments at 3.2 GPa. Overall, the
average quench compositions can be described as silica-undersaturated low-alkali olivine
tholeiites, as evidenced by their positions on the oxygen-normalized pseudo-ternary
projection scheme of Tormey et al. (1987) (Figure 4). Because of the variability in the
average quench compositions, an additional technique for determining the quenched melt
compositions was explored. For this technique, the quench products were carefully
removed from half of experiment D194 (1200'C, 3.2 GPa) and the remaining half
capsule was mounted in epoxy and analyzed on the microprobe. The removed quench
was melted in a platinum foil packet in a 1 atm furnace at -1500 C for 3 hours and
quenched to an anhydrous glass. The composition of this anhydrous glass is very similar
to the hydrous glass for experiment D194 (Table 1), with slightly higher A120 3 , FeO,
Na 2O, K2 0 and lower TiO 2 , Cr 2O 3 and CaO when both are normalized on an anhydrous
basis. This remelted quench composition also plots in a location on the Oliv-Cpx-Qtz
ternary consistent with the formation of a melt similar to the bulk composition at
increased extents of melting. Because of their internally consistent variations in
composition with increasing temperature, the average melt compositions produced by
Filter 3 at <1200'C and the remelted quench at 1200'C are our best estimates of the
equilibrium melt compositions (referred to hereafter as the "preferred melts").
In addition to the tholeiitic quench, there is evidence for a vapor phase quench in the
experiments. Spheres found in the voids between quenched melt wisps in experiments at
1 100 C (Figure 1) produce compositions with >70 wt.% SiO 2 at <2 wt.% MgO (Table 1).
There is chemical evidence for this high-silica quench phase in experiments at 1100,
1 125'C and 1 175'C at 3.2 GPa, where several quench analyses plot on mixing lines
between the high-silica composition and what we interpret to be the equilibrium melt
composition. Instead of the high-silica spheres, there is a carbonate quench phase present
in the voids spaces between the quenched melt at 4 GPa (Figure 1). The 4 GPa
experiments were conducted more than a year and half after the preparation of the
starting material and the carbonate quench is likely the result of progressive carbonation
of the synthetic oxide mixture over time by atmospheric C0 2, as observed in other
experiments where the starting material is not kept under vacuum (M. Schmidt, personal
communication, 2010). This small amount of CO 2 in the starting material (<100 ppm) is
not sufficient to change the solidus temperature of peridotite by more than 20'C
(Dasgupta and Hirschmann, 2007). We interpret the spheres and carbonate as quench
products of a vapor phase that was either present with the melt phase at pressure or
exsolved during cooling, similar to the interpretation of White and Wyllie (1992).
4. Discussion
4.1 PhaseProportions& Inferred Melting Reactions
The proportions of the crystalline phases and melt in each experiment were calculated
by fitting the bulk composition of the starting material to a linear combination of the
analyzed phase compositions. The phase proportions produced using the three different
quench filters are illustrated in Figure 5, in addition to the proportions of the equilibrium
minerals calculated without a melt composition for experiments at <11 00 C and 3.2 GPa
from Till et al. (2011). Both the ternary positions and the phase proportions produced
with our preferred melt compositions (Filter 3 compositions at 1100-1 175'C and the
1200'C remelted quench composition) imply a melting reaction prior the exhaustion of
clinopyroxene and garnet (<1 100 C) where olivine, clinopyroxene and garnet are
consumed to produce melt and orthopyroxene. This is illustrated in Figure 6, where the
tie line between the 1 100 C orthopyroxene composition and the 1 100 C average
quenched melt composition pierce the plane defined by coexisting clinopyroxene, olivine
and garnet (the garnet in these experiments plots above the Plag apex of the Olv-Cpx-Qtz
pseudo-quaternary). The proportions of the crystalline phases in all the experiments at
<1 100 C have also been renormalized in Figure 6 to include the predicted proportion of
melt. At temperatures above 11 00 C, only a harzburgitic residue remains and the location
of the preferred melts on the pseudoternary suggest olivine and orthopyroxene are
consumed in a eutectic melting reaction.
4.2 Comparison to PreviousExperimental Studies
The preferred melts are lower in FeO and MgO and higher in SiO 2 and CaO than
similar extent anhydrous pyrolite melts at 3 GPa (Walter, 1998) (Figure 7; Table 1). The
increased SiO 2/FeO+MgO ratio in the hydrous preferred melts is a result of the lower
melting temperature relative to anhydrous peridotite as demonstrated by Gaetani and
Grove (1998; 2003) and Parman and Grove (2004). The melting reactions and the switch
in melting reactions at clinopyroxene-out observed here are identical those observed in
anhydrous melting experiments at 3 GPa, as well as anhydrous melting experiments at
lower pressure (Kinzler and Grove, 1997; Longhi, 2002). This similarity in melting
behavior is encouraging, as it lends credibility to the equilibrium melt compositions and
melting reactions we interpret.
The fact that our melt compositions are silica-undersaturated tholeiites is significant.
Early work on low extent H20-saturated peridotite melts by Kushiro (1968, 1969, 1972)
suggests that melts coexisting with olivine, orthopyroxene, clinopyroxene (± amphibole)
at 0.75 GPa, 1000 0 C and 1.95 GPa, 11 000 C are silica-saturated quartz tholeiites (-60
wt.% SiO 2 ,Mg# 0.79-0.83) with compositions similar to calc-alkaline andesites (Figure
7). The composition of H20-saturated peridotite melts in equilibrium with olivine,
orthopyroxene, clinopyroxene and spinel at 1.6 GPa, 940'C in the Till et al. (2011) study
and at 1.2 GPa, 1020'C by Grove et al. (2006) are both also silica-saturated quartz
tholeiites (58-59 wt.% SiO 2, Mg# 0.75-0.91). These observations suggest that low-extent
melts at lower pressures are silica-saturated but become increasingly silicaundersaturated at higher pressures.
4.3 Melt Productivity
The melting extents predicted by our preferred melt compositions suggest the rate of
isobaric melt production (dF/dT) pat 1100-1200'C and 3.2 GPa is -0. 13% melt/0 C
(Figure 8). Gaetani and Grove (1998) also calculate (dF/dT)p = ~0.05%/'C at 2% partial
melt for a peridotite with 0.15 wt.% H20 at 1.5 GPa and find the productivity increases
continuously to 0.80%/'C at 20% partial melt. Studies of anhydrous peridotite melting
determine an average melt productivity of -0.30%/'C for ~20-40% partial melt at both 1
and 3 GPa (Baker and Stolper, 1994; Walter, 1998; Falloon and Danyushevsky, 2000)
and values of 0.8-1.27%/'C for fertile peridotite at -2 GPa (Kinzler, 1997). These
observations indicate the concentration of dissolved H20 in the melt exerts an important
control on melt productivity. The extremely low melt production rate determined here
for the H20-saturated 20-30% melts of Till et al. (2011) and the H20-saturated 2% partial
melts of Gaetani and Grove (1998), suggest that a melt production rate of~0.050.13%/'C may be a minimum bound on peridotite melting rates common to H2 0saturated melts. These low rates of melt production suggest a significantly longer
melting column is required to produce the same amount of partial melt generated in
anhydrous partial melting environments. When extrapolated from the 1100-1200'C
experiments of Till et al. (2011) to 0% melt, the H20-saturated melt productivity predicts
a H20-saturated solidus temperature of -800'C, consistent with that documented by Till
et al. (2011) (Figures 6 & 8).
4.4 Comparisonof the Experimental Melts to Arc Lavas
H20-saturated partial melts formed at the base of the mantle wedge buoyantly ascend
into the hotter overlying mantle, where these superheated melt will re-equilibrate by
dissolving the surrounding lower pressure peridotite minerals (Grove et al., 2002). As
these melts subsequently ascend through the top half of the wedge and lower crust, they
will experience fractional crystallization and to varying extents, assimilation and magma
mixing. Krawczynski and Grove (2011) conducted an experimental study of mantle-
derived parental arc magma compositions from Mt. Shasta in the presence of H 20 that
demonstrates near liquidus crystallization of amphioble and clinopyroxene and olivine at
lower crustal pressures (500-800 MPa) can produce calc-alkaline liquids with
compositions typical of arc lavas.
Our preferred melt compositions are compared to one of the parental arc magmas
studied by Krawczynski and Grove (2011) in Figure 9, as well as quenched mafic
inclusions and calc-alkaline basalts-andesitic lavas from Cascades arc volcanoes. The
preferred melts compositions are only slightly lower in silica and alkali contents than the
parental arc magma and the most mafic of the arc lavas. If the preferred melt
compositions are approximations of the melts formed at the base of the mantle wedge,
their ascent to the shallow crust will likely cause them to evolve to the more SiO 2-rich
compositions typical of arc magmas. Thus the melts of H20-saturated peridotite at -3
GPa generated in the experiments of Till et al. (2011) may be the forefathers of arc lavas.
4.5 Where Do We Go From Here?
Although we are able to determine potential equilibrium melt compositions from the
quench products in the experiments of Till et al. (2011) and make cogent inferences about
H20-saturated peridotite melting behavior, there is still uncertainty regarding the exact
composition of the equilibrium melt. The three different methods for filtering the quench
compositions produce average compositions that vary by as much as 6 wt.% SiO 2 and 11
wt.% MgO for the experiment at 1 100'C, and 1 wt.% FeO and 4 wt.% CaO for the
experiment at 11 75'C, for example. As a result, several of the average quench
compositions plot at locations on the pseudoternary diagrams that are untenable, as well
as inconsistent with the other quench compositions (e.g., the Filter 1 1100 0 C average
quench). Na 20 is highly volatile and notoriously difficult to analyze on the EMP. Our
mass balance calculations suggest that up to 75% of the bulk Na 2O was lost in the EMP
analyses of the experimental phases, the majority likely from the quench analyses.
One explanation for these discrepancies could be that the analyzed quench in these
experiments is not an equilibrium melt, but instead represents a complicated super-critial
fluid with a high dissolved silicate content. Till et al. (2011) demonstrate that the
changes in the composition of the equilibrium minerals in these experiments with
temperature could only be the result of solid-melt partitioning and not solid-fluid
partitioning. The presence of both the high-silica quench and the more mafic quench
wisps further suggests these experiments are below the second critical end point, which is
poorly constrained in the peridotite-H 20 system.
Therefore to determine the equilibrium melt composition of 1 20-saturated peridotite
at pressures >3 GPa with greater precision and accuracy, further work is required.
Experimentalist have developed and struggled with a number of methods to segregate and
analyze low extent (<10%) hydrous melts, these methods include sandwich experiments,
melt traps, and thermal segregation (Stopler, 1980; Takahashi and Kushiro, 1983; Hirose
and Kushiro, 1994; Baker and Stolper, 1994; Walter, 1998; Laporte et al., 2004). The
simplest of these, the sandwich experiment, consists of a layer of the assumed melt
composition (the meat) to be surrounded by two thicker layers of the bulk rock
composition (the bread). During the experiment, the melt will approach equilibrium with
the minerals in the bread through diffusion and dissolution and may quench to a
homogenous material that can be chemically analyzed at the end of the experiment.
Several problems can arise using this method, as the length scales of the capsule can be
too great and the dissolution rates too slow to reach chemical equilibration, a layer of
reaction minerals can form between the two materials, impeding chemical exchange
and/or the addition of the melt can shift the bulk composition of the experiment such that
the resulting melt composition is not in equilibrium with the original bulk composition.
An iterative version of the sandwich experiment, where the melt composition from each
experiment is used as the initial meat of the next experiment can relieve the first two sets
of pitfalls. Hirschmann and Dasgupta (2007) propose a new modified iterative sandwich
experiment technique that approaches equilibrium in a small number of experiments and
successfully produces the composition of finite near-solidus melt fractions. This
technique may be an important next step in determining the equilibrium melt
compositions in <1 100 C Till et al. (2011) experiments. Using a diamond wire saw to
thinly slice the experimental capsules could create additional surfaces to expose
additional analyzable quench pockets in some experiments, as well. The use of LA-ICPMS to analyze the major element composition of quench by Humayun et al. (2010)
suggests this technique may also be useful in experiments with pooled melt to further
refine estimates of the equilibrium melt composition of H20-saturated peridotite at
pressures >3 GPa.
5. Concluding Remarks
EMP analyses of the quench products in the H20-saturated peridotite melting
experiments of Till et al. (2011) are suitable to document the presence of both a low
alkali tholeiitic melt and a high-silica vapor-phase quench. The composition of the
equilibrium melts can be determined if the Mg-Fe partition coefficients and co-variations
in oxide abundances are used to filter the quench compositions. The melting reactions
and melt production rates compare favorably with previous studies of hydrous and
anhydrous peridotite melting behavior and are compatible with the solidus temperature of
-800'C at 3.2 GPa interpreted by Till et al. (2011). The low extent hydrous melts at 3.24 GPa are silica-undersaturated and do not resemble the SiO 2-rich compositions of melts
documented at lower pressures. Several additional techniques are recommended to refine
the estimates of the equilibrium melt compositions. These preliminary results support
previous interpretations that low extent melts of H20-saturated peridotite at 3-4 GPa are
the forefathers of magmas erupted at volcanic arcs.
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White, B.S., and Wyllie, P.J., 1992, Solidus reactions in synthetic lherzolite-H20-CO2 from 2030 kbar, with applications to melting and metasomatism, in Cox, K.G., and Baker, P.E., eds.,
Essays on Magmas and Other Earth Fluids (a volume in Appreciation of Professor P.G.
Harris), Volume 50, Journal of Volcanology and Geothermal Research, p. 117-130.
Yogodzinski, G.M., Kay, R.W., Volynets, O.N., Koloskov, A.V., and Kay, S.M., 1995,
Magnesian andesite in the western Aleutian Komandorksy region: Implications for slab
melting and processes in the mantle wedge: Geological Society of America Bulletin, v. 107,
p. 505-519.
Figure and Table Captions
Figure 1. Electron back-scatter images of experiments with analyzed quench.
Experiment D214 at 1 100 C, 3.2 GPa exhibits a pool of the segregated melt at the hotter
end of the experimental capsule, where both the quench wisps and high-silica quench
spheres are visible. Experiment D230 at 4 GPa, exhibits similar features but with quench
carbonate in place of the high-silica quench.
Figure 2. SiO 2 vs. MgO for quench analyses with EMP totals >65 wt.%. All
compositions were normalized as anhydrous prior to plotting and averaged for to produce
the average Filter 1 compositions in Table 1. Only the analyses with 0.20<KDMg-FeOliv-Liq
<0.42 (red squares) are averaged to produce the Filter 2 quench compositions. In Filter 3
(green circles), the quench analyses that we interpret sampled multiple phases (i.e.,
quench + equilibrium mineral or high-silica quench) are removed.
Figure 3. KDMg-FeOliv-Liq vS. MgO for quench analyses with EMP totals >65 wt.%.
Symbols are the same as in Figure 2.
Figure 4. Average quench compositions for the experiments at 3.2 GPa determined using
the three filters plotted on the oxygen-normalized Oliv-Cpx-Qtz pseudotemary (projected
from garnet) of Tormey et al. (1987).
Figure 5. Proportions of melt and the crystalline phases in the 3.2 GPa experiments
calculated with each of the average quench compositions determined from the three
filtering techniques. The proportions of the equilibrium minerals calculated without a
melt composition for experiments at <1 100 C and 3.2 GPa from Till et al. (2011) are also
shown in each plot.
Figure 6. Phase proportions and melting reactions predicted by our preferred melt
compositions. The proportions of the crystalline phases in all the experiments at
<11 00 C have also been renormalized to include the predicted proportion of melt.
Figure 7. Anhydrous peridotite melts of Walter (1998) at 3 GPa (gray) and low-pressure
hydrous peridotite melts of Mysen et al. (1974) (green), Grove et al. (2006) (blue) and
Till et al. (2011) (orange) plotted on the oxygen-normalized Oliv-Cpx-Qtz pseudoternary
(projected from garnet) of Tormey et al. (1987). The bulk composition for all of these
experiments resembles that from Till et al. (2011) (black).
Figure 8. Isobaric melt production rates (dF/dT)p as determined from peridotite melting
experiments at -
or 3 GPa.
Figure 9. Comparison of volcanic rocks from the Cascades volcanic arc and the average
quench compositions from the experiments of Till et al. (2011). The compositions of
quenched mafic inclusions in andesite lavas from Mt. Shasta are from Krawczynski and
Grove (201l b). Star indicates the composition of a parental arc magma from Mt. Shasta
from Grove et al. (2005). Samples are plotted on the IUGS classification for extrusive
rocks. The range of melt compositions illustrated are corrected for the loss of alkalis
during EMP analyses.
Table 1. Average quench compositions determined from H20-saturated peridotite
melting experiments of Till et al. (2011). EMP quench analyses were filtered using three
different methods to determine their average compositions for each experiment. The
composition of similar extent anhydrous pyrolite melts at 3 GPa from Walter (1998) are
shown for comparison.
D214 (11000C, 3.2 GPa)
D230 (1100*C, 4 GPa) j
BED230qsnch
Figure 1.
0
D194 (1200*C, 3.2 GPa)
D209 (11750C, 3.2 GPa)
50
0
40
0
0
-30
20
0
X remelted
10
quench
0
0
0
20
40
60
80
0
10,
20
8102 (wt. %)
40
60
SiO2 (wt.%)
80
10
60
D200 (11500C, 3.2 GPa)
D207 (11250C, 3.2 GPa)
50
40
0
30
20
10
0
0
20
40
60
SiO2 (wL%)
80
0
101
40
60
80
10
SbO
2 (wt.%)
60
60
D214 (11000C, 3.2 GPa)
* olivine
50
20
D230 (11000C, 4 GPa)
50
40
40
Aorthopyroxene
-30
-30
0
0
20
garnet.
20
clinopyroxene
10
10
high-silica quench
0
0
20
Figure 2.
40
60
SIO2 (wt.%)
80
100
0
20
40
60
8102 (wt.%)
80
100
LEGEND
Filter 1:all analyses with totals >65 wt.%
Filter 2: analyses from Filter 1 with 0.20<KoFe-Mgoi-Uq<0.42
Filter 3: analyses with >65 wt.% totals and reasonable Ko's, that also
do not fall on mixing lines with the equilibrium minerals or the highsilica quench
1
1
D194 (1200*C, 3.2 GPa)
0.9
D209 (11750C, 3.2 GPa)
0.9
0.8
0.8
0.7
0.7
0.6
20.6
5
0.5
010.5
L 0.4
0.4
gwo
0.3
a
0.3
0.2
0.2
0.1
0.1
0
15
10
5
0
20
25
30
,,,,,-.
0
35
0
MgO (wt.%)
5
10
25
20
15
MgO (wt.%)
30
35
1.0
1
D200 (11500C, 3.2 GPa)
0.9
D207 (11250C, 3.2 GPa)
0.9
0.8
0.8
0.7
0.7
0.6
40.6
0.5
010.5
LL0.4
U. 0.4
U.
0.3
0.2
0.1
0.0
0
5
20
15
MgO (wt.%)
10
25
30
5
35
10
15
20
25
30
35
MgO (wt.%)
D214 (1100*C, 3.2 GPa)
D230 (1100*C, 4 GPa)
0.9
0.8
0.7
0.6
0
010.5
0.6
**
05
a 0.5
LL0.4
U
L 0.4
*
4
0.3
0.2
0.1
0
5
Figure 3.
10
20
15
MgO (wt.%)
25
30
35
0
35
15
20
25
30
MgO (wt%)
LEGEND
Filter 1: all analyses with totals >65 wt.%
Filter 2: analyses from Filter 1 with 0.20<KoFo-MgIiv-Uq<0.42
Filter 3: analyses with >65 wt.% totals and reasonable Ko's, that also
do not fall on mixing lines with the equilibrium minerals or the highsilica quench
0
5
10
Figure 4.
1
0.5
Ft 0.4'
n32 GPamet
0
0. 3
03.2 GPaolivine
3.2 GPacpx
II.
3.2GPaopx
0
0.1
Filter 1.
0
-
-
900
850
800
-
--
-
-
1050
1000
950
Temperature (*C)
1100
1150
1200
b.
0.5
04
0. 3
U.
02
0.1
Filter 2.
800
900
850
1050
1000
950
Temperature (*C)
1100
1150
1200
0.6
ooooo
0.5
0.4
0
0.3
U.
&0.2
Filter 3
0
800
850
Figure 5.
-
900
1050
1000
950
Temperature (*C)
1100
1150
1200
3.2 GPagamet
<11000C
>11000C
Oliv + Cpx + Garnet -- > Opx +Melt
Oliv + Opx -- > Melt
Cpx
1100*CM
CPa
10c
I75C
remeltedque
remeltedque S1200'
bulkcomp
1200 ME
0 bulkcamp
Oliv
I
Oliv
Qtz
utz
Opx
0.6
0.5
t 0.4
llm3.2
0.3
GPa melt
4"-13.2 GPa olivine
41 3.2 GPa cpx
dwWW3.2
GPa opx
0.2
IN3.2 GPa garnet
0.1
low
0
800
850
900
950
1000
1050
Temperature (*C)
Figure 6.
1100
1150
1200
Q
Cpx
Oliv
to
3d I %# 'L
Opx
L& 0.
0.,
@1.2 GPa
1020C
Figure 7.
X Qtz
1700
Legend
Anhydrous Experiments
" Falloon et al., 1999 (1 Gpa)
1600
1500,
" Baker & Stolper, 1994 (1 GPa)
1400.
" Falloon & Danyushevsky,
2000 (1.5 GPa)
o Walter, 1998 (3 GPa)
1300-
H 20-Saturated Experiments
* Grove et al., 2006 (1.2 GPa)
1200.
* This study (3.2 GPa)
1100.
H20-Undersaturated Experiments
0000.0
* Gaetani & Grove, 1998 (1.2 GPa)
1000U
900800
0
10
20
30
40
Melt Fraction (wt%)
Figure 8.
50
60
12
Thriphonolite
Parental
from
R
Mt.
arc
lava composition
(85-44)
Shasta
10
Phonotephrite
Trachyandesite
Fod0
Rhyolke
Basaltic
Trachy-
+
Tephrite
andesite
0Basanit*c
salt
4
Picro-
Dacite
esite
c
Fgeste\
2basa
hi hi siliCa
nt
meltq
36
40
44
45
152
56
60
SiO2 (Wt.%)
Figure 9.
64
68
72
7
0
84
Table 1. Average composition of quench in experiments at 1100-1200"C and 3.2-4 Gpa from Till et al. (2011).
Cr203
FeO
MnO
MgO
CaO
SiO2
TiO2
A1203
0.40
7.56
0.10
37.82
0.18
4.06
46.20
Hart & Zindler (1986) PM
Filter 1 Results: Quench with totals >65 wt.%
16.99
48.97
0.98
3.2 Gpa, 1100 D214
6.42
18.09
7.40
14.02
49.77
0.83
3.2 Gpa, 1125 D207
13.91
7.88
11.50
17.01
47.95
0.73
3.2 Gpa, 1150 D200
7.62
51.97
0.48
9.80
3.2 Gpa, 1175 D209
21.93
19.99
8.35
48.65
0.60
9.75
3.2 Gpa, 1200 D194
48.27
0.43
11.09
3.2 Gpa, 1200 D194 remelted
10.38
20.65
21.65
53.26
0.49
12.66
7.37
4 Gpa, 1100*C D230
Filter 2 Results: Quench filteredby Kd only
1.04
18.31
48.52
11.40
5.95
3.2 Gpa, 1100 D214
0.83
13.42
47.59
7.93
3.2 Gpa, 1125 D207
15.07
0.71
10.91
7.82
47.92
3.2 Gpa, 1150 D200
17.23
0.48
10.23
49.82
3.2 Gpa, 1175 D209
8.52
20.60
48.19
0.64
9.95
8.47
3.2 Gpa, 1200 D194
18.39
10.71
50.30
1.47
4 Gpa, 1100'C D230
7.73
14.34
Filter 3 Results: Quench filtered by Kd and mixing lines
3.2 Gpa, 1100 D214
42.78
1.29
17.91
6.05
10.42
3.2 Gpa, 1125 D207
47.02
0.82
13.02
7.97
15.27
3.2 Gpa, 1150 D200
47.10
0.83
12.25
16.32
7.21
3.2 Gpa, 1175 D209
48.64
0.54
11.76
8.58
18.71
3.2 Gpa, 1200 D194
48.11
0.67
9.45
8.43
15.52
4 Gpa, 1100*C D230
7.03
47.19
2.25
9.56
11.67
High Silica Quench
3.2 Gpa, 1060 D206
82.28
0.19
11.46
0.09
0.29
0.05
0.50
3.2 Gpa, 1100 D214
74.75
0.38
17.85
0.05
0.37
0.06
0.30
AnydrousPeridotite Melts of Walter (1998)
3 Gpa, 1540*C 24% melt
46.91
0.64
12.46
8.86
0.17
18.22
3 Gpa, 1580*C 37% melt
48.98
0.48
11.06
9.45
0.18
19.71
Na20
K20
Total
Mg#
3.22
0.33
0.03
7.61
13.31
14.35
7.53
11.97
8.37
4.12
0.45
0.28
0.21
0.09
0.12
0.30
0.23
0.79
0.76
0.78
0.79
0.80
0.78
0.84
13.66
14.44
14.88
9.76
13.70
15.05
0.46
0.29
0.18
0.09
0.13
0.10
0.77
0.77
0.80
0.81
0.79
0.76
20.69
15.18
15.75
11.11
17.17
22.00
0.34
0.25
0.20
0.17
0.14
0.11
0.76
0.77
0.80
0.79
0.76
0.75
3.82
5.72
0.51
0.20
0.47
0.27
10.86
8.78
0.82
0.77
0.34
0.23
0.90
0.74
0.58
99.7
110.2
Kd Oliv-Liq Fe-Mg
0.28
0.39
0.14
Chapter 3:
Primary basaltic magmas from variably metasomatized
mantle: The effects of pressure, temperature, and composition
on melting plagioclase and spinel lherzolite
100
Abstract
Here we develop a new model that simulates melting of variably metasomatized mantle
peridotite and explore the effects of variation in mantle pressure, temperature and
composition on melt composition. This new model combines the approaches of Kinzler
and Grove (1 992a,b) and Kinzler (1997) to predict the temperature and major element
composition of a broad spectrum of primary basalt types produced under anhydrous
conditions at upper mantle pressures. The model can also be used to calculate the
temperature and pressure at which primary magmas were produced in the mantle, as well
as to model both near-fractional adiabatic decompression and batch melting. Our
experimental compilation definitively locates the pressure interval of the plagioclase to
spinel transition on the solidus and shows that it is narrow (~0. 1 GPa) for melting of
natural peridotite compositions. A comparison of our melting models to other mantle
thermometers reveals that the choice of fractionated phase assemblage(s) used to correct
primitive liquids back to primary magmas has a significant effect on the calculated source
temperature (>200'C). Our model is used to examine the petrogenesis of a suite of
Holocene basaltic lavas from Diamond Crater in Oregon's High Lava Plains (HLP) and
demonstrate that the primary magmas inferred for Diamond Crater are ~6 wt.% batch
melts of a spinel lherzolite mantle. After correction for fractional crystallization, our
model predicts these basaltic liquids were last in equilibrium with the mantle at 1.2-1.6
GPa and 1322-1340'C.
1.
Introduction
Primary basaltic magmas are the principal instrument for understanding the depth,
temperature and style of melt generation in the Earth's upper mantle. Beneath mid-ocean
ridges (MOR's), near-fractional adiabatic decompression melting is the dominant process
of melt generation (Klein and Langmuir, 1987; 1989; Kinzler and Grove, 1992b, Niu and
Batiza, 1993). In contrast, petrologic and geophysical constraints show that melting
beneath the thick crust of continental margins and interiors occurs under near-batch
melting conditions (Bartels et al., 1991; Holbig and Grove, 2008), often close to the
crust-mantle boundary. An additional important but little explored factor is the effect of
variations in the composition of mantle lherzolite. A number of major element mantle
melting models have brought us closer to a rigorous quantitative description of melting
focused on the formation of mid ocean ridge basalts (MORB) (e.g., Klein and Langmuir,
1987; McKenzie and Bickle, 1988; Niu and Batiza, 1991; Kinzler and Grove, 1992a, b;
Langmuir et al., 1992; Kinzler, 1997). However, a significant portion of primitive basalts
not erupted at MOR's, including many that form in continental back arc settings and at
ocean islands, exhibit a greater range of K2 0 (up to 2.0 wt.%), and P2 0 5 (up to 0.50
wt.%) relative to MORB, which is thought to result from melting variably metasomatized
peridotite. Here we seek to build on this prior work and examine mantle melting
behavior over a wide range of upper mantle compositions and specifically constrain the
effects of K 2 0 on mantle melting equilibria. Small extent melts of the upper mantle are
saturated with the complete upper mantle phase assemblage of olivine + orthopyroxene +
clinopyroxene + plagioclase and/or spinel at genesis, and our models are restricted by
incorporating only these multiply-saturated mantle melts in their calibration. We will
102
compare our estimates of the temperature and pressure of mantle melting for MORB and
ocean island basalts (OIB) to those produced by other popular thermometers and
barometers. The compilation of experimental melt compositions in equilibrium with an
upper mantle phase assemblage also facilitates observations of the location of the
plagioclase to spinel lherzolite boundary and we explore how the stability of plagioclase
+ spinel lherzolite varies with pressure, Mg# (Mg/(Mg+Fe T )), Na 20, and K20. Our
models for melting the upper mantle can also be used to identify at what pressure and
temperature a primitive basaltic magma was last in equilibrium with the mantle and the
extent to which it crystallized on the way to the surface. We provide a worked example
of these calculations by examining a suite of primitive basaltic lavas erupted in the
Diamond Crater area of Oregon's High Lava Plains (HLP).
11.
Experimental Data
a. ExperimentalData From the Literature
A query was conducted in the Library of Experimental Phase Relations (LEPR)
database (Hirschmann et al., 2008) for experiments containing the phases glass, olivine,
orthopyroxene, clinopyroxene, and plagioclase or spinel. Starting materials for the
experiments from the literature range from synthetic analogs of fertile or depleted mantle
to natural MORB and high alumina olivine tholeiite (HAOT) lava samples (Table 1).
The experiments returned by our query were divided into two categories: one for the
plagioclase-bearing experiments, the other for spinel-bearing experiments (Table 1).
Experiments that contained both plagioclase and spinel were included in both categories
for model development. In addition to the experiments downloaded from LEPR, we
103
added to our datasets spinel and plagioclase lherzolite melting experiments from an
electronic supplement to Kinzler (1997), 0.1 MPa experiments from Grove and Juster
(1989), CMAS experiments from Presnall et al. (1979), and CMASN experiments from
Walter and Presnall (1994).
Several different criteria were used to evaluate the data. First, experiments saturated
with garnet were removed from the dataset, these were experiment L92 from Kinzler
(1997), experiments DPI-48 and DPI-44 from Draper and Johnston (1992), and
experiments 3402.11, 3403.11, 3503.11 from Walter and Presnall (1994). Second,
experiments conducted in the presence of water were also removed from the datasets and
include experiments B277g, B304, B329, B330, B333, B342, B348, B359 from Gaetani
and Grove (1998), 85-41-7b from Grove et al. (2003) and experiments 71,72 and 76 from
Feig et al. (2006). Third, single experiments that were clear outliers from the surface
defined by the bulk of the data in pressure, temperature and composition space were
removed. These experiments include a 2 kilobar experiment from Takagi et al. (2005)
and four 10 kilobar experiments from Meen (1990).
b. New Experiments
New experiments were also undertaken as part of this study (Tables 2, 3, 4) to
determine the phase relations and compositions for a high K2 0 basaltic liquid in
equilibrium with a both a plagioclase and spinel lherzolite mantle phase assemblage. To
this aim three sets of experiments were conducted. The first set of experiments utilized a
high K20 basaltic lava composition from Jordan Valley Volcanic Field in eastern Oregon
(JC-30B), which between 8-14 kbars and 1250-1310 C saturated with olivine,
104
clinopyroxene, and plagioclase but failed to crystallize orthopyroxene. A second set of
experiments, conducted with a derivative of the original starting material (JC30B+opx),
were multiply-saturated with olivine, orthopyroxene, clinopyroxene and plagioclase
between 8-10 kbar. The third set of experiments was conducted on a composition
calculated to be in equilibrium with a spinel lherzolite mantle composition at 17 kbar
using our model equations (SpLM). This starting composition was saturated with olivine,
clinopyroxene, and spinel at 17 kbar and 1313'C but again failed to crystallize
orthopyroxene between 1310-1353'C. Here we report the results of the second and third
set of experiments but use only the second set of experiments in our model that yield
insights relevant for our understanding melts saturated with a complete upper mantle
phase assemblage.
c. Startingmaterial
A primitive basaltic composition with 48.43 wt% SiO 2 , 0.96 wt% K20 and a Mg#
of 0.64 from Coffee Pot Crater in the Jordan Valley Volcanic Field in eastern Oregon was
chosen as the first starting composition for our experiments (JC-30B; Table 2). A
synthetic oxide mixture of this composition was prepared and homogenized in an
automatic agate mortar and pestle for three hours. A pellet of the sample was then
prepared using polyvinyl alcohol as a binder and attached to a Pt loop for use in a
DelTech 1 atm gas mixing furnace. The sample was held at 1000'C at the quartzmagnetite-fayalite buffer in the furnace to ensure the iron in the starting material was
dominantly FeO, following the methods of Grove (1981) and Grove and Juster (1989).
The sample was then reground into a powder with the automatic agate mortar and pestle.
105
A second starting composition was subsequently fabricated using the original starting
material and Kragero orthopyroxene (JC30B+opx; Table 2). 15 wt. % hand-picked,
cleaned, and crushed orthopyroxene was added to the original starting material and
homogenized for three hours in the automatic agate mortar and pestle. A third starting
material with 47.09 wt% SiO 2, 0.62 wt% K20 and a Mg# of 0.73, which is an estimate of
a liquid in equilibrium with a spinel lherzolite at 1.7 GPa calculated with our model
(SpLM; Table 2) was prepared following the methods described above for JC-30B. Prior
to each experiment, approximately 4 mg of the starting material was packed into a
graphite capsule and held at ~125'C for at least 24 hours to assure the experiments were
conducted under volatile-free conditions.
d. Experimentalmethods
Experiments were performed in a 0.5" end-loaded solid-medium piston cylinder
apparatus (Boyd and England, 1960) in the MIT Experimental Petrology Laboratory
using the setup detailed by Medard et al. (2008). All experiments were conducted in
graphite capsules that totaled -0.15" in length. The capsules were placed in a nonporous
A120 3 sleeve, positioned in the center of a graphite furnace using MgO spacers, and
encased in a sintered BaCO 3 pressure medium. The temperature was monitored and
controlled using W97Re 3-W75Re 25 thermocouples positioned ~1.5 mm above capsule.
The experimental temperatures reported in Table 2 are corrected for the 18 ± 6'C
temperature difference between the thermocouple temperature (colder) and the average
capsule temperature (hotter) (Medard et al., 2008). The pressure was calibrated using the
breakdown reaction of Ca-tschermak pyroxene to anorthite + gehlenite + corundum and
106
the spinel to garnet transition in CMAS (Hays, 1966; Longhi et al., 2005). Experimental
conditions are reported in Table 3.
e. Analytical methods
Major element concentrations of the experimental products were analyzed with
the JEOL 733 or 8200 microprobes at Massachusetts Institute of Technology. All
analyses of the individual phases presented in Table 3 were conducted with a 15kV
accelerating voltage, a 10 nA beam current and a beam diameter of 1 pm for solid phases
and 10 pm for silicate glasses. On-line data reduction used the CITZAF correction
package (Armstrong, 1995) and analytical precisions were estimated from replicate
analyses of a basalt glass working standard reported in Gaetani and Grove (1998).
g. Composition of Experimentally ProducedMelts Saturated With a Mantle Residue
Assemblage
The composition of the melts in equilibrium with spinel lherzolite from the
previously published and new experiments is illustrated in Figure 1. The lower pressure
experiments (i.e., 1-1.5 GPa) cover a wider range of liquid composition space relative to
the higher pressure experiments because of the greater diversity of starting compositions
used for the lower pressure experiments. At a given pressure, melts coexist with an upper
mantle assemblage over a temperature range of ~50-150'C, which narrows toward higher
pressures. The dataset is well suited to explore the effects of oxides such as K20 and
TiO 2 on mantle melting as highlighted by the range of these oxides in the experimental
liquids relative to CMAS liquids, which plot at an Mg# of unity. The experimental
107
liquids in these compositionally complex peridotitic systems exhibit an overall
covariation in NaK# ((Na 2 O+K2 0)/(Na 2O+K 2 0+CaO)), wt.% TiO 2, and wt.% A12 0 3 with
Mg#.
The composition of the minerals and melts from our experiments are given in
Table 4. Three liquids in equilibrium with a plagioclase lherzolite assemblage have 1-1.3
wt.% K 2 0 and Mg#'s of 0.48-0.55. The liquids are very similar in their oxide content to
other plagioclase lherzolite liquids compiled from the literature (Figure 1).
III.
Parameterization of Plagioclase and Spinel Lherzolite Melting Equilibria
Melting of plagioclase and spinel lherzolite is a 5-phase equilibrium involving
liquid + olivine + orthopyroxene + clinopyroxene + plagioclase or spinel. The
thermodynamic variance of the melting equilibrium can be obtained from the Gibbs
phase rule, F= C + 2 - #, where F is the number of degrees of freedom or variance, C is
the number of components and # is the number of phases. In the CMAS system, these
melting reactions have one degree of freedom and are univariant (F =1 = c + 2 - *).
Therefore, specifying the pressure of melting fixes the temperature and composition of all
the phases. If an additional chemical component is added (e.g., Na 2 0: Walter and
Presnall, 1994), the thermodynamic variance increases by one and the Gibbs method
(Spear, 1993) can be applied to constrain the compositions of all coexisting phases. We
choose to extend the univariant CMAS composition space into the natural system in our
model by exploring the effects of additional compositional degrees of freedom on the
melting equilibria, following Longhi (1991), Grove and Juster (1989), Kinzler and Grove
(1992a) and Kinzler (1997). The number of chemical components was chosen to be nine
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(this includes the dominant major and minor oxides required to calculate mineral
components: SiO 2 , TiO 2, A12 0 3 , FeO, MgO, CaO, Na 20, K20, P2 0 5) and there are five
mantle phases involved (olivine, orthopyroxene, clinopyroxene, plagioclase/spinel, melt),
therefore the thermodynamic variance of our system is six. Thus five compositional
variables in addition to pressure should ideally be fixed to constrain the variability of
upper mantle melts in P-T-X space. In previous studies, wt.% TiO 2, K20, Na 20 and Mg#
have been identified as important compositional variables that systematically influence
the composition of multiply-saturated liquids (e.g., Kushiro, 1975; Grove and Juster,
1989; Walter and Presnall, 1994). We chose to express the composition of the spinel and
plagioclase lherzolite melts using oxygen-based mineral components, because these vary
systematically from their CMAS univariant values in response to changes in
compositionally important exchange vectors like Mg-Fe and Ca-Na. We used the same
component choice as Kinzler and Grove (1992a): Olivine (Oliv), Clinopyroxene (Cpx),
Plagioclase (Plag) and Quartz (Qtz). After extensive testing, we settled on an empirical
model with only four compositional variables in addition to pressure. The five
independent variables in our model are 1-Mg#, NaK#, wt.% TiO 2, wt.% K20 and
pressure (in GPa) (Table 5). Principal component analysis of these variables indicates
that >97% of the chemical variance in the dataset is accounted for in two principal
components in which K 20 and TiO 2 are the dominant coefficients.
The results of linear least squares regressions of the experimental dataset using
the five independent variables are illustrated in Figure 2. Here we show the difference
between the observed value of the mineral components Oliv, Cpx, Plag and Qtz for
experimental liquids in equilibrium with spinel lherzolite and the same value calculated
109
using our equations and the pressures and temperatures of the experiments. Figure 2a-d
illustrate the mineral component values returned by the regressions of the full dataset,
while Figure 2e-h illustrate the mineral component values returned by regressions of only
those experiments analyzed on the microprobe at MIT and experiments done in the
simple systems CMAS and CMASN. There is significantly more scatter in the mineral
component values calculated from the full dataset equations relative to the MIT +
CMAS(N) equations. One explanation for the scatter in Figure 2a-d is likely systematic
differences in the pressure calibration and the microprobe standardization used to
produce the experimental liquid compositions at different labs. For example, the
synthetic peridotite melting experiments of Kinzler (1 992a) and Kinzler (1997) were both
analyzed using the same microprobe and standards at MIT but conducted at different
piston cylinder labs. When the pressure of these experiments are calculated using the
spinel lherzolite olivine equation from the MIT + CMAS(N) regression, pressure is
systematically underestimated for experiments run at the MIT experimental lab and
overestimated for the experiments conducted at Lamont Doherty experimental lab (Figure
3a), pointing toward discrepancies in the pressure calibration between these two labs.
Here we report only the regression of experiments in the CMAS(N) system plus those
analyzed at MIT (Table 5; Figure 2e-h), which more closely reproduces the known
experimental glass composition in each case. The regressions of the full experimental
dataset are presented for comparison in Appendix 1.
The expressions in Table 5 are qualitatively similar to the plagioclase and spinel
lherzolite melting equilibria expressions of Kinzler and Grove (1992a). All the
coefficients have the same sign and relative effect on temperature and the composition of
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the liquid expressed in mineral components with the exception of the I -Mg# coefficient
in the Cpx expression, which is 0 in the Kinzler and Grove (1992a) expression, and the
TiO 2 coefficient in the Plag expression, which is the most poorly resolved coefficient in
all expressions. Our calculated 1 atm solidus temperature for the metastable spinel
lherzolite assemblage is -40'C higher than that calculated by Kinzler and Grove (1992)
for the regression of the MIT + CMAS(N) experiments (Table 5) and ~30'C higher for
the full dataset regression (Appendix 1), while our 1 atm plagioclase lherzolite solidus
temperatures are very similar for both regressions (122 0 'C, 1216'C) and ~20'C lower
than Kinzler and Grove's (1992a). In our expressions, 1-Mg#, NaK# and K20 lower the
solidus temperature for both plagioclase and spinel lherzolite saturated liquids, and all
other components serve to increase the solidus temperature. Pressure exerts a positive
control on temperature and the Oliv and Plag mineral components for a liquid in
equilibrium plagioclase lherzolite, as well as a positive control on temperature and the
Oliv and Cpx mineral components for a liquid in equilibrium with a spinel lherzolite.
Principal component analysis of the mineral component variance reveals that
-99% of the variance resides in two principal components, one in which Plag has the
dominant coefficient, the other in which Oliv has the dominant coefficient. Of the four
potential barometers, both the Plag and the Qtz equations are the most sensitive to small
variations in analytical error, which manifests itself in the NaK# coefficient. Because of
the coordination of sodium to three silica ions in feldspar, small changes in the sodium
content of the melt have large effects on the calculated Plag and Qtz components. This is
evident in the larger average errors returned by the Plag and Qtz equations relative to the
Oliv equation. In spite of the lower r 2 for the spinel lherzolite Oliv equation (r2 = 0.63),
this equation recovers pressure with the lowest average error (0.25 GPa) and thus it is our
preferred equation for calculating the pressure of a spinel lherzolite primitive melt. The
Cpx component exhibits the poorest fit to the experimental liquid dataset, and has the
least variance, suggesting the majority of variability may stem from interlab calibration
issues. Therefore a single value for the Cpx component may be an adequate
approximation for the range of bulk compositions and pressures considered here.
The plagioclase and spinel lherzolite thermometers (Figure 3b, Table 5) have r 2
values of 0.97 and 0.94, respectively, and reproduce the experimental temperatures with
average errors that are equivalent to the temperature uncertainty of the experiments
(-10 C). The plagioclase thermometer is calibrated over the temperature range of-
100-
1300'C and overlaps for -50'C with the calibrated 1250-1550'C range of the spinel
lherzolite thermometer. The overlap in these calibrations reflects the range of
temperatures where both spinel and plagioclase are stable in the experimental dataset.
IV. Plagioclase Lherzolite to Spinel Lherzolite Boundary
The compilation of experimental plagioclase and spinel lherzolite liquids allows
us to revisit the location and primary controls on the plagioclase to spinel transition (PST)
in the upper mantle. Presnall et al. (1979) locate the invariant PST in the CMAS system
at 0.9 GPa and 1300'C. Additional experiments by Walter and Presnall (1994) in CMAS
have both spinel and plagioclase present at 0.93 GPa, with only plagioclase present at 0.7
GPa and only spinel at 1.1 GPa. Because there is a negligible temperature-dependence of
the PST (Kushiro and Yoder, 1966), the experiments in the CMASN system can be used
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to define the location of the univariant PST in P-NaK# space. The equation of the
resulting line is,
P = 1.56 (NaK#) + 0.999
(1)
where the pressure of the PST is in GPa. This result reinforces earlier observations that
the Na 20/CaO content of the liquid, a proxy for the degree of depletion of the source, is
positively correlated with the pressure of the PST (Walter and Presnall, 1994; Falloon et
al., 2008).
The PST expands from a line in two dimensions into a field where both
plagioclase and spinel are stable in three dimensions. This can be visualized in Figure 4
where the plagioclase~plagioclase+spinel and the plagioclase+spinel4spinel
boundaries are depicted as two bent screens. We also can approximate the plagioclase +
spinel*spinel boundary as a plane for practical use, where the pressure of this boundary
in GPa is given by the expression,
P = (0.615 NaK# + 0.3255)/0.35
(2)
Although the experimental liquid compositions suggest an apparent positive correlation
between pressure and Mg#, the third dimension in our illustration, it is actually an artifact
of the Na 20/CaO control on the location of the PST, as evidenced by the covariance in
NaK# and Mg# in our dataset of experimental liquids (Figure 1). Therefore Mg# does
not appear in this planar approximation of the PST. Experimental liquids in equilibrium
with both spinel and plagioclase with the highest NaK#'s and lowest Mg#'s in the dataset
suggest the maximum width of the PST field in this composition space is ~0.3 GPa.
Therefore low extent mantle melts are in equilibrium with both plagioclase and spinel
over a relatively narrow pressure interval. Typical mantle melts with an Mg# equal to
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0.75 plot close to the Mg-axis, therefore plagioclase + spinel are stable in these liquids
over a pressure range much smaller (~0. 1 GPa: 1.1-1.2 GPa) than 0.3 GPa. Metasomatic
enrichment of mantle alkali contents will increase the NaK# of the mantle and slightly
increase the pressure interval over which plagioclase and spinel are stable.
V.
Implementation of the Plagioclase and Spinel Lherzolite Models
Our mantle melting model can be used in two ways. The first is to estimate the
depth and temperature of origin of a fractionation-corrected magma composition. The
second is to model mantle melting processes given a mantle bulk composition and
melting model. The second step can provide useful information on the compositional
characteristics of the mantle residue and the processes that can lead to the production of a
melt with a given set of characteristics. Ideally, both of these tasks should be carried out
when one is trying to understand the origin of mantle melts.
a. FractionalCrystallization Corrections
The first step in estimating the composition of melts of plagioclase or spinel
lherzolite is to correct for the effects of fractional crystallization. This step requires
knowledge of the conditions under which the melt was modified by fractionation, which
in turn requires knowledge of the phase equilibria that control fractional crystallization
under those conditions. The phase assemblage chosen for the fractional crystallization
correction can have a significant effect on the corrected liquid composition and therefore
the estimates of source pressure and temperature, as evidenced by the comparison of the
thermometers and barometers for MORB compositions (see Section VI below). For our
fractionation corrections, we used the methods of Grove et al. (1992) (see their Appendix
114
2). Yang, Kinzler and Grove (1996) updated this methods and we recommend using their
approach if clinopyroxene is a fractionating phase. To chose the appropriate
fractionating phase assemblage, the bulk sample compositions are plotted in the oxygennormalized Olivine-Plagioclase-Clinopyroxene-Quartz pseudo-quaternary projection
scheme of Tormey et al. (1987). The location of the sample is compared to the location
of the multiple saturation points (MSP) for that sample (i.e., where a liquid of that
composition would be multiply-saturated with a spinel or plagioclase lherzolite mantle
phase assemblage) with varying pressure, which can be calculated using our model
equations (Table 5). This allows the fractionation path of the sample to be visualized in
2D in the Plag-Cpx-Oliv pseudo-ternary projection (Figure 5a). The trace element
composition of a sample, especially the sign of the Eu-anomaly, is a good indication of
whether fractional crystallization has shifted a liquid away from the plagioclase or spinel
lherzolite MSP. Figure 5a illustrates that an olivine-only fractionation correction will
drive a liquid toward the spinel MSP for a much higher pressure than olivine +
plagioclase fractionation. This is exemplified by our pressure estimates for the olivineonly and olivine + plagioclase corrected MORB compositions (Figure 6). Therefore to
use our model to calculate the pressure and temperature a liquid was last saturated with a
mantle lherzolite, an independent constraint on the type of fractionation experienced by
the liquid is required. For example, the presence of olivine + plagioclase or olivine +
plagioclase + clinopyroxene phenocrysts in a basalt sample are a good indication of the
fractionating phase assemblage, as are the chemical composition of these minerals. This
type of data is not available for the MORB glass compositions used in Section VI from
the PetDB database, therefore we plot these compositions on a series of pseudo-ternary
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projections. All the MORB glass compositions plot on the Oliv-Plag cotectic on the
Plag-Cpx-Oliv ternary (Figure 5a) within the analytical uncertainty, which suggests they
all experienced olivine + plagioclase fractionation, with the exception of the EPR and
Azores compositions which may have experienced olivine + plagioclase + clinopyroxene
fractionation. Correcting for olivine + plagioclase fractionation moves the liquid
composition in a vertical direction along the Oliv-Plag cotectic toward the Oliv-Plag axis
as shown in the Figure 5a. The relative abundance of plagioclase and olivine for the
fractionation correction can be estimated by projecting the liquid onto the Oliv-Plag axis
and converting from volume percent to weight percent using the density of the
crystallizing phases. This estimate of the crystallized phase abundances can then be used
to add increments of the crystallized phases to the bulk composition of each sample until
it can be approximated as a liquid in equilibrium with a mantle olivine with a forsterite
content (Fo) of 90.
Here we emphasize that the choice of fractionating mineral assemblage (oliv; oliv
+ plag; oliv + cpx; or oliv + plag + cpx), any changes in the fractionating minerals with
decreasing pressure and temperature, and when the liquid reaches equilibrium with the
mantle, all have a significant effect on the thermometric and barometric calculations for a
primitive basalt. Liquids calculated by Kinzler (1997) are used to illustrate this in Figure
7. As shown in Figure 7, a liquid that has experienced multi-phase fractionation but is
corrected for only olivine fractionation until it is in equilibrium with a Fo90 olivine
significantly overestimates the FeO* and the MgO content of the parental liquid. When
corrected for olivine-only fractionation, the 1.5-0.3 GPa post fractional crystallization
liquid composition from Kinzler (1997) yields a temperature and pressure estimate from
116
our regression that is 145'C and 1.1 GPa greater than the actual parental liquid and the
3.0-0.3 GPa liquids yield temperature and pressure estimates that are up to 225 0 C and
1.75 GPa more than the parental liquid. This figure also illustrates how the choice of
olivine Fo-content that represents equilibrium with the mantle can have a significant
effect on the FeO and MgO of the fractionation corrected liquid and thus any temperature
and pressure calculations using that composition.
b. CalculatingMelts of Mantle Lherzolites
The forward model of plagioclase and spinel lherzolite melting we use here is an
update of the model presented by Kinzler and Grove (1992a,b) that incorporates the
improvements of Kinzler (1997). In order to use the melting model, the Mg# and
abundances of minor components in each melt increment must be inferred. Estimates of
the minor components are obtained using the non-modal batch melting equation,
Ci=Co/DB+F(1-PB)
(3)
where Ci is the concentration of the minor component (e.g., K2 0) in the melt, Co is the
concentration of that component in the bulk solid, DB is the partition coefficient for the
component between solid and melt, F is the extent of melting, and PB is the bulk solidmelt partition coefficient for the component weighted by the stoichiometric coefficients
in the melting reaction (wt.% units). Partition coefficients have been calculated using the
experimental data that is incorporated in the model. The partition coefficient models are
temperature- and pressure-dependent, following the method of Kinzler (1997), with the
important exception of our modeling being conducted in weight units.
Our model uses the pressure-dependent spinel lherzolite melting reaction
calibrated by Kinzler (1997) and the plagioclase lherzolite melting reaction of Kinzler
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and Grove (1 992b). The Mg# of a melt coexisting with a lherzolite composition is
determined by solving equations for the conservation of mass for FeO and MgO between
lherzolite and a given fraction of melt, following the method of Kinzler and Grove
(1993). These equations require knowledge of the partitioning of Fe and Mg between
melt and all the coexisting phases, and the exchange KD's for Fe and Mg are calculated
from the experimental dataset. Since the value for KD
Fe-Mg
can vary with pressure and
melt composition (Ulmer, 1989; Toplis, 2005), the model sets the KDFe-Mg values for
clinopyroxene, orthopyroxene and spinel to 0.94, 0.97 and 1.5 times the KDFe-Mg value for
olivine-melt.
Melting in the model can be specified to be either batch or polybaric near-fractional
adiabatic decompression melting and the melt production rate for polybaric melting can
be varied depending on preference for the heat of fusion, we recommend a value of~0.01
wt.% per 0.1 GPa of decompression. In the case study presented in Section VII, we use
this forward model to test models of both batch and fractional upper mantle melting to
generate basalts from Oregon's HLP.
VI.
Comparison of Parameterization to Other Thermometers and Barometers
Previous major element models of upper mantle melting focus on the generation
of MORB (Klein and Langmuir, 1987; McKenzie and Bickle, 1988; Niu and Batiza,
1991; Kinzler and Grove, 1992a, b; Langmuir et al., 1992; Kinzler, 1997) and do not
consider melting of variably metasomatized peridotite. Other thermometers, barometers
and melting models aim to investigate the production of a wider range of basalt
compositions (e.g., OIB) but are calibrated on experimental liquids that are in equilibrium
118
with olivine only (e.g., Ford et al., 1983; Putirka, 2007, 2009) or olivine + orthopyroxene
only (Lee et al., 2009). To explore the differences that might arise in estimating melting
conditions for basaltic liquids with a model calibrated exclusively on liquids saturated
with the complete upper mantle phase assemblage of olivine + orthopyroxene +
clinopyroxene + plagioclase and/or spinel, we compare our thermometric and barometric
estimates for five different locations of MORB genesis and three locations of OIB
genesis to those from a variety of popular mantle thermometers and barometers.
MORB compositions were compiled from PetDB ocean rock database using the
latitude and longitude of the selected areas, filtered for fresh glass analyses (MgO > 8
wt.%) and averaged before they were corrected for olivine-only and alternatively olivine
+ plagioclase fractionation until they were in equilibrium with a Fo90 mantle olivine
following the methods described in Section V. Pressure estimates are required for the
thermometers of Ford et al. (1983), Kinzler and Grove (1992a), Langmuir et al. (1992),
Putirka et al. (2007) and the thermometer presented here (Table 5), and were calculated
for the fractional crystallization corrected compositions at each location using the
barometers of Lee et al. (2009) and the spinel lherzolite olivine expression from this
study. In the case of MORBs, which are thought to be polybaric near fractional melts
(McKenzie, 1984; McKenzie and Bickle, 1988; Johnson et al., 1990), the calculated
pressure represents an average pressure for the entire melting column. Our barometric
estimates for the olivine + plagioclase corrected MORB compositions are within 0.3 GPa
of the Lee et al. (2009) pressures for the olivine only corrected compositions and both are
within the range of accepted values for MORB generation (Kinzler and Grove, 1992b)
(Figure 6). Our pressure estimates for the olivine-only corrected compositions are quite
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high relative to average MORB generation pressures, because our Oliv barometer is
sensitive to the increase of 0.02-0.04 in NaK# and 0.02-0.03 in Mg# for the olivine-only
corrected liquid relative to the olivine + plagioclase corrected liquid. The temperature
estimates for each location vary up to 185'C between thermometers when calculated with
olivine-only corrected compositions (Figure 6) or 117 C when calculated with the olivine
+ plagioclase corrected compositions. It is interesting to note that all temperatures for the
MORB localities cluster around 1300'C when the olivine + plagioclase corrected
compositions are used with the exception of the Ford et al. (1983) thermometer (Fig. 6b).
This suggests the variety of methods employed by the different thermometers compared
here are all similarly sensitive to the chemical variability of the liquid produced by
different types of fractionation corrections. It also appears that when corrected for the
accurate type of crystal fractionation the various thermometers converge on similar
source temperature estimates. This comparison of thermometers also reveals that
calculations of mantle temperature for MORB depend heavily on the type of fractionation
correction and the source pressure.
Determining the temperature and pressure for OIB genesis is significantly more
complicated due to the paucity of primitive OIB glasses that can be positively identified
as liquids and the abundance of compositions that have experienced complex fractional
crystallization and other pre-eruptive processes after they segregated from a peridotitic
source. Glasses from Kilauea volcano in Hawaii represent some of the only picritic
glasses from a hotspot locality and work by Wagner and Grove (1998) reveals these
liquids were never in equilibrium with garnet lherzolite assemblage on their liquidus.
Rather, these tholeiites appear to have been in equilibrium with a harzburgitic source at
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lithospheric depths. Wagner and Grove (1998) propose a model where the Kilauea
tholeiites formed by melting of garnet lherzolite in a mantle plume, followed by
equilibration of the liquids with a harzburgitic mantle at shallow depths. In addition,
many primitive OIBs, such as those from Samoa (Workman et al., 2004), Marquesas
(Desonie et al., 1993) and the Ontong Java plateau (Herzberg et al., 2007), exhibit
evidence for olivine and/or clinopyroxene accumulation. Therefore reconstructing the
parental liquid composition of an OIB in order to calculate the source pressure and
temperature is not simply a matter of determining the right phase assemblage for the
fractionation assemblage, as is the case for MORB. A number of techniques have been
developed to circumvent these issues and calculate the parental liquid for a given OIB
compositions (e.g., Putirka et al., 2007; Herzberg and Asimow, 2008). Parental liquid
compositions for selected OLBs calculated by Putirka et al. (2007) plot in a location on a
series of pseudo-ternary ternary projections that are not consistent with liquids in
equilibrium with plagioclase or spinel lherzolite as determined by our model, or garnet
lherzolite as determined by the work of Walter (1998). Rather they appear to have
experienced olivine addition since last in equilibrium with a spinel lherzolite. However
for the sake of comparison, we choose to use the Putirka et al. (2007) calculated
compositions to contrast thermometric and barometric estimates, rather than attempt to
correct a fictitious liquid composition for olivine addition.
Pressures for the calculated parental OIB liquids determined with our olivine
spinel lherzolite barometer and that of Lee et al. (2009) are quite different (Figure 6f).
Both sets of pressure estimates yield temperature estimates that vary by -150'C at a
given locality (Figure 6c & 6d). All the thermometers yield similar temperature estimates
for the Hawaiian and Samoan compositions and temperatures that are ~100-200'C lower
for the Icelandic composition depending on the pressure. The thermometer of Lee et al.
(2009) is not pressure-dependent; therefore the OIB source temperature estimates are
identical at either pressure. The parameterization of Putirka et al. (2007) is the least
pressure sensitive after Lee et al. (2009), and an increase in source pressure of -2 GPa for
the Hawaii and Iceland compositions results in an increase of temperatures of -50'C. At
the other end of the spectrum, the parameterization of Kinzler and Grove (1992a) is the
most pressure-sensitive and an increase in source pressure of -2 GPa results in an
increase of temperatures of -300'C. The pressure-dependence of the remaining
thermometers falls between these two end members, with our model tending toward more
pressure-dependent. It is difficult to draw conclusions about the actual source
temperature and pressure for the three OIB localities from these calculations, which
reinforces the observation that the determination of the parental liquid compositions and
estimates of source pressure have profound effects on the temperature estimates for OIB
genesis.
V1I.
Case Study of Basaltic Lavas from Oregon's HLP
A suite of primitive basaltic lavas from Diamond Crater and surroundings in
Oregon's eastern HLP were selected to demonstrate the utility of our model for
calculating the pressure and temperature where a primitive basaltic liquid was last in
equilibrium with an upper mantle phase assemblage. Diamond Crater (43'05', 118'43')
is a suite of basaltic lava flows emplaced around a central vent between 8,400-6,400 yr.s
ago (Sherrod, written comm., 2009). The lavas are relatively pristine diktytaxitic high
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alumina olivine tholeiites (HAOT) with varying proportions of olivine and plagioclase
phenocrysts visible in hand sample. Major and trace element geochemistry of all samples
presented here were conducted at the Washington State University GeoAnalytical Lab
(Table 6). In addition, electron microprobe analyses of the phenocrysts and groundmass
of samples from Diamond Crater were conducted at the Massachusetts Institute of
Technology following the same analytical procedures outlined for the experiments
presented above.
The petrology and geochemistry of the Diamond Crater samples demonstrate they
are low extent anhydrous melts of spinel lherzolite. The trace element abundances of our
most primitive samples are compared to modeled 1-15% batch melts of a primitive spinel
lherzolite mantle in Figure 8. We modeled batch melting of primitive upper mantle using
the appropriate equations from Shaw (2006), major element abundances from Hart and
Zindler (1986), trace element abundances from Hoffman (1988), partition coefficients
from Halliday et al. (1995) with additional spinel partition coefficients from Canil (2004)
and Kinzler and Grove (1 992a) and mantle modes and melting reactions from Kinzler
(1997). This model suggests samples ch8329 and ch8343 are 2 - 6 wt.% melts of a
primitive spinel lherzolite mantle, samples DC 07-16 and DC 07-12 are 6 -10 wt. % melts
and DC 07-13 and ch8357b are 10-15 wt.% melts. All samples are relatively enriched in
Ba, Sr and La and depleted in Th and Pb relative to equivalent degree model melts. The
positive or negligible Eu-anomaly demonstrates these samples were not a product of
melting plagioclase lherzolite. The bulk rock compositions, as well as the chemical
composition of the plagioclase and olivine phenocrysts in thin sections of the Diamond
Crater samples indicate these liquids did not form in the presence of hydrous fluids.
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Thus these samples are excellent candidates for our regression because they appear to be
anhydrous low extent mantle melts.
For use in our model, the major element compositions of the selected samples
were first corrected for minor fractional crystallization following the methods discussed
in Section V. The presence of either olivine + plagioclase phenocrysts or olivine +
plagioclase + minor clinopyroxene microcrysts in the Diamond Crater samples and the
chemical composition of these minerals imply these samples experienced a period of low
pressure olivine + plagioclase fractionation, in some cases followed by a late episode of
olivine + plagioclase + clinopyroxene fractionation. Alternatively, we tested correcting
the samples for a period of olivine-only fractionation, followed by a period of lower
pressure olivine + plagioclase fractionation. However, correcting for this scenario drives
the liquids toward the olivine apex on a Plag-Cpx-Oliv ternary and beyond the spinel
MSP long before they reach equilibrium with Fo90 olivine. Therefore we conclude, the
most logical fractional crystallization history for these samples is a period of lowpressure olivine + plagioclase fractionation, and in some cases followed by a period of
olivine + plagioclase + clinopyroxene fractionation.
Once corrected for olivine + plagioclase fractionation, the corrected compositions
and their corresponding mineral component values were plugged into our spinel
lherzolite Oliv barometer and the spinel lherzolite thermometer (Table 5) and the results
are presented in Table 7. The samples were generated at temperatures of -1323-1370'C
and pressures of 1.22-1.6 GPa, which corresponds to depths of ~42-54 km based on
density estimates for the crust and mantle from recent seismic refraction profiles near
Diamond Crater (Cox, personal communication, 2011). Recent work on the crustal
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structure of Oregon's HLP using teleseismic P-to-S receiver functions finds the Moho is
located at -30 km in the region of Diamond Crater when the H-K method is applied to
individual stations, or -36 km if the maximum amplitudes from GCCP stacks are used
(Eager et al., 2011). In either case, the primitive basalts erupted in the Diamond Crater
area appear to be formed by shallow decompression melting of spinel lherzolite, with the
shallowest depths of melting almost immediately below the Moho.
We can also apply our forward model of spinel lherzolite melting to model both
batch and near-fractional mantle melting over the range of calculated depths for the
Diamond Crater parental liquids. The trace element modeling indicates Diamond Crater
basalts are relatively enriched in several of the large ion lithophile elements typically
carried by subduction zone fluids and/or melts, such as Ba and Sr, but are depleted in
others, such as Rb and Th, relative to the primitive mantle (Figure 8), suggesting they
may have originated from variably metasomatized mantle. Therefore, we specified the
bulk mantle composition in our models to be either that of the primitive mantle from Hart
and Zindler (1986) or a xenolith of metasomatized depleted mantle from southern
Patagonia (Ntaflos et al., 2007). Models were run with initial pressures of 1.0-1.8 GPa,
an olivine-melt KDe
of 0.3, and either 100% of the melt was removed for batch
melting or 90% in the case of near-fractional melting. Batch melting models were run for
2-10 wt.% melting. For near-fractional melting, we chose the melt production rate
(dF/dP) to be 1% per 0.1 GPa of decompression and track the composition of the melt
produced at each step, as well as the accumulated melt. The modeling results that best
replicate the composition of the fractionation-corrected Diamond Crater basaltic lavas are
presented in Figure 9 and Table 8. Models that use the Patagonia xenolith mantle
125
composition exhibit a much better fit to the Diamond Crater liquids than those that use
the bulk primitive mantle composition (Figure 9). Overall, the 6% batch melts generated
at either 1.3 or 1.6 GPa and the accumulated melt generated during near-fractional
melting between 1.6-1.0 GPa are closest to the Diamond Crater liquids. The 1.3 GPa 6%
batch melt probably exhibits the single best fit, falling within <0.5 wt.% of all the key
oxides, as shown in Figure 9. This is an exciting result, as two completely independent
sets of calculations (i.e., the thermo-barometric calculations and the forward mantle
melting model) both suggest the Diamond Crater basalts were generated during small
degrees of melting of a metasomatized, depleted spinel lherzolite at 1.3-1.6 GPa. This
example also illustrates the power of our spinel lherzolite forward melting model to
constrain the melting style, in addition to the pressure, temperature and composition of
the bulk mantle during melting. Our result that Diamond Crater HAOTs likely formed
under batch melting conditions immediately below the Moho in Oregon's HLP agrees
with an increasing number of studies on HAOT magma generation at continent margins
and interiors (e.g., Bartels et al., 1991; Bacon et al. 1997; Holbig and Grove, 2008). Thus
the plagioclase and spinel lherzolite melting models presented here represent two new
tools to understand the melting of variably metasomatized mantle and the petrogenesis of
HAOT magmas.
126
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132
Figure & Table Captions
Figure 1. Composition of the experimental liquids in equilibrium with a plagioclase
(white circles) or spinel lherzolite (black circles) phase assemblage (olivine + cpx + opx
plag ± spinel) modeled in our regression (see Table 1 for references). Additional
experiments conducted as part of this study are shown for comparison (triangles).
Figure 2. Observed versus calculated mineral component values for spinel lherzolite
experimental liquids empirically fit by our regression. Dashed line represents 1:1
correspondence. Plots a-d show the mineral component values returned by equations
calculated from the full dataset (see Appendix 1), while plots e-h show only the mineral
component values returned from equations calculated with only those experiments done
in the simple systems CMAS and CMASN and those analyzed on the microprobe at MIT.
Figure 3. a) Experimental versus calculated pressure for experiments from Kinzler and
Grove (1992a) and Kinzler (1997) using the spinel lherzolite olivine expression. These
two sets of experiments were conducted in different labs but analyzed on the same
microprobe with the same standards, therefore the systematic differences in pressure
between the two studies likely result from interlab pressure calibration, not analytical,
differences. b) Experimental versus calculated temperature for the plagioclase and spinel
lherzolite experimental liquids determined with our plagioclase lherzolite and spinel
lherzolite thermometers.
133
Figure 4. Illustration of the plagioclase-spinel lherzolite boundary in Pressure-Mg#-(1NaK#) space. The plagioclase-spinel transition (PST) can be conceptualized as two bent
screens (the plag-only to plag + spinel and plag+spinel to spinel-only boundaries) that
define a three dimensional field with a height of -0.3 GPa at high NaK#s and low Mg#s.
This field collapses to a line in both the P-(1-NaK#) and P-Mg# planes. We plot 1-NaK#,
rather than NaK#, to allow the CMASN experiments that constrain the slope of the
univariant PST to be visualized on the plane where Mg#=1. Also shown are experiments
in the CMAS system that constrain the pressure of the PST at an Mg# and 1-NaK# of
one. A typical upper mantle primitive liquid is illustrated by the red X's.
Figure 5. Rational for fractional crystallization corrections depicted on pseudo-ternary
projections. a) Uncorrected average primitive glass compositions from the five MORB
localities plotted with the plagioclase and spinel lherzolite multiple saturation points
calculated for the av. Reykjanes MORB composition. Insert illustrates how different
types of fractional crystallization corrections drive the glass compositions in very
different directions with regards to the spinel multiple saturation points. Gray numbers
indicate pressures for the spinel multiple saturation points. b) Uncorrected MORB
compositions on Oliv-Cpx-Qtz pseudo-ternary. Grey dashed line illustrates the 1 atm
Oliv-Cpx cotectic. The EPR and Azores average compositions lie on the cotectic
indicating the possibility of 1 atm oliv+plag+cpx fractional crystallization for these
liquids. Liquids below the cotectic experienced oliv+plag fractionation. c) Average
MORB glass compositions corrected for olivine and plagioclase crystallization using the
methods described in Section IV.
134
Figure 6. Comparison of temperature estimates from popular thermometers and this
study for five MORB and three OIB locations. MORB compositions are primitive
glasses from the PetDB database and OIB compositions are calculated by Putirka et al.
(2007). Temperatures were calculated using pressure estimates from the Lee et al. (2009)
barometer and our spinel lherzolite olivine expression.
Figure 7. Effect of olivine-only fractionation correction on mantle melts that
experienced a multiphase fractionation history. Parental liquids are polybaric, near
fractional aggregate magma compositions calculated by Kinzler (1997). The composition
of these parental liquids (black lines and dots) are shown during olivine fractionation,
followed by olivine + plagioclase fractionation and in the case of the high-pressure
parental liquid, a third period of olivine + plagioclase + clinopyroxene fractionation, until
MgO in the liquid equals 8 wt.%. The fractionated liquid is then corrected for olivineonly fractionation until it is in equilibrium with Fo90 olivine using a Herzberg and
Asimow (2008) olivine-liquid Fe-Mg Kd. Also shown are contours of the liquid
composition using an olivine-liquid Fe-Mg Kd of 0.3 (dashed gray lines).
Figure 8. a) Trace element abundance of basalts from the Diamond Crater region,
Oregon and modeled 1-15% melts of a spinel lherzolite mantle. b) Rare earth element
abundances of the same samples and modeled melts.
135
Figure 9. Results of our forward spinel lherzolite melting model relative to fractionationcorrected basaltic liquids from the Diamond Crater area, Oregon.
Table 1. Experiments used in our plagioclase and spinel lherzolite melting model
calibration. Star indicates experiments used in the MIT + CMAS(N) regression presented
in Table 5.
Table 2. Composition of experimental starting materials for experiments conducted in
this study.
Table 3. Experimental run conditions and phases present for experiments conducted in
this study.
Table 4. Normalized major element composition of experimental phases from our
experiments, as determined by electron microprobe analysis.
Table 5. Model equations produced by multiple linear regression of experimental
plagioclase and spinel lherzolite liquids. Number in italics indicates the standard error for
the coefficient above.
Table 6. Geochemistry of primitive basalts from Diamond Crater area, Oregon. Major
element XRF and trace element ICPMS analyses were conducted at the WSU
GeoAnalytical Lab.
136
Table 7. Pressure and temperature estimates for Diamond Crater area samples produced
by our spinel lherzolite thermometer and barometer (Oliv expression).
Table 8. Starting compositions, run conditions and results from our spinel lherzolite
forward melting model, as used to replicate the parental liquids for basaltic lavas from the
Diamond Crater area, Oregon. Starting bulk peridotite compositions used in the
modeling are a xenolith of depleted metasomatized mantle from Patagonia (Ntaflos et al.,
2007) and the primitive mantle composition of Hart and Zindler (1986), renormalized to
exclude MnO, NiO and P2 0 5 . The Diamond Crater compositions were corrected for
olivine + plagioclase fractionation using the methods discussed in Section V until they
reached equilibrium with Fo90 olivine.
137
Appendix 1.
Both the plagioclase and spinel lherzolite expressions presented in Section III and
Table 5 are the product of regressing only those experiments analyzed on the microprobe
at MIT and experiments done on the simple systems CMAS and CMASN. Here we
present an additional set of lherzolite regressions determined from the full dataset minus
several experiments as discussed below. There are only five additional experiments that
can be included in the full dataset for plagioclase lherzolite melting. These experiments
are from Dunn and Sen (1994), Falloon et al. (1997), and Falloon et al., (1999).
Figure 2a-c reveals that the suite of experiments from Schwab and Johnston
(2001) and Draper and Johnston (1992) define a slope <1, that does not correspond to the
slope of ~1 defined by the majority of the data. For this reason, we have chosen to
exclude the data of Draper and Johnston (1992) and Schwab and Johnston (2001) from
our regressions of the full dataset.
Full Dataset Regression:
r2
Plag Lherzolite
T = 1216 + 98.8 (P) - 73.5 (1-Mg#) - 168.0 (NaK#) + 24.8 (TiO 2)- 9.70 (K20)
(8.36)
(10.0)
(29.0)
(30.2)
(7.47)
Oliv = 0.128 + 0.077 (P) + 0.165 (1-Mg#) - 0.271 (NaK#) - 0.009 (TiO 2) + 0.017 (K20)
(0.004) (0.005)
Cpx
-
(0.014)
(0.014)
(0.003)
(0.010)
(0.029)
(0.030)
(0.007)
(0.031)
(0.032)
(0.008)
(0.98)
(0.009)
Qtz = 0.216 - 0.152 (P) + 0.018 (1-Mg#) - 0.287 (NaK#) - 0.010 (TiO 2) + 0.043 (K20)
(0.009) (0.011)
(0.032)
(0.034)
(0.008)
(0.010)
138
(0.82)
(0.009)
Plag = 0.411 + 0.139 (P) - 0.130 (1-Mg#) + 0.639 (NaK#) + 0.014 (TiO 2) - 0.035 (K 2 0)
(0.009) (0.006)
(0.96)
(0.004)
0.244 - 0.064 (P) - 0.053 (1-Mg#) - 0.081 (NaK#) + 0.005 (TiO 2) - 0.025 (K 2 0)
(0.008)
(0.95)
(8.65)
(0.97)
Spinel Lherzolite
T = 1187 + 135.8 (P) - 41.9 (1-Mg#) - 97.1 (NaK#) - 1.63 (TiO2 ) - 13.92 (K2 0)
(8.94)
(4.94)
(32.8)
(25.2)
(10.6)
Oliv = 0.116 + 0.090 (P) + 0.195 (1-Mg#) - 0.262 (NaK#) - 0.009 (TiO 2) + 0.011 (K20)
(0.010)
(0.006)
(0.038)
(0.029)
(0.012)
(0.008)
(0.051)
(0.039)
(0.016)
(0.73)
(0.004)
Cpx = 0.191 - 0.003 (P) + 0.038 (1-Mg#) - 0.243 (NaK#) + 0.015 (TiO 2) + 0.012 (K2 0)
(0.014)
(0.93)
(3.25)
(0.35)
(0.005)
Plag = 0.559 - 0.046 (P) - 0.199 (1-Mg#) + 0.945 (NaK#) - 0.007 (TiO 2) - 0.042 (K20)
(0.015) (0.009)
(0.057)
(0.044)
(0.018)
(0.006)
(0.86)
Qtz = 0.134 - 0.042 (P) - 0.034 (1-Mg#) - 0.439 (NaK#) + 0.001 (TiO 2) + 0.019 (K20)
(0.009) (0.005)
(0.034)
(0.026)
(0.011)
(0.003)
(0.89)
139
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1
0.2
0.3
0.5
0.4
0.6
Mg#
3
e.
A
0
0000
A
00
0
o0
00 00
0.8
0.9
1
0
00
0
2.5
O
00
00
2
0
0
00
1
0
00
0
3-
:1.5
0
1
0
3.5
0
0.9
4
2.5
2
0.8
.
4
0
0
0.7
Mg#
0
.
000 0
0
1.5
0
0
0
1.
*0
0.5
O
43
Val
0 00
0.5
***
0 00
C4 n
CBS
0
0
0.2
0.3
0.4
0.5
0.6
0.7
0.8
0.2
0.9
0.3
0.4
0.5
0.7
0.6
Mg#
Mg#
100
o
Liquid in equilibrium with
g.
plagioclase lherzolite
1600
Liquid in equilibrium with
spinel lherzolite
Liquid in equilibrium with
plagioclase lherzolite (this study)
1500
00
99
1400
0 41063l
le
Figure 1.
1000
0
0.5
1
1.5
2
P (GPa)
140
2.5
3
3.5
08
0.75
0.9 -
a.
*Other
*CMAS(N)
0.851
*Schwab & Johnston,2201
61
*Draper & Johnston, 1992
o
.
+
0.75
0.65
0 75208
0.55
05
e. c'
0'-- -~
10.65
04
R2 =0 .86
Plag Error
**
045
-Av.
04
0,45
0.5
0.55
0.6
065
0.7
075
08
Av. Plag Error = 0.02
Av. P Error =0.39 GPa
045
=
0.03
0.85
c
0.55
047
0.9
ExperimentalPlag
0.51
R20 8
0.4
0.45
20.4
0.4
0.35
0
S035
.
0.3
0.25j
0.2~
0L5
ols5
0,25
r3
01 r
0.1
015
0,2
0.25
0.3 0.35 0.4 0.45
Experimental011y
0.5
0.55
R2 =0.63
Av. Oliv Error = 0.02
Av. P Error = 0.25 GPa
0.2
R2=0.73
Av. Oliv Error = 0.02
.
0.1
0.1
0.6
015
0.2
025
0.3
0.35
0.4
0.45
0,5
055
0.6
ExperimentalOlily
0.5
A6.
0.45
- -
- - -
_-
-
-
0.45
0.4
0.4
0.35
035
0.3
0.3
0,25
025
0.15
0.15
0.2
0.1
R2=0.35
Av. Cpx Error =0.03
05
0
005
0.1
0.15
02
025
0.3
0.35
04
045
R2=0.20
Av. Cpx Error = 0.01
Av. P Error = 4.35 GPa
0.05
0
0.5
0.05
01
0.15
ExperimentalCpx
0.2
025
0.3
0.35
04
0.45
0
ExperimentalCpx
0.35
0.35-
0.3
0.3
025
0.25
o2
0,2
0.15
0.15
-Y
01
0'1
0.05
.
0
0.05
0A
-0,05
4
-0.1
-0 * .5
R 2=0 .89
;
Av. Qtz Error =0.02
-0.15
-01
-005
0
0.05
01
0.15
0,2
0.25
0.3
-0.15
-0,15
0.35
ExperimentalQt
-01
-0.05
0
R2=0.92
Av. Qtz Error = 0.01
Av. P Error = 0.45 GPa
0.05
0.1
0.15
ExperimentalQtz
Figure 2.
141
0.2
0.25
03
0.35
**
PP(GPa)
spine1
pk1giole
Figure 4.
143
a.
Cpx
Oliv+Cpx
doOiv+Plag
5
b.
21 kbar
Oliv
%
Plag
13
Cpx
Y
e
y y yvYy
1
EMP Error
Qtz
Oliv
C.
Cpx
Oliv
Plag
Figure 5.
144
1600
-
m
M
m
m
m
a.
Fordetal.(1983)
Kinzler
and Grove(1992a)
Langmuiretal. (1992)
Putirkaet al. (2007)
Leeetal. (2009)
ThisStudy
b. -
1500
-
1500
1400
1400 -
1300
13001-
1200
Olivine Only Correction
Lee et al. 2009 Pressures
Azores
I
EPR 9N
I
Kane FZ
I
Reykjanes
I
SWIR
1700
Azores
II
Kane FZ
EPR 9N
Olivine + Plag Correction
This Study Pressures
I
I
SWIR
Reykjanes
C.
1600
1500
1400
I
4
I
I
Hawaii
Iceland
Lee et al. (2009) Pressures
I
Samoa
e.
Comps,
Leeet al. (2009)
OlivineOnlyCorrected
Comps,Leeet al. (2009)
Olivine+ PlagCorrected
Camps,ThisStudy
OlivineOnlyCorrected
ThisStudy
Comps,
Olivine+ PlogCorrected
Average Ptndling hW
Azores
I
EPR 9N
-0-
Leeet aL.(2009)
-()-
ThisStudy
f.
0
I
Kane FZ
I
Reykanes
I
R
I
SWIR
Figure 6.
145
Iceland
I
Hawaii
I
Samoa
I
-
12
10
MgO wt%
Figure 7.
146
0
1% melt
2% melt
6% melt
--
10% melt
15% melt
4E
02
-e-DC 07-12
"DC 07-13
DC 07-16
-
*ch8329
z
ch8343
-
ch8357b
1
Rb Ba Th U Nb Ta K La Ce Pb Pr Sr P Nd Zr Hf Sm Eu T Gd Dy Y Ho Er Yb
100
-1% melt
% melt
% melt
10%melt
-15%
10
melt
*DC 07-12
DC 07-13
-DC 07-16
**ch8329
z0
- h8343
"$"sch8357b
La
Ce
Pr
Nd
Sm
Eu
Figure 8.
147
Gd
Dy
Ho
Er
Yb
-
K2 0 (Wt.%)
FeO (wt.%)
.5
g
+
.0
OLi
El
.5
+
03
+
Na20 (wt.%)
.5
2.5
2.0
3.0
3.5
MgO (wt.%)
7.0
4. 0
10.0
10.5
11.0
11.5
12.0
12.5
1)
CaO (wt.%)
KEY
+
11.5-
+ +
11.0-
Patagonia xenolith 1.8-1.3 GPo accumulated
near-fractionalmelt
Patagonia xenolith 1.6-1.0 GPaaccumulated
near-fractionalmelt
M Patagonia xenolith 1.3GPa 6%batch melt
03
+*
*
+
10.5 -
o
o
Patagonia xenolith 1.3 GPa8%batch melt
0
Patagonia xenolith 1.6GPa 6%batch melt
Patagonia xenolith 1.6GPa 4% batch melt
SH&Z Primitive Mantle 1.3GPa6%batch melt
L
10.0-
0S
9.5
46.5
H &Z Primitive Mantle 1.6GPa 6%batch melt
+
S'02 (Wt*%
47.0
47.5
48.0
H &Z Primitive Mantle 1.3GPa 8% batch melt
48.5
49.0
49.5
Figure 9
148
Fractionation-corrected Diamond Crater area
basalts
1
Table 1. Experiments used in our plagioclase and spinel lherzolite melting model calibration. Star indicate the experiments used in the MIT + CMAS(N) regression presented in
Table 5.
IReference
Plagioclase Lherzolite
Baker, M.B., Grove, T.L., Price, R. (1994)*
Dunn, T., Sen, C. (1994)
Falloon, T.J., Green, D.H., O'Neill, H.St.C., Hibberson, W.O. (1997)
Grove and Juster, 1989*
Kinzler, R.J., Grove, T.L. (1992)*
Presnall et al. 1979*
Walter and Presnall (1994)*
Spinel Lherzolite
Bartels, K.S., Kinzler, R.J., Grove, T.L. (1991)*
Falloon, T.J., Danyushevsky, L.V., Green, D.H. (2001)
Falloon, T.J., Green, D.H., O'Neill, H.St.C., Hibberson, W.O. (1997)
Gaetani, G.A., Grove, T.L. (1998)*
Holbig, E. S., Grove, T.L. (2007)*
Kinzler, R.J. (1997)*
Kinzler, R.J., Grove, T.L. (1992)*
Laporte, D., Toplis, M.J., Seyler, M., Devidal, J-L. (2004)
Pichavant, M., Mysen, B.O., MacDonald, R. (2002)
Pickering-Witter, J., Johnston, A.D. (2000)
Presnall et al. 1979*
Villiger, S., Ulmer, P., Muntener, 0., and Thompson, A.B. (2004)
Walter and Presnall (1994)*
Wasylenki, L.E., Baker, M.B., Kent, A.J.R., Stolper, E.M. (2003)
Startina Materials
basalts, basaltic andesites, andesites, Mt. Shasta, CA
basalt and basaltic andesite, Mt. Adams, Washington
peridotite and corresponding low degree melts synthesized from oxides
basaltic andesite, Medicine Lake, CA and seeds of natural minerals
MORB, natural samples and synthetic oxide mixtures
CMAS synthesized from oxides
CMASN synthesized from oxides
high alumina basalts, Medicine Lake, CA
peridotite and corresponding low degree melts synthesized from oxides
see above
basalts, Medicine Lake, CA + synthetic oxide peridotite
shoshonite, Tibet
peridotite synthesized from oxides
see above
depleted peridotite synthesized from minerals from a spinel harzburgite, Kane Fracture Zone
high-MgO basalt, Soufriere, St. Vincent
peridotite synthesized from minerals from Kilbourne Hole, Hawaii
see above
tholeiitic basalt, synthetized from oxides and seeds of natural minerals
see above
peridotite synthesized from minerals from Kilbourne Hole, Hawaii + synthetic diopside
Table 2. Composition of experimental starting materials for
experiments conducted in this study.
JC-30B Kragero Opx JC-30B + 15% OPX
48.92
57.3
50.18
SiO2
1.49
1.74
0.06
TiO2
13.73
0.1
A1203
16.13
0
Cr203
0
0
9.41
9.49
FeO
9.39
0.15
0.14
MnO
0.15
12.86
MgO
9.22
33.5
8.5
0.26
CaO
9.95
2.64
Na20
0
3.1
0.82
K20
0.97
0
0.37
0
P205
0.43
0.71
Mg#
0.86
0.64
150
SpLherz
47.09
1.01
16.79
0
8.41
0
12.62
10.84
2.4
0.62
0.22
0.73
Table 3. Experimental run conditions and phases present for experiments conducted in this study
Expt.
Starting Mat. P (kbars)
T (*C)
Duration (hrs) Phases
ol
opx
B1204
JC30B+opx
8
1270
13.3
x
x
B1205
JC30B+opx
10
1250
75.3
x
x
B1149
JC30B+opx
10
1270
10
x
x
B1201
JC30B+opx
14
1270
28.8
x
C449
C450
C448
SpLM
SpLM
SpLM
17
17
17
1313
1333
1353
97
67.3
71.8
x
cpx
x
x
x
x
x
x
plag
x
x
x
sp
glass
x
x
x
x
x
x
x
x
x
Expt
Phase
SpLM
C450
cpx
17 kbar, 1333C
spinel
#
Si02
Ti02
A1203
Cr203
FeO
MnO
MgO
CaO
Na20
5
50.01
0.21
0.55
0.27
0.6
0.17
0.21
0.02
11.5
1.68
67.71
0.71
0
6.52
0.61
9.8
0.3
0.04
0.06
0
16.65
2.04
21.45
0.25
13.43
0.8
0.19
0.03
1.25
0.42
39.69
0.24
50.57
0.28
0.11
0.03
47.16
0.68
48.87
1.34
0.05
0.02
0.65
0.06
0.18
0.02
1.09
0.04
0.67
0.64
0.16
0.01
10.41
0.27
68.04
0.6
18.57
0.14
19.07
0.98
13.18
0.14
4.7
0.4
8.71
0.23
8.38
0.11
9.05
0.61
0.02
0.02
0.01
0.01
0.01
0.02
0.01
0.01
0.2
0.17
46.62
0.4
16.72
0.39
22.8
0.17
9.26
0.14
7.56
1.72
0.27
0.03
15.62
0.64
0.11
0.04
8.36
0.09
7.26
1.63
39.38
0.61
51.36
0.88
52.41
0.67
53.97
1.71
51.62
0.22
0.09
0.02
1.13
0.16
0.60
0.23
0.24
0.23
5.74
0.94
8.48
0.61
25.70
4.82
17.27
0.19
0.00
0.01
0.00
0.01
0.00
0.00
16.96
0.49
6.34
0.49
10.37
1.29
2.21
3.51
8.90
0.13
0.27
0.02
0.19
0.03
0.21
0.02
42.67
0.77
18.58
1.27
23.81
3.01
1.51
2.47
6.24
0.11
0.38
0.16
16.17
1.41
3.42
2.30
10.62
0.93
8.77
0.14
0.54
0.16
0.67
0.78
4.94
0.61
4.03
0.17
0.09
0.01
6.34
0.85
7.15
0.67
0.01
0.01
0.04
0.01
0.00
0.00
0.29
0.02
0.22
0.05
0.21
0.01
0.18
0.02
39.77
0.51
17.57
0.70
23.72
0.58
0.13
0.03
5.48
0.17
0.29
0.03
15.32
1.33
4.22
0.46
10.82
0.66
7.09
0.07
0.58
0.07
0.83
0.09
5.10
0.43
5.39
0.21
8
0
K20
NiO
P205
Total
100.25
100.07
0.06
0.09
glass
C449
oliv
17 kbar, 1313C
cpx
8
10
spinel
9
glass
15
C448
glass
17 kbar, 1353C
JC-30B + OPX
B1204
oliv
8 kbar, 1270C
cpx
6
9
7
opx
5
plag
11
glass
10
1.98
0.10
0
0.01
0.03
0.01
0.02
0.01
0.04
0.08
0.02
0.02
0.16
0.02
1.32
0.15
0.01
0.01
6.42
0.19
6.09
1.98
0.58
0.02
0.77
0.18
0.02
0.03
10
32.28
17
opx
3
plag
9
glass
9
51.27
0.42
53.03
0.53
54.79
0.94
50.17
0.47
0.10
0.01
0.99
0.14
0.78
0.01
28.51
0.76
2.61
0.03
17.38
0.20
0.00
0.00
21.14
0.29
7.66
0.57
10.05
0.25
0.34
0.05
10.40
0.19
50.13
0.67
51.70
0.41
51.54
0.32
0.91
0.14
0.48
0.06
2.06
0.03
9.18
0.78
8.32
0.38
19.22
0.23
0.01
0.01
0.00
0.00
0.03
0.03
8.19
0.34
11.27
0.10
9.67
0.12
0.21
0.03
0.21
0.03
0.15
0.01
17.10
0.68
25.95
0.46
4.59
0.41
13.31
0.67
1.82
0.08
6.85
0.15
0.95
0.15
0.23
0.08
4.59
0.17
38.48
0.18
52.19
0.18
55.81
1.87
54.25
0.30
50.00
0.39
0.07
0.01
0.67
0.06
0.24
0.27
0.09
0.01
2.05
0.16
0.11
0.03
5.33
0.51
2.10
2.89
28.54
0.18
18.08
0.18
0.01
0.01
0.01
0.01
0.02
0.01
0.04
0.02
0.01
0.01
19.17
0.63
7.21
0.19
9.47
0.40
0.41
0.02
9.83
0.28
0.19
0.03
0.15
0.02
0.09
0.08
0.01
0.01
0.16
0.02
41.60
0.49
20.03
0.32
30.08
3.98
0.11
0.01
6.41
0.12
0.36
0.02
13.91
0.35
1.53
1.75
11.32
0.22
7.98
0.04
0.51
0.04
0.28
0.47
4.90
0.13
4.40
0.26
glass
18
4
9
B1149
oliv
10 kbar, 1270C
cpx
5
opx
26
plag
7
glass
8
5
101.29
100.76
99.79
0.22
0.02
0.31
0.08
100.07
99.19
100.80
100.96
99.96
100.41
0.46
0.33
1.01
0.03
100.87
0.02
0.02
B1205
oliv
10 kbar, 1250C
cpx
B1201
cpx
14 kbar, 1270C
opx
0.01
0.01
0
0
0
0
0
0.01
0.1
0.14
101.39
100.80
99.83
0.31
0.04
1.30
0.05
100.00
99.95
100.57
0.38
100.71
1.30
0.06
100.26
0.27
0.01
0.02
99.72
101.34
101.14
0.34
0.01
1.10
0.05
100.90
100.73
Table 5. Model equations produced by multiple linear regression of experimental plagioclase and spinel Iherzolite
liquids. Number in italics indicates the standard error for the coefficient above
Plagioclase Lherzolite
T = 1220 + 90.8 (P) - 21.1
(1-Mg#) - 160.9 (NaK#) + 6.28 (TiO 2 ) - 13.4 (K 2 0)
0.97
7.63
Oliv = 0.130
0.016
Cpx = 0.252
0.008
Plag = 0.410
0.01
Qtz = 0.213
0.009
10.7
+ 0.076
0.009
- 0.088
0.011
+ 0.138
0.014
- 0.144
0.012
29.9
(P)
-+ 0.151
33.6
8.74
10.23
(1-Mg#) - 0.272 (NaK#) - 0.003 (TiO 2) + 0.016 (K20)
0.96
0.051
- 0.018
0.005
0.006
(P) - 0.047 (1-Mg#) - 0.005 (NaK#) + 0.005 (Ti0 2 ) - 0.005 (K20)
0.03
0.033
0.009
0.01
(P) - 0.135 (1-Mg#) + 0.647 (NaK#) + 0.016 (TiO 2 ) - 0.036 (K20)
0.039
0.044
0.011
0.013
(P) - 0.005 (1-Mg#) - 0.303 (NaK#) - 0.003 (TiO 2 ) + 0.050 (K 0)
2
0.034
0.034
0.009
0.01
0.88
0.97
0.97
Spinel Lherzolite
T = 1200 +
10.62
Oliv = 0.136 +
0.016
Cpx = 0.148 +
0.013
Plag = 0.603 0.019
Qtz = 0.113 0.011
128.6 (P) - 82.06 (1-Mg#) - 104.2 (NaK#) + 9.72
5.83
33.8
29.8
10.52
(TiO 2) - 8.05
(K20)
0.94
(K20)
0.63
7.81
0.079 (P) + 0.197 (1-Mg#) - 0.251 (NaK#) - 0.016 (TiO 2) + 0.026
0.009
0.051
0.045
0.016
0.012
0.004 (P) - 0.024 (1-Mg#) - 0.097 (NaK#) + 0.012 (TiO 2) + 0.008 (K20)
0.007
0.041
0.036
0.013
0.009
0.057 (P) - 0.187 (1-Mg#) + 0.852 (NaK#) + 0.002 (TiO 2) - 0.047 (K20)
0.01
0.06
0.053
0.018
0.014
0.025 (P) + 0.013 (1-Mg#) - 0.504 (NaK#) + 0.003 (TiO 2) + 0.013 (K20)
0.006
0.034
0.03
0.011
0.008
153
0.2
0.88
0.92
Table 6. Geochemistry of samples from Diamond Crater area, Oregon. Major element XRF and trace element ICPMS analyses were
conducted at the WSU GeoAnalvtical Lab.
DC07-12
DC07-13
DC 07-14
DC 07-15
DC 07-16
ch832 9
ch8343
ch8357b
43.05
43.05
43.05
43.05
Latitude
43.05
44.9
44.10
43.15
Longitude
-118.43
-118.43
-118.46
-118.43
-118.43
-117.7
-118.08
-118.60
0
XRF Normalized Major Elements (Wt /):
48.33
Si02
47.98
47.90
48.94
47.83
49.7
47.70
48.00
1.171
1.247
1.710
TiO2
1.297
1.377
1.25
1.23
0.95
A1203
17.48
17.71
17.07
15.40
17.18
15.5
15.50
16.50
11.89
10.19
FeO*
10.11
9.93
10.22
9.27
9.81
9.63
0.181
0.215
0.17
MnO
0.181
0.176
0.181
0.15
0.17
6.74
MgO
8.71
8.82
8.20
9.12
10.30
9.65
9.36
11.46
11.38
11.10
CaO
10.72
11.15
10.39
9.53
10.80
Na20
3.14
2.92
2.86
3.01
2.38
2.44
2.71
2.17
K20
0.30
0.39
0.45
0.40
0.28
0.61
0.78
0.32
P205
0.216
0.140
0.161
0.241
0.31
0.17
0.201
0.33
Total
99.99
99.20
99.41
100.22
100.32
99.92
99.40
100.16
ICP-MS Trace Elements (ppm):
La
7.363
4.986
5.655
7.177
8.257
12.31
5.35
16.63
14.132
Ce
17.767
12.267
19.872
12.73
17.988
26.38
33.58
Pr
2.601
1.904
2.153
2.766
2.889
4.25
1.89
3.63
Nd
12.181
9.287
10.573
13.620
13.510
15.70
17.70
9.00
Sm
3.385
2.762
3.076
4.068
2.59
3.671
3.75
4.10
Eu
1.247
1.433
1.367
1.145
1.564
1.29
1.43
1.04
Gd
3.909
3.315
3.667
4.961
4.123
4.23
3.13
3.88
Tb
0.676
0.590
0.665
0.887
0.717
0.70
0.65
0.57
4.352
3.878
4.523
Dy
4.321
5.727
4.25
3.79
3.91
0.824
0.919
1.225
0.927
Ho
0.946
0.88
0.82
0.81
Er
2.552
2.277
2.519
3.368
2.583
2.17
2.33
2.25
Tm
0.368
0.332
0.367
0.492
0.34
0.379
0.31
0.33
Yb
2.294
2.082
2.308
3.093
2.367
1.94
2.16
2.10
Lu
0.364
0.330
0.365
0.490
0.371
0.30
0.34
0.34
144.222
199.480
193.546
Ba
154.809
203.529
255
269
158
0.350
0.428
Th
0.476
0.304
0.502
0.96
2.26
0.55
Nb
5.810
3.412
4.238
5.123
6.887
12.13
26.48
4.14
Y
22.661
30.525
23.547
22.865
20.585
20.75
22.21
20.16
2.260
1.766
Hf
1.985
2.617
2.486
2.03
2.48
1.73
Ta
0.386
0.221
0.277
0.334
0.430
0.129
0.60
0.173
0.110
U
0.159
0.189
0.35
0.13
Pb
0.946
0.879
0.769
1.009
1.49
1.016
2.60
0.93
2.968
3.242
17.3
4.219
Rb
3.711
5.029
7.1
2.5
0.035
0.039
0.057
0.22
0.11
0.053
Cs
0.054
0.02
Sr
308.900
272.669
273.986
253.609
398
315
321.051
223
47.886
29.4
31.222
34.169
Sc
40.970
30.236
33.9
34.8
101.476
99.128
Zr
90.391
66.597
74.123
76
98
66
Ba/Nb
33.314
42.275
36.526
38.938
154
29.554
21.000
10.140
38.090
Table 7. Pressure and temperature estimates
for Diamond Crater area samples produced
by our spinel lherzolite thermometer and
barometer (Oliv expression).
P(GPa)
T(*C)
DC 07-12
1.50
1357
DC 07-13
1.39
1345
DC 07-14
1.32
1336
DC 07-15
1.39
1346
DC 07-16
1.60
1370
ch8329
1.38
1341
ch8343
1.29
1330
ch8357b
1.22
1323
155
Table 8.
Total F
dF/dP
Patagonia Xenolith
H & Z Primitive Mantle
T
0.47
FeO
8.8
7.58
MgO
43.6
37.94
CaO
1.32
3.22
Na20
0.14
0.33
K20
0.02
0.032
17.56
17.91
18.04
17.92
17.65
17.48
0.13
0.13
0.14
0.15
0.13
0.14
8.50
8.14
7.94
7.93
8.40
8.49
12.19
11.70
11.89
12.27
12.28
12.75
10.69
11.13
11.37
11.73
10.61
11.09
2.42
2.22
2.04
1.70
2.46
2.03
0.38
0.33
0.32
0.24
0.47
0.32
1329
1333
1362
1367
1.04
0.93
1.15
18.34
18.24
17.74
0.11
0.11
0.11
7.42
7.57
8.04
10.54
11.05
11.46
10.00
10.61
9.92
3.48
2.93
3.27
0.53
0.38
0.53
1320
1326
1361
46.90
46.84
46.94
46.70
46.92
48.95
47.11
47.49
0.78
0.73
0.72
0.75
0.84
0.87
0.84
0.61
19.56
19.64
19.55
19.55
19.28
18.38
18.54
19.53
7.64
7.68
7.59
7.75
7.76
7.50
7.73
7.44
11.53
11.44
11.37
11.54
11.72
11.12
11.58
11.14
10.94
11.26
11.45
11.38
10.68
10.43
11.62
11.45
2.29
2.15
2.17
2.08
2.39
2.08
1.83
2.02
0.24
0.17
0.17
0.17
0.27
0.43
0.53
0.21
Si02
45
46.16
TiO2
0.06
0.18
A1203
1.65
4.08
Cr203
47.56
47.94
47.80
47.65
47.46
47.23
0.56
0.51
0.45
0.42
0.54
0.48
48.54
48.52
47.79
Patagonia Xenolith
1.8-1.3 Gpa
Near-Fractional
6%
1%/0.1 Gpa
1.6-1.0 Gpa
Near-Fractional
7%
1%/0.1 Gpa
1.3 Gpa
1.3 Gpa
Batch
Batch
6%
8%
1.6 Gpa
Batch
4%
1.6 Gpa
Batch
6%
H & Z Primitive Mantle
1.3 Gpa
1.3 Gpa
Batch
Batch
1.6 Gpa
Batch
6
%
8%
6%
Fractionation-Corrected Diamond Crater Area Samples
DC 07-12
DC 07-13
DC 07-14
DC 07-15
DC 07-16
ch8329
ch8343
ch8357b
157
Chapter 4:
Constraints on the Depths and Temperatures of
Anhydrous Asthenospheric Melting and
the Location of the Lithosphere-Asthenosphere Boundary
Beneath southern Oregon and northern California
158
Abstract
An extensive geochemical dataset of young (<10.5 Ma) primitive basaltic lavas from
across the northernmost Basin and Range and the central - southern Cascades is used to
calculate the depths and temperatures of asthenospheric melting in this region. In
conjunction with recent geophysical observations, these calculations place primary
constraints on the tectono-magmatic processes driving anhydrous melting in a modem
convergent margin and back-arc. The basalts have high Mg#s (1OOMg/(Mg + FeT) at low
SiO 2 (<52 wt.%), low phenocryst abundances, and trace element concentrations such that
we consider them to be primitive. The minimum depth of melting below Oregon's High
Lava Plains, a geographic sub-province of the Basin and Range, decreases towards the
west, with melting occurring at -45 km below Jordan Valley volcanic center, -30 km
below Newberry volcano, and -25 km below Crater Lake volcano on the Cascades arc
axis. To the south, the minimum depth of melting also decreases towards the west, from
-40 km below the Modoc Plateau to 33 km below Medicine Lake and Mt. Shasta
volcanoes. Further south at Mt. Lassen, the southernmost Cascades volcanic center, the
minimum depth of melting is 35 km. All basalts in this study originated at 1250-1320'C
and the calculated minimum depths of asthenospheric melting are very close to Moho
depths as determined from a number of regional geophysical studies. These observations
point to the hot nature of the mantle immediately below the Moho and suggest the Moho
and lithosphere-asthenosphere boundary occur at approximately the same depth in this
region. <10.5 Ma anhydrous mantle melting in southern Oregon and northern California
was likely driven by a combination of corner flow in the mantle wedge, toroidal flow
around the southern edge of the subducting Juan de Fuca and Gorda plates, and crustal
159
extension-related upwelling, not a mantle plume. Geodynamic models of mantle flow
also indicate the thin mechanical lithosphere is an important factor in the generating the
observed conditions in the mantle.
160
1. Introduction:
The goal of this study is to use the chemical characteristics of primitive basalts from
southern Oregon and northern California in the northwestern Unites States to further our
understanding of the processes driving anhydrous mantle melting beneath convergent
margins and extending into the back-arc. Starting in the Oligocene, a significant plate
reorganization occurred in the region that is now the western coast of North American,
resulting in a period where the relationship between the plate tectonic driving forces and
the composition and location of volcanism is poorly understood (e.g., Atwater, 1970;
Christiansen and McKee, 1978; Hart and Carlson, 1987). Specifically, a notable number
of volcanic centers have formed since 25 Ma in the area between the southern Cascades
volcanic arc and the Oregon-Idaho border that have been ascribed to a variety of tectonomagmatic processes including, extension in the Basin and Range Province (Christiansen
and McKee, 1978; Cross and Pilger, 1982; Christiansen et al., 2002), lithospheric
delamination (Hales et al., 2005; West et al., 2009), back-arc extension (Christiansen and
McKee, 1978) and mantle flow related to the geometry of the subducting Juan de Fuca
slab (Carlson and Hart, 1987; Humphreys et al., 2000; Faccenna et al., 2010). Recent
studies have also attributed the volcanic centers in question to mantle plume-related
volcanism (e.g., Geist and Richards, 1993; Camp and Ross, 2004) because of the nearby
eruption of >10 5 km 3 of evolved basaltic - basaltic andesite lavas ca. 16.5-14 Ma, which
formed the Steens Mountain and Columbia River Plateau flood basalt provinces.
However, <10.5 Ma basaltic volcanism in the area of interest is more limited in volume
(600-1000 km 3) and consists predominantly of anhydrous high alumina olivine tholeiites
(HAOT) that have experienced very little crystal fractionation and crustal contamination
161
(Hart et al., 1984; Draper, 1991; Hart et al., 1997). By utilizing primitive basaltic lavas
from these volcanic centers that approximate liquid compositions and correcting for the
effects of small amounts of fractional crystallization, we can infer the locations and
conditions in the mantle from which these <10.5 Ma magmas segregated, and
subsequently identify the tectonic and magmatic driving forces responsible for their
formation.
To calculate the depth and temperature of melt extraction from the mantle for the
primitive basaltic lavas, we use a new model for melting variably depleted and
metasomatized upper mantle (Till et al., 2011*). The calculated depths of melt extraction
are then compared to the location of the Moho and the lithosphere-asthenosphere
boundary (LAB) in the region as constrained by teleseismic receiver functions data (e.g.,
Gashawbeza et al., 2008; Eager et al., 2011) and Rayleigh Wave tomography (Wagner et
al., 2010) and other seismic refraction surveys (e.g., Leaver et al., 1984; Zucca et al.,
1986), in order to place new constraints on the thickness of the mechanical lithosphere.
In the broadest sense, the LAB can be defined as the boundary between the rigid
material that composes the Earth's tectonic plates (i.e., the lithosphere) and the
underlying material that undergoes solid-state creep such that it behaves as viscous fluid
on geologic timescales (i.e., the asthenosphere). However, in practice there is no unified
theory for what comprises the lithosphere and it is alternately defined by its mechanical,
seismic, electric, thermal, or chemical properties. As such multiple depths for the LAB
can be calculated in a given area depending on the property used define the LAB and the
method used to calculate that property. The mechanical lithosphere is defined as the
Same as Chapter 3 in Till, C.B., 2011, Melt Generation in the Earth's Mantle at Convergent
Plate Margins, MIT PhD Thesis.
162
portion of the Earth's upper thermal boundary layer that behaves elastically, and the
depth of the mechanical LAB is usually inferred from the depths of intraplate earthquakes
and plate flexure models. By calculating depths of melt extraction from the
asthenosphere in southern Oregon and northern California, we produce an independent
estimate of the maximum thickness of the mechanical lithosphere in the region that can
be used to further evaluate the nature of the lithosphere, especially in regions with a
prolonged history of volcanism.
2. Tectonic and Volcanic History
The western United States has a rich and varied volcanic and tectonic history (see
reviews in Humphreys and Coblentz, 2007). Extensive volcanism in the western United
States began in the early Cenozoic with the ignimbrite flare-up, a period of large volume
silicic magmatism that swept through the western United States between 50 and 25 Ma
(e.g., Lipman et al., 1972; Christiansen and Yeats, 1992). In eastern Oregon, a period of
silicic volcanism was followed by a period of massive flood basalt eruptions near the
Oregon/Idaho/Nevada border that produced the Steens Mountain basalts from 16.6-15.5
Ma (Brueseke et al., 2007). This flood basalt volcanism migrated north along the western
edge of Precambrian North America to form the Columbia River Plateau basalts between
16.5 and 6 Ma (Waters, 1961; Swanson and Wright, 1979; Carlson, 1984; Tolan et al.,
1989). To the west, Cascadian volcanism produced by oblique subduction of the Juan de
Fuca plate began ca. 37 Ma and continues to the present. The Basin and Range province
impinges on the volcanic highlands of southern Oregon and Cascades arc, and as such
163
these regions have been characterized by NW-SE to N-S normal faulting and associated
extension-related volcanism since -14 Ma.
This study in part focuses on lavas from the High Lava Plains (HLP) of central and
southeastern Oregon, a geographic sub-province of the northern Basin and Range. This
10 Ma bimodal volcanic field seemingly connects the Yellowstone-Snake River Plane
(YSRP) track to the east, with the Cascades arc to the west and its tectono-magmatic
origin has been a topic of debate since the 1970's (e.g., MacLeod et al., 1975; Draper,
1991; Jordan et al., 2004). The volcanic rocks in the HLP are intercalated basalt lava
flows and rhyolite ash flow tuffs and related sediments with widely distributed rhyolite
dome complexes. Silicic volcanism in the HLP migrated in a northwesterly direction
starting in southeastern Oregon at ca. 10 Ma and ending near Newberry volcano at ca. 2
Ma where volcanism continues today (Jordan et al., 2004). Several periods of increased
basaltic volcanism in the HLP occurred at 7.8-7.5 Ma, 5.9-5.3 Ma and 3-2 Ma, however
the basaltic volcanism did not exhibit age progressive migration similar to the rhyolitic
volcanism in the HLP (Jordan et al., 2004). HLP basaltic lavas are dominantly olivine
and plagioclase phyric high-alumina olivine tholeiites, which often exhibit a diktytaxitic
texture (Hart et al., 1984; Draper, 1991; Jordan et al., 2004; Grove et al., Petrology of the
HLP, special volume). The abundant volcanic and sedimentary cover in the HLP make it
difficult to unravel the basement geology. Exposures of oceanic and arc terranes in the
Klamath Mountains to the southwest and Blue Mountains to the north, suggest the HLP is
built on accreted Mesozoic-Cenozoic basement composed of island-arc assemblages,
geosynclinals deposits and paleo-microplates assembled on the western margin of the
North American plate over time (e.g., Dickinson, 1979). The eastern margin of the HLP
164
is located at the Oregon-Idaho border, which roughly coincides with western margin of
Precambrian North America and an abrupt change in the initial
87Sr/ 86 Sr of Mesozoic
and
Cenozoic magmatic rocks from <0.706 west of the boundary to >0.706 to the east
(Armstrong et al., 1977; Manduca et al., 1992).
The Cascades arc is a subduction-related volcanic arc extending 1300 km N-S from
southern British Columbia to northern California. Today, the Cascades are characterized
as a "hot" subduction zone due to the subduction of the young (<28 Ma) Explorer, Juan
de Fuca and Gorda Plates and high heat flow within the central portion of the arc (e.g.,
Schmidt et al., 2008). Convergence of these plates is occurring at 40-45 mm/yr and
varies from orthogonal convergence in British Columbia to more oblique northwesterly
convergence in southern Oregon and California (Wilson, 1988), where we focus in this
study. Geophysical studies image the slab at ~100-200 km depth below this southern
extent of the Cascades (Zucca et al., 1986; Harris et al., 1991; Benz et al., 1992). In
northern California, subduction of the Mendocino triple junction suggests the presence of
a slab window below the southernmost Cascades (Beaudoin et al., 1996). The Tertiary
volcanic deposits exposed in the southern Cascades arc can be divided into a western
faulted Eocene-Miocene sequence, unconformably overlain by the High Cascades
Pliocene-Holocene sequence. In the southern Cascades, abundant diffuse mafic scoria
cones are erupted in close proximity to more proto-typical andesitic arc volcanic rocks
(Sherrod and Smith, 1990). Mt. Shasta is a andesite stratovolcano near the southern end
of the Cascades arc that grew in four eruptive stages between 250 ka and 2 ka along with
isolated flank eruptions and cinder cones (Christiansen et al., 1977). Work by Baker et
al. (1994) reveals Quaternary HAOTs erupted in proximity to Mt. Shasta represent
165
anhydrous low degree (6-10%) partial melts of a depleted mantle source that has been
slightly enriched by a fluid component derived from the subducted slab. Similar to the
HLP, the structure and composition of the southern Cascades basement can be difficult to
unravel but appears to be composed of thrust sheets of Paleozoic to Eocene accreted
oceanic and immature continental terrains, such as the Paleozoic ultramafic Trinity
peridotite, and the Paleozoic-Mesozoic marine metasedimenary and volcanic rocks
exposed in the Klamath Mountains (Griscom, 1980; Blakely et al., 1985; Zucca et al.,
1986; Snoke and Barnes, 2006).
Newberry Volcano in central Oregon and Medicine Lake Volcano in northern
California are both broad, shield-shaped volcanoes > 1 km in height with central calderas,
flanks dotted with multiple vents and extensive aprons of basaltic lava (Donnelly-Nolan,
2008). Both volcanoes are located just east of the Cascades arc axis at the intersection of
major tectonic features. Newberry is located at the westernmost terminus of the HLP
volcanic trend and the Brother's Fault zone, the northern bounding fault of the Basin and
Range (Lawrence, 1976), while Medicine Lake volcano is located at the northwestern
extension of the Walker Lane fault zone (Grose et al., 1989; Hildreth, 2007), the southern
extension of the Klamath Graben, and the volcanic highlands that extend WSW toward
Mt. Shasta (Donnelly-Nolan et al., 2008). Both volcanoes have also erupted primitive
basaltic through rhyolitic lavas during their ~half million year histories. Newbery
Volcano is the larger of the two volcanoes (~3000 km2 ) and has erupted >500 km3 of
bimodal basalt and high-silica rhyolite since -500 ka (Jensen et al., 2009), including at
least three significant ash-flow tuffs, the most recent at ~80 ka. Overall Newberry lava
compositions are unlike those from elsewhere in the HLP and primitive basalts at
166
Newberry are notably higher in K2 0 that those from Medicine Lake. Volcanism at
Medicine Lake (-2000 km 2) began at ~475 ka (Donnelly-Nolan and Lanphere, 2005;
Donnelly-Nolan et al., 2008) and the early eruptive history of the volcano produced
abundant rhyolitic to dacitic lava until 300 ka. The majority of the volcano was built by
eruptions of basalts between 100 and 13 ka that was followed by postglacial episodic
bimodal silicic and basaltic volcanism. Petrologic and isotopic work indicates arc fluids
are present in both volcanic systems, such that both erupted wet calc-alkaline basalts, as
well as dry HAOT's throughout their histories, often in temporal and spatial proximity to
one another. Studies of HAOT's erupted at Medicine Lake suggest they are anhydrous
partial melts of a MORB-like mantle lherzolite source enriched by a fluid component
derived from the subducted slab (Bartels et al., 1991).
The Modoc Plateau is a block-faulted Tertiary volcanic terrain in the northwest Basin
and Range (NWBR), located in the extreme northeast corner of California, east of the
Cascades arc and south of the HLP. Although the Basin and Range province (BRP) of
eastern, central and western Nevada has been well studied, comparatively little is known
about the diffuse northern margin of the province that extends into northeastern
California and southern Oregon. Limited data suggest that BRP-related extension
migrated northwestward from central Nevada since the late Oligocene (see summary by
Wernicke, 1992) and began after 12 Ma in the NWBR of northeast California and
southern Oregon (e.g., Colgan et al., 2004). The numerous models for the driving forces
for BRP extension largely fall into two categories: (1) partitioning of the relative motion
between the North American plate and the subducting Farallon plate (e.g., Humphreys,
1994; Atwater and Stock, 1998; Wesnousky, 2005) or (2) collapse of the Laramide-age
167
crust in the region (e.g., Glazner and Bartley, 1984; Humphreys and Coblentz, 2007;
West et al., 2009). In general, the Modoc Plateau is composed of flat-lying pyroclastic
rocks with intercalated and capping basalt flows that are believed to be Miocene and
younger in age (Duffield and Fournier, 1974). The Warner Range, located in the eastern
Modoc Plateau, is a N-S horst bounded by the Surprise Valley fault, the westernmost
large-offset fault of the BRP. The Warner Range exposes over 3000 m of rhyolitic to
basaltic pyroclastic deposits and lava flows erupted between 32 and 14 Ma (Duffield and
McKee, 1986; Fosdick et al., 2005).
3. Methods
3.1 Lava Sample Selection
Primitive basaltic lavas from across the northernmost Basin and Range and the central
- southern Cascades (Figure 1) were selected for thermometric and barometric
calculations, as these liquids are the least likely to have experienced significant
modification via fractional crystallization and crustal assimilation since their origin in the
asthenosphere. We consider lavas to be primitive if they are rich in Mg relative to Fe
(Mg# (Mg/(Mg+Fe ) >0.50) with Si0 2 <52 wt.% and relatively phenocryst poor. Most
of the samples are HAOT and they are typically aphyric, containing rare
microphenocrysts of olivine and plagioclase. Many samples, approach compositions of
primary magmas with Mg#s> 0.70 (Figure 2) and require very little correction to bring
their compositions into equilibrium with a mantle mineral assemblage. In addition, we
restricted our calculations to Pliocene or younger basalts, with the majority of the
168
samples erupted within the last 15,000 years, in order to constrain modem processes to
the extent possible. In some cases, more than one sample was selected from a given
volcanic vent or unit and therefore each sample does not necessarily represent a distinct
melting event. These criteria are met by lavas from across the HLP, including Jordan
Valley Volcanic Field (Hart, unpublished data), Diamond Crater (Till et al., 2011),
unnamed vents near Steens Mountain and the Blue Mountains (Carlson, unpublished
data), Newberry Volcano (Donnelly-Nolan, unpublished data) and the central and
southern Cascades arc, including South Sister (Schmidt et al., 2008), Crater Lake (Bacon,
1990; Bacon et al., 1997), Mt. Bailey (Barnes, 1992), Mt. Shasta (Baker et al., 1994;
Elkins-Tanton et al., 2001), Medicine Lake (Baker et al., 1991; Bacon et al., 1997;
Elkins-Tanton et al., 2001; Donnelly-Nolan et al., 2008; Donnelly-Nolan, 2010) and Mt.
Lassen (Bacon et al., 1997) volcanoes, as well as various unnamed basaltic vents in the
Modoc Plateau (Carlson, unpublished data). The basaltic lavas considered here are
HAOTs to calc-alkaline basalts. The trace element compositions of the basalts suggest
they represent melts of depleted mantle, in some cases with the addition of a subduction
component, which may consist of subduction-related fluids, melts and/or previously
metasomatized lithosphere.
3.2 Basalt Thermometry and Barometry
These data enable the examination of the pressure and temperature of origin for
young basaltic volcanism in northern California and Oregon. For the thermometric and
barometric calculations, the major element compositions of all samples were first
corrected for minor fractional crystallization following the methods described by Till et
al. (2011) until the bulk compositions could be approximated as a liquid in equilibrium
169
with a mantle olivine with a forsterite content (Mg 2 SiO4 ) of 90. The spinel lherzolite
melting model of Till et al. (2011) was then used calculate the pressure and temperature
where the basaltic melts were last multiply saturated with an upper mantle assemblage of
olivine, orthopyroxene, clinopyroxene, and plagioclase and/or spinel, also known as the
multiple saturation point (MSP). The olivine equation of Till et al. (2011) was
demonstrated to most closely reproduce the pressure of experiments and is therefore used
here for our pressure calculations. These basaltic melts likely represent batch melts as
demonstrated by Till et al. (2011) and therefore the calculated pressure represents the
shallowest depth of mantle equilibration, or melting, for a given melt. Density profiles
for the crust and upper mantle below the HLP (Cox and Keller, unpublished data) and
northern California (Zucca et al., 1986) were used to determine a pressure-depth
relationship at the basalt sample localities and convert to depth.
Experiments by Gaetani and Grove (1998) illustrate that silicate melts formed in the
presence of water are multiply saturated with a mantle assemblage at greater pressures
than analogous anhydrous melts. Specifically these experiments demonstrate the
multiple saturation point for a liquid in equilibrium with a spinel lherzolite mineral
assemblage shifts to greater pressure by 0.4 GPa with the addition of 6 wt.% H20. The
melting model of Till et al. (2011) was calibrated with only anhydrous plagioclase and
spinel lherzolite melting experiments, therefore the pressures of multiple saturation
calculated with the Till et al. (2011) barometer are inaccurate for hydrous melts. Olivineplagioclase hygrometry suggests basalts from Newberry Volcano with -1000 ppm Sr
contained -4 wt.% H20 prior to eruption (Barr, 2010). The Sr/H 20 ratio from the
Newberry basalts was therefore used to estimate the magnitude of the pressure correction
170
required for HLP and Cascades samples that likely involved H20 in their formation based
on their petrology (e.g., olivine and plagioclase compositions) and/or geochemistry (e.g.,
calc-alkaline basalts with >400 ppm Sr).
4. Pressure and Temperature of Mantle Melt Extraction for Primitive HLP and
Cascadian Basalts
The depths of melting for primitive southern Oregon and northern California lavas are
illustrated in two E-W transects: one transect below the HLP and Oregon Cascades at
-43.5'N that covers ~5' longitude and another below the California Cascades and Modoc
Plateau at -41.5'N that covers ~3' longitude (Figure 3). The depth of Quaternary melt
extraction along the HLP track overall shallows to the west, from ~45 km below
Diamond Crater and Jordan Valley Volcanic Field in the east, to ~30 km below
Newberry volcano and ~25 km below Crater Lake in the west. The minimum depth of
melting is -35 km for a suite of basalts from unnamed vents near Steens Mountain and
the Blue Mountains in eastern Oregon, approximately 10 km shallower than the
Quaternary basalts at Jordan Valley and Diamond Crater. Two primitive basalt samples
from the South Sisters area in the Cascades arc (Schmidt et al., 2008), which is found at
the far western end of the HLP track, originated at 45-53 km similar to the average depth
of basalts from Newberry (Donnelly-Nolan, unpublished data; Barr, 2010). The
maximum depth of melt extraction is located at -55 km for the northern transect and -50
km for the southern transect.
Our thermometry indicates that the basalts originated at temperatures between 12501380'C in the asthenosphere. Our temperature calculations are dependent on both
pressure and the compositional characteristics of the fractional crystallization-corrected
liquid, including its Mg# and wt.% K20. Therefore the temperatures shown in Figure 3
are an average for the liquids that originated at the specified depth. Phase equilibrium
studies of primitive basalts lavas from Medicine Lake (Baker et al., 1991), Mt. Shasta
(Baker et al., 1994), and Jordan Valley (Till et al., 2011) find that HAOT's from these
localities were separated from the mantle at -1 300'C and 10-11 kbar, a pressure range
equivalent to -35-39 km depth in this region, corroborating our calculations.
5. Constraints on Crustal and Lithospheric Structure
A number of geophysical studies have focused on the crustal structure and Moho
depths for the regions of interest in this study and reveal remarkably consistent results.
Most recently, Eager et al. (2011) analyzed teleseismic P-to-S receiver functions to image
the crustal structure below the HLP and determined Moho depths for the region using
both a H-K stacking and a new Gaussian-weighted common conversion point (GCCP)
stacking technique. The H-K and the GCCP stacking techniques overall yield similar
results for the topography of the Moho with ;> 40 km thick crust below the Cascades and
thinner 31-36 km crust below the HLP and northern Great Basin, where GCCP Moho
depths average -5 km deeper than those determined from H-K stacking (Figure 3). A
reverse seismic refraction profile 30 km southwest of Crater Lake volcanic field defined a
crustal thickness of 44 km (Leaver et al., 1984). Gashawbeza et al. (2008) also analyzed
172
teleseismic data using H-K stacking to determine Moho depths below the Modoc Plateau
and the areas to the east in the Great Basin. Two stations in the Modoc Plateau yield
Moho depths of 35 and 37 km, compared with H-K Moho depths of 36 -37 km
determined by Eager et al. (2011) for the same region. A seismic refraction survey
conducted by the USGS in 1981 characterized the crustal structure of the Klamath
Mountains, Cascade Range, and Basin and Range province of Northern California (Zucca
et al., 1986). This survey estimates Moho depths of 36-45 km below the southern
Cascades arc, specifically 33-37 km below Mt. Shasta, and 38-45 km below the Modoc
Plateau. In addition, a teleseismic tomography experiment focusing on Medicine Lake
Volcano (Ritter and Evans, 1997) inferred a Moho depth of ~36-37 km, similar to
estimates by Mooney and Weaver (1989) of 38-40 km below Mt. Shasta to Medicine
Lake and 38±4 km below Lassen volcano. These results are summarized in Figure 3. The
variable depth of the Moho below the HLP and Oregon Cascades relative to the constant
depth of the Moho below northern California likely reflects differences in the nature of
the Cenozoic basement to the north and the Paleozoic-Mesozoic basement to the south, as
well as variability in the amount of crustal extension experienced by both regions.
The shallowest depths of melting for the primitive basalt samples from all regions in
this study appear to hug the location of the Moho as determined from the various
geophysical studies. In the northern transect, primitive mantle melts from Crater Lake,
unnamed vents to the east, and Newberry volcano originated at such shallow depths
according to our calculations that they plot above some or all estimates of the present-day
geophysical Moho. A fundamental premise of this work is that the studied primitive
basalts originated in the asthenospheric mantle. Therefore our calculated depths can be
173
used to infer an alternative petrologic Moho. These depths also suggest that the
mechanical lithosphere cannot be significantly thicker than the continental crust here.
Several other studies of the thickness of the mechanical lithosphere in the overriding
plate at subduction zones draw a similar conclusion. The geodynamic models of
Rowland and Davies (1999) find that mechanical lithosphere thickness at subduction
zones is controlled by compositional buoyancy and therefore closely related to the
thickness of the crust. Plank and Langmuir (1988) also find a correlation between crustal
thickness and the major element parental magma at subduction zones that can be
explained if crustal thickness controls the temperature and depth of asthenospheric
melting, as observed here.
6. What Is Driving Asthenospheric Melting Below Southern Oregon and Northern
California?
The extremely low S-wave velocities observed in the shallow mantle below a
significant portion of the Pacific Northwest support our finding of anomalously warm
sub-lithospheric mantle in these regions (Wagner et al., 2010). The high temperatures
required to generate magmas at such shallow depths in this region are startling and raise
the obvious question, what is driving asthenospheric melting?
We find no evidence for hotspot or plume-related melting processes to generate the
<25 Ma basalts or the thin mantle lithosphere in the region of interest. Geochemically,
these basalt compositions do not resemble ocean island or flood basalts that are attributed
174
to hot spots (Figure 2; Carlson, 1984). The basalts studied here are relatively depleted in
the high field strength elements, such as U, Th, and Zr, and last equilibrated with the
mantle at shallow depths, inconsistent with an origin in a mantle plume (Figure 3).
Furthermore, the HLP HAOTs do not have the high 3He/ 4He values measured in Snake
River Plain basalts, which is often taken as the best geochemical indication of a plume
source (Graham et al., 2009). A number of recent geophysical studies also find little
evidence for a mantle plume beneath the HLP, where very low seismic velocities directly
below the crust are restricted to the upper 100 km of the mantle (e.g., Warren et al., 2008;
Wagner et al., 2010). Therefore plate motions likely provide the driving force for the
observed conditions in the mantle beneath southern Oregon and northern California.
In a study similar to ours, Elkins-Tanton et al. (2001) found the pressures of mantle
wedge melting decrease from east to west below Medicine Lake and Mt. Shasta and
inferred these pressures of melting paralleled mantle flow in the upwelling limb of corner
flow as calculated by Furukawa (1993). In this study we observe a similar pattern in the
depth of melting below Mt. Shasta and Medicine Lake volcanoes, but at overall shallower
depths and corresponding lower temperatures. Differences in our calculated depths and
temperatures of melting are the result of using the new lherzolite melting model of Till et
al. (2011), relative to the model of Kinzler and Grove (1992) used in their study. In our
northern transect, the depth of melting also appears to decrease from east to west. Other
studies of primitive basalts erupted above a subduction zone also find evidence for
anhydrous adiabatic decompression melting induced by corner flow in the mantle wedge
(Sisson and Bronto, 1998; Righter, 2000; Cameron et al., 2003). Geodynamic models
that include realistic temperature-dependent viscosities (Furukawa, 1993; Eberle et al.,
175
2002; Conder et al., 2002; Kelemen et al., 2003) produce the significant upwelling due to
corner flow in the mantle wedge required to generate these anhydrous basalts, as
discussed by Wiens et al. (2008). Furthermore, the close spatial and temporal association
of nominally anhydrous basaltic lavas to products of hydrous flux melting at South Sister,
Crater Lake, Newberry, Medicine Lake and Mt. Shasta volcanoes support the
interpretation that shared patterns of mantle flow are responsible for their production.
If melting below the HLP and southern Cascades is initiated by upwelling due to
corner flow in the mantle wedge, an important consideration is the oblique nature of the
North America-Juan de Fuca (NA-JdF) and North America-Gorda (NA-G) plate
convergence. A component of horizontal flow in the mantle wedge may therefore be
aligned with the SW-direction of NA plate motion, rather than solely perpendicular to the
NA-JdF and NA-G plate boundaries. Measurements of seismic anisotropy in the region
do in fact suggest an ESE-directed component of mantle flow that becomes more
pronounced north of the HLP (Long et al., 2009). The oblique nature of mantle flow
vectors relative to our E-W cross-sections could also explain why the minimum melting
depth across southern Oregon has such a shallow slope relative to the steep dip of mantle
flow lines in the top half of the wedge predicted by geodynamic models with
temperature-dependent viscosity (e.g., Condor et al., 2002; Kelemen et al., 2003).
The southern boundary of subducting slab is located just north of ~40'N latitude
based on the location of the Mendicino Triple Junction and tomographic images of the
slab (James et al., EPSL, in revision). Laboratory models of subduction illustrate a
toroidal flow around a slab edge, which includes a pronounced vertical component when
slab rollback is occurring, as in the northwestern U.S. (Funiciello et al., 2006; Druken
176
and Kincaid, in preparation). Therefore toroidal flow around the southern edge of the slab
could contribute to upwelling and decompression melting in the mantle below northern
California. The complex interplay of corner and toroidal flow at the southern edge of the
slab may also produce the narrow E-W surface expression volcanism in California
relative to a more proto-typical pattern of corner flow producing the wider E-W location
of volcanism in southern Oregon.
An alternative interpretation is that anhydrous melting in southern Oregon and
northern California is caused by extension-driven mantle upwelling. Recent geological
and geophysical studies suggest that the evolution of the NWBR was characterized by
voluminous volcanism without significant tectonism, followed by low-magnitude (520%)
extension along high-angle normal faults (Lerch et al., 2008) that began ca. 12 Ma (see
compilation in Scarberry et al., 2010). Extension also occurred within the main Cascades
arc during this time period, as indicated by discontinuous north-striking graben faults
(Hughes and Taylor, 1986; Smith et al. 1987). Paleomagmatic estimates at 42'N latitude
suggest that eastern Oregon has extended -17% since 15 Ma (Wells and Heller, 1988)
and Eager et al. (2011) use crustal thicknesses to estimate that the HLP has experienced
~16% extension since 10 Ma. The depths of melting we calculate for basalts from the
Modoc Plateau and unnamed vents near Steens Mountain in Oregon, are inconsistent
with the overall pattern of decreasing melting depths from east to west in both of our
transects. Because several of these basalts appear to originate at or above the present-day
geophysical Moho, they are good candidates for melting triggered by crustal extension.
Thus there is evidence for extension-driven asthenospheric upwelling in southern Oregon
and northern California, which is likely the result of partitioning of the relative motion
177
between the North American plate and the subducting JdF-G plates (e.g., Humphreys,
1994; Atwater and Stock, 1998; Wesnousky, 2005).
Therefore anhydrous melt generation below southern Oregon and northern California,
could be the result of, 1) upwelling due to corner flow, 2) toroidal flow around the
southern edge of the slab, and/or 3) extension-driven upwelling. Ultimately, all these
processes are driven by plate subduction and indicate the <10.5 Ma primitive basalts
erupted in the region are the result of purely tectonic processes. As such, it is probable
these three processes operate in unison or vary in importance through time. In many
geodynamic models of mantle flow at subduction zones, it is extremely difficult to
produce the high temperatures required at 30-50 km depth required to generate
asthenospheric melts, as constrained in this and other petrologic studies (see review in
Kelemen et al., 2003). However, Kelemen et al. (2003) show that if the thermal
boundary layer at the base of arc crust (i.e., the mechanical lithosphere) is thinner than
originally thought, mantle models incorporating temperature-dependent viscosity and
widely accepted values for activation energy and asthenospheric viscosity produce
temperature-depth relationships consistent with the petrologic estimates, including those
from this study. Therefore, no matter which process triggers mantle melting, the thin
nature of the mechanical lithosphere in this region also plays an important role in
generating the conditions we observe in the mantle below southern Oregon and northern
California, which may be more prevalent at convergent margins than previously thought.
178
Acknowledgements
The ideas and geochemical data presented in this paper are an outgrowth of the
Continental Dynamics Project on Oregon's High Lava Plains and earlier foundational
research by a number of the authors. A great deal of thanks is extended to Rick Carlson,
Julie Donnelly-Nolan, Matt Fouch, Anita Grunder, Randy Keller, David James, Chris
Kincaid, Bob Duncan, Jenda Johnson and all the other scientists and personalities that
brought this project to fruition.
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186
Figure Captions
Figure 1. Geologic and tectonic map of the Pacific Northwest, United States. (a)
Locations of Holocene volcanism indicated by the white triangles with locations of
interest in this study labeled in bold italic: Crater Lake Volcanic Center (CL), Diamond
Crater (DC), Jordan Valley Volcanic Center (JV), Lassen Volcanic Field (LS), Medicine
Lake Volcano (MV), Mount Bailey (MB), Mount Shasta (SH), and South Sister Volcano
(SS). The location of Quaternary basaltic volcanism is shown by the dark gray shaded
regions and the mid-Miocene Steens and Columbia River flood basalts by the white
shaded area. The white dashed lines represent depth contours to the top of the slab (km)
(Harris et al., 1991) and the black dashed lines the approximate location of the
87Sr/ 8 6
Sr
0.704 and 0.706 lines (Armstrong et al., 1977; Manduca et al., 1992). Isochrons (Ma) for
the migration of rhyolitic volcanism in the High Lava Plains (Jordan et al., 2004) and
Snake River Plain (Christiansen et al., 2002) denoted by the black dotted lines. Gray
bold lines A-A' and B-B indicate the locations of the cross-sections in Figure 3. (b)
Location of primitive basalt samples used in this study.
Figure 2. Chemical characteristics of <10.5 Ma primitive basaltic lavas from southern
Oregon and northern California selected for this study.
Figure 3. Depth of asthenospheric melt segregation calculated for primitive basaltic lavas
from southern Oregon and northern California. White symbols denote basalt samples
corrected for the effect of H20. The base of the green field illustrates the depth of the
lithosphere-asthenosphere boundary as predicted from our calculations and the dashed
black lines indicate geophysical estimates of the depth of the Moho in the area and the
corresponding reference number. Location of the cross-sections as shown in Figure 1.
Figure 4. Cartoon illustrating the three types of asthenospheric flow that could be
responsible for the generation of <10.5 Ma primitive basalts in southern Oregon and
northern California: (I) ESE-directed corner flow in the mantle wedge, (II) toroidal flow
around the southern termination of the subducting slab, and (III) additional upwelling
187
triggered by Basin and Range crustal-extension. Gray dashed lines represent depth
counters for the subducting slab (Harris et al., 1991) and black dashed lines are the
approximate location of the
87Sr/ 86Sr 0.704
and 0.706 lines (Armstrong et al., 1977;
Manduca et al., 1992).
Table 1. Primitive Basalt Samples used in this study. Samples pressures in italics were
corrected for the effect of H20 as described in the text. Sample pressures were converted
to depth using crustal density profiles from regional seismic refraction experiments
(Zucca et al., 1986; C. Cox, personal communication, April 2011).
188
0
48* '0"N
0
-46 0'0"N
-44" 0'0"N
-42*0 '0"N
-40*0'0"N
48*0'0"N
4600'0"N.
440*0"N
42*0'0"N
40*0'0"N
Figure 1.
189
0.75
b.
0.70 -
4
m~
0.65 -
0.60
0.55
5e a
0.50
46
48
47
50
49
51
52
53
47
SiO2 (Wt%)
48
49
50
. . . . . . . . . . .
51
52
53
Si02 (wt%)
300
C.
250
4
U
200
z 150
EU
100
U
50
200
400
600
800
1000
1200
1400
1600
Rb Ba Th U Nb Ta K La Ce Pb Pr Sr P Nd Zr Hf Sm Eu Ti GdDy Y Ho Er Yb
Sr(ppm)
0
KEY
Jordan Valley Volcanic Center
+ Diamond Crater
A Newberry Volcano
* Unnamed Vents, Oregon
* South Sister Volcano
* Crater Lake Volcanic Center
Mt. Bailey Volcano
* Mt. Shasta & Medicine Lake Volcanoes
* Lassen Volcanic Field
l Modoc Plateau
* Av. MORB
Figure 2.
190
10
jjc
Newberry
Cascodes
1200
20
1250
E30
1300
S40
I
500
4..'
1350
50
1400
60
to
10
120*W
121OW
1220W
e
Mt. Sdoc
118"W
119"W
\oteu
Lo
1200
20
1250
30
Inferred LA8
S1300
Q
40
'
-
117"W
Primitive Basalt Samples
0 Jordan Valley Volcanic Center (PCorrected)
* Diamond Crater
A Newberry Volcano
A Newberry Volcano (PCorrected)
U Unnamed Vents, Oregon
2 South Sister Volcano
south Sister Volcano (PCorrected)
* Crater Lake Volcanic Center
0 Crater Lake Volcanic Center (PCorrected)
- Mt. Bailey Volcano
0 Mt. Shasta & Medicine Lake Volcanoes
* Lassen Volcanic Field
CL Lassen Volcanic Field (PCorrected)
Modoc Plateau
Moho: Geophysical References
1350
50
1a. Eager et al. (2011) - H-k (errors ±2-5 km)
1b. Eager et al. (2011) - GCCP
2. Leaver et al. (1984) - seismic refraction
60
122*W
121*W
120OW
1400
119*W
Figure 3.
191
3. Zucca et al. (1986) - seismic refraction
4. Ritter & Evans (1997) -tomography
5.Mooney & Weaver (1989)
6.Gashawbeza et al. (2008) - H-k
Figure 4.
192
Table 1. Samples used for Pressure & Temperature Calculations
Sample ID P (GPa) Depth (km Temp (*C) Ref ISample ID P (GPa) Depth (km) Temp (*C) Ref Sample ID P (GPa)
Depth (km) Temp ('C) Ref
Jordan Valley Volcanic Center, OR
Jordan Valley (Cont.)
Mt. Bailey, OR
753
1.47
49.3
1349
1
n1403
1.51
50.5
1362
1
64
1.37
46.9
1342
755
1.41
1342
1
47.3
n1503
1.45
48.6
1351
1
98
1.36
46.6
1340
759
1.55
51.6
1363
1
1.45
48.5
1346
1 Crater Lake, OR
jc30b
8528b
1.63
54.1
1373
1
n1703
1.34
45.1
1334
1
81C621
0.92
32.4
1287
h818
1.59
52.8
1368
1
1.48
owy22
49.5
1356
1
82C894
0.81
1271
28.7
h824
1.59
52.8
1366
1
1.41
47.3
1343
1 88C1143
owy23
0.66
23.5
1250
h858
1.54
51.4
1362
1
1.45
1347
1 88C1530
48.6
owy3l
1.57
53
1361
h860a
1.47
49.2
1349
1
44.1
1.31
owy36
1331
1 88C1540
1.38
47.3
1342
h861
1.63
54.2
1376
1 Newberry Volcano, OR
Mt. Shasta- Medicine Lake, CA
h862
1.63
54.1
1375
1 RC05-2
1.46
49.7
1350
2
83-43
0.93
32.6
1282
h863
1.59
52.8
1370
1 RC05-3
1.46
49.6
1349
2
82-72f
1.12
38.7
1311
h864
1.68
55.6
1380
1 RC05-10
1.27
43.7
1329
2
840M
1.22
42
1321
h865
1.66
55.0
1379
1 RC05-11
1.39
47.3
1341
2
1788M
1.04
36.3
1295
h866
1.61
1372
1 RC05-13
53.5
1.39
1329
2
47.5
1629
1.23
1225
42.5
h943a
1.67
55.2
1379
1 RC06-1
1.45
49.3
1336
2
1630
1.36
46.5
1340
h943b
1.66
54.9
1378
1 RC06-3
1.41
48.2
1344
2 Lassen Volcanic Field, CA
h971b
1.72
56.8
1386
1
1.51
51.1
1682j
1356
3 LC86-1046
1.41
48.1
1347
h8610
1.33
44.8
1333
1
1.11
38.7
l j
1309
3 LC82-970
1.45
49.3
1352
h8612
1.62
53.9
1375
1
1.45
49.4
15j
1330
3 LC88-1311
1.3
44.6
1333
jc3
1.33
44.9
1333
1
1.15
40.0
1565j
1315
3 LC86-1005
1.13
39.4
1306
jc5
1.36
45.8
1338
1
1564j
1.16
40.0
1315
3 LM87-1384
1.73
53
1378
jc6
1.64
54.4
1373
1
1.10
38.3
1568j
1310
3 LC88-1312
1.14
35
1310
jc15
1.65
54.6
1378
1 RC05-4
0.93
32.8
1278
2 LC82-905
1.45
45.8
1343
jc30a
1.45
48.6
1347
1 RC05-9
1.50
50.7
1346
2 LC86-855
47.7
1.52
1353
jc3l
1.43
47.9
1345
1 RC06-2
1.05
36.7
1290
2 LC85-671
1.49
46.5
1345
jc32
1.41
47.2
1342
1 RC06-4
1.59
53.6
1358
2 Modoc Plateau, CA
jc33
1.36
45.8
1337
1 RC06-5
1.58
53.2
1357
2
a8946a
1.23
42.2
1326
jc34
1.36
45.8
1337
1
1.34
45.8
1327
222j
3
a8946b
1.05
36.3
1299
jc35
1.70
56.1
1382
1
1.27
43.6
1060j
1319
3
a8946c
1.29
43.9
1333
jc36a
1.58
52.6
1370
1
1.28
43.9
1319
761j
3
a8946d
1.09
37.7
1303
jc36b
1.52
50.9
1363
1
1.02
35.7
700j
1289
3
ch8311
1.02
35.7
1297
jc36c
1.63
54.2
1
1377
1023
0.96
33.8
1284
3
a8948b
1.08
37.6
1309
jc37
1.65
54.8
1380
1
1.02
1014j
35.6
1291
3
a8948c
1.13
39.1
1310
jc38
1.36
45.7
1337
1
1.02
868j
35.6
1290
3 ch8779a
1.14
39.5
1313
jc39
1.63
54.1
1374
1
1.04
990j
36.4
1290
3
a8932
1.32
44.9
1335
jc4O
1.62
54.0
1376
1
1.03
989
36.0
1288
3
a8927
1.35
45.8
1340
jc4la
1.61
53.6
1371
1
1017
0.89
31.2
1274
3 Diamond Crater, OR
jc4lb
1.69
55.9
1381
1
1.64
55.0
1576j
1360
3
G 07-12
1.50
50.4
1357
jc42
1.56
51.9
1364
1
1.12
38.9
710j
1299
3
G 07-13
1.39
47.2
1345
jc43a
1.37
46.2
1338
1
1.06
712j
37.1
1292
3
G 07-14
1.32
44.8
1336
jc43b
1.37
46.2
1339
1
968j
1.20
41.4
1313
3
G 07-15
1.39
47.2
1346
3
jv991
1.38
46.3
1339
1
75 j
1.36
46.5
1330
3
G 07-16
1.60
53.5
1370
jv992
1.35
45.4
1335
1
1.01
35.3
146j
1286
3 Unnamed Vents, OR
n104
1.49
50.0
1355
1
1.68
751j
56.4
1370
3
ch8343
1.29
43.9
1330
n603
1.55
51.8
1363
1
1.18
988j
40.9
1307
3
ch8342
1.39
47.1
1343
n703
1.29
43.7
1329
1 Three Sisters, OR
ch8351
1.08
37.2
1305
n1103
1.59
52.8
1367
1
CAB
1.33
45.6
1335
4
ch8353
1.02
35.1
1298
n1303
1.31
1316
44.2
1
LKT
1.56
52.9
1356
4
ch8357b
1.22
41.7
1323
ch8358
0.99
34.4
1295
1 Hart, Unpublished Data
2 Carlson, Unpublished Data
3 Donnelly-Nolan, Unpublished Data
4 Schmidt et al., 2008
5 Barnes
6 Bacon et al., 1997
7 Elkins-Tanton et al., 2001
8 Till et al., Submitted
italics corrected for the effect of H20
193
,
.29
Chapter 5:
A Mechanism for Low Extent Melts at the
Lithosphere-Asthenosphere Boundary
194
Abstract
Recent studies have imaged sharp vertical drops in shear wave velocity at the lithosphereasthenosphere boundary (LAB). In some regions, the magnitude of the negative velocity
gradient at the LAB is too large to be explained by changes in temperature alone. This
study demonstrates that small amounts of partial melt in the shallow asthenosphere are a
viable model for this sharp seismic boundary. In particular, we examine melting in the
upper asthenosphere at the edge of thick cratonic lithosphere, using the example of
eastern North America where a sharp LAB velocity gradient has been observed. Finite
element modeling of asthenospheric flow at an abrupt lateral decrease in lithosphere
thickness indicates that this geometry, together with lateral plate motions, produces edgedriven convection and asthenospheric upwelling at the continental margin. A key
component of this work is a comparison of the locations and extents of melting produced
by using different models for the depression of the peridotite solidus with varying H2 0
content. In addition, we develop a simplified parameterization of the H20-undersaturated
peridotite solidus for a constant degree of H20 saturation in nominally anhydrous
minerals. The patterns of mantle flow produced by our numerical modeling and various
solidus parameterizations predict less than 0.1 wt% to 2.8 wt% (0.0 1-3.3 vol%) melting
at depths between 102 and 126 km for an asthenosphere with a mantle potential
temperature of 1350'C and 150 ppm H2 0, or between 91 km and a maximum of 200 km
for a mantle at 1350'C and 450 ppm H2 0. If the asthenosphere has a mantle potential
temperature <1340'C or contains less than 150 ppm H2 0 at 3 GPa, no melting will occur.
This process of generating melt in the asthenosphere to produce a sharp vertical velocity
gradient at the LAB is viable in other locations where convective upwelling occurs in the
195
shallow asthenosphere although it is dependent on asthenospheric potential temperature,
composition, and H20 content. Because the asthenosphere may be heterogeneous in
composition and H20 content, the onset of melting below the LAB may fluctuate with
time and space, as may the magnitude of the shear velocity drop at the LAB.
1. Introduction
In the most general sense, the lithosphere-asthenosphere boundary (LAB) is
defined as the boundary between the rigid material that composes the Earth's tectonic
plates (i.e., the lithosphere) and the underlying material which undergoes solid-state creep
such that it exhibits fluid-like behavior on geologic times scales (i.e., the asthenosphere).
Yet the first-order physical and chemical properties of the mantle, which change at the
LAB to produce this rheological boundary, are poorly understood.
The LAB is marked by a negative gradient is seismic velocity from the fast
lithosphere to the slow asthenosphere.
Slow velocities in the asthenosphere have been
explained with a variety of models, some of which employ temperature gradients with or
without grain size variations [e.g., Faul and Jackson, 2005; Stixrude and LithgowBertelloni, 2005; Priestley and McKenzie, 2006], while others invoke a contrast between
a dry depleted lithosphere and a hydrated fertile asthenosphere [e.g. Hirth and Kohlstedt,
1996; Karato andJung, 1998; Gaherty et al., 1999] or the presence of melt within the
asthenosphere [e.g. Anderson, 1989; Kawakatsu et al., 2009].
The sharpness of the velocity gradient at the LAB in many regions poses a
challenge for purely thermal models [Fischeret al., 2010]. In the last decade, resolution
of the LAB has been improved with a variety of seismic data, including teleseismic
phases such as Ps and Sp [Farraand Vinnik, 2000; Levin and Park, 2000; Li et al., 2000;
196
Collins et al., 2002; Oreshin et al., 2002; Li et al., 2004; Rychert et al., 2005; Vinnik et
al., 2005; Kumar et al., 2005a; Kumar et al., 2005b; Chen et al., 2006; Kumar et al.,
2006; Mohsen et al., 2006; Sodoudi et al., 2006; Hansen et al., 2007; Heit et al., 2007;
Kumar et al., 2007; Li et al., 2007; Rychert et al., 2007; Wittlinger and Farra,2007;
Chen et al., 2008; Snyder, 2008; Hansen et al., 2009; Rychert and Shearer,2009; Abt et
al., 2010; Fordet al., 2010], ScS reverberations [Bagley and Revenaugh, 2008] and
multiple S surface reflections [Tan and Helmberger, 2007] (reviews may be found in
Rychert et al. [2010] and Fischer et al. [2010]). In some cases, the LAB velocity gradient
has been explicitly modeled. Rychert et al. [2005; 2007] utilized inversions of P-to-S
(Ps) and S-to-P (Sp) receiver functions to resolve a 5-10% decrease in shear wave
velocity over less than 11 vertical km at depths near 100 km beneath the northeastern
U.S. and southeastern Canada. Using experimentally-derived relationships between
temperature and shear velocity [FaulandJackson, 2005] and a minimum olivine grain
size of 1 mm [Evans et al., 2001], to produce this velocity gradient with temperature
alone, a temperature gradient of more than 20'C/km is required at the LAB [Rychert et
al., 2007]. In contrast, temperature gradients are less than 10 0 C/km in numerical models
for comparable tectonic settings (the boundary of thick cratonic and thinner continental
lithosphere) in which viscosity varies with temperature and pressure but is not affected by
other factors such as composition [King andRitsema, 2000; Cooper et al., 2004].
Moreover, in these and other models for this type of tectonic setting [Lenardic and
Moresi, 1999; Korenaga and Jordan,2002], thermal gradients from the lithosphere to the
asthenosphere are distributed over 60-80 km or more, as opposed to the much smaller
depth range spanned by the gradients in seismic velocity. Rychert et al. [2005; 2007;
197
2009] therefore conclude that a mechanism that involves more than temperature is
required to explain the observed velocity gradients: a more hydrated asthenosphere,
perhaps in combination with a downward decrease in chemical depletion or grain size or
with a coincident change in anisotropy, or simply a small amount of partial melt in the
asthenosphere.
Evidence for sharp LAB velocity gradients has also been found in other regions.
Using Ps phases in eastern China, Chen et al. [2006] inferred shear velocity decreases of
3-7% over 10 km or less and reached conclusions similar to those of Rychert et al. [2005;
2007]. Kawakatsu et al. [2009] obtained velocity drops of 7-8% and argued for horizontal
melt layers in the asthenosphere beneath the Pacific and Philippine Sea plates. Using Sp
phase alone, Li et al. [2007], Abt et al. [2010], and Fordet al. [2010] were able to limit
LAB velocity gradients to depth ranges of less than 30-40 km beneath the western U.S.,
large portions of the non-cratonic eastern U.S., and eastern Australia. Although
constraints on vertical velocity gradients from Sp phases alone are looser than those
possible with Ps data, 30-40 km is still much smaller than the depth range of lithosphereasthenosphere thermal gradients in numerical models [Lenardic and Moresi, 1999; King
and Ritsema, 2000; Korenaga andJordan, 2002; Cooper et al., 2004]. Thus these results
also point to compositional differences between the lithosphere and asthenosphere or the
presence of partial melt in the asthenosphere.
To further explore the origins of the LAB, this study examines the case of the
LAB below eastern North America. In this region, surface wave tomography studies
indicate that the lithosphere has a thickness of approximately 200 km below the
westernmost Appalachians and beneath the craton to the west [van der Lee, 2002; Li et
198
al., 2003; Nettles and Dziewonski, 2008; Bedle and van der Lee, 2009; Yuan and
Romanowicz, 2010]. Moving east, the lithosphere abruptly thins to thicknesses of 90-110
km beneath New England, based on Sp and Ps analyses [Rychert et al., 2005; Rychert et
al., 2007; Abt et al., 2010] and most surface wave tomography [van der Lee, 2002; Li et
al., 2003; Nettles and Dziewonski, 2008; Bedle and van der Lee, 2009] (in one exception,
Yuan and Romanowicz [2010] argue for a deeper LAB beneath New England based on
vertical variations in azimuthal anisotropy). An abrupt eastward decrease in lithospheric
thickness together with a WSW absolute motion for the North American plate [Gripp and
Gordon, 2002] is ideal to produce edge-driven convection. The work of King and
Anderson [1995; 1998] originally modeled small-scale convection in an isothermal
mantle resulting from discontinuities in the thickness of the lithosphere (i.e., edge-driven
convection). Newer models of edge-driven thermal convection suggest it could be a local
mechanism for melt formation or an alternative explanation to mantle plumes for
intraplate volcanism [King andRitsema, 2000; Korenagaand Jordan, 2002].
We have conducted numerical experiments using an abruptly thinning lithosphere
to test whether edge-driven convection can produce a small melt fraction at depths below
the seismically-imaged LAB [Rychert et al., 2005; Rychert et al., 2007; A bt et al., 2010].
We examine the range of compositions and H20 contents required to produce melting,
given the range of calculated mantle flow patterns, by comparing a variety of solidi
relevant to the asthenosphere, including the anhydrous peridotite solidus and several
parameterizations for peridotite solidus depression with varying H20 content. These
models for the melting behavior of peridotite in the presence of H2 0 express solidus
depression as a function of the concentration of water in the melt based on the bulk
199
partition coefficient for H2 0 between solid peridotite and melt. In addition, we derive a
simplified parameterization for solidus depression at a constant degree of H20-saturation
in nominally anhydrous minerals, which is only a function of pressure. Our
parameterization is based on the method of Hirth and Kohlstedt [1996] and includes new
experimental data on the H20-saturated solidus and on H20 partitioning between
nominally anhydrous peridotite minerals and silicate melts. The parameterization yields
comparable melting temperatures for a given H20 content, between 2 and 7 GPa, to preexisting H20-undersaturated melting models, including those of Hirschmannet al.
[2009], Katz et al. [2003] and a model for olivine liquidus depression by Medard and
Grove [2008]. The similarity of these expressions for solidus depression is not surprising
given they are all derived from or calibrated with high-pressure melting experiments on
peridotite or primitive basalts. We also explore the implications of the similarity of these
models for solidus depression in the presence of H20 and their implications for the
presence of small amounts of partial melt in the asthenosphere in generalized tectonic
settings.
2. Methods
2.1 Solidi Used in Melting Calculations
For our modeling we use the anhydrous peridotite solidus parameterization of
Hirschmann [2000], which is based on an up-to-date compilation of the experimental
data. The experiments used to bracket the solidus are depicted in Figure 1, along with
additional experiments at 0.9 - 1.6 GPa by Kinzler and Grove [1 992b] and at 1 atm by
Grove andJuster [1989]. As noted by Hirschmann [2000], the bulk composition of a
200
peridotite (specifically the alkali content, Mg# and cpx mode) plays an important role in
determining the solidus temperature. Therefore we use the Hirschmann[2000] equation
that has been filtered to exclude enriched peridotites and includes the experiments
conducted prior to 1988, which is as follows,
T(0 C)
=
-5.104P 2 -132.899P + 1120.611
(1)
where P is in GPa. We choose to include experiments conducted prior to 1988 because of
the lack of more recent experiments to constrain the solidus at <1 GPa. The above
equation yields a 1 atm solidus temperature of 1120.661 'C, whereas the parameterization
filtered for enriched and depleted peridotites, using only data that post-dates 1988, yields
a 1 atm solidus temperature of 11070 C. Using the equation of Grove and Juster [1989]
for the variation in the 1 atm anhydrous perdotite solidus temperature with composition,
typical glass compositions from MORB melting experiments at 0.9 GPa by Kinzler and
Grove [1992b] predict 1 atm solidus temperatures of 1150 - 1177'C; H20-saturated
peridotite melting experiments on a Hart andZindler [1986] primitive mantle
composition by Grove et al. [2006] at 1.2 GPa and Till et al. [2007] at 3.2 GPa predict 1
atm solidus temperatures of 11 96'C and 1205'C, respectively. Therefore, the
Hirschmann [2000] parameterization with an intercept of 1120.61 PC most closely
matches the constraints from 1 atm and other low pressure melting experiments. An
updated version of the Hirschmann [2000] parameterization for the pressures beyond that
required in our models (>10 GPa) can be found in Hirschmann et al. [2009].
H2 0-saturated peridotite melting experiments by Grove et al. [2006] from 1.2 to
3.2 GPa and Till et al. [2007] from 3.2 to 6 GPa are illustrated on Figure 1 and are the
source of the H 20-saturated solidus used for our parameterization. All regressions for the
201
following equations were calculated using the method of least squares. Peridotite solidi
used here represent the temperature of the first melt at a given pressure in a typical
undepleted upper mantle composition. In general, the solidi are suitable for a typical
MORB-source upper mantle composed of approximately 60% olivine, 15%
clinopyroxene, 15% orthopyroxene and 0-10% garnet (depending on pressure). A fourthorder polynomial fit most closely matches the H20-saturated experimental data
(R2=0.9816), with pressure in units of GPa. This solidus is applicable over the range of
experimental data it is derived from, which is 0 to 6 GPa.
T(OC)= -0.0842P 4 -1.7404P
3
+ 36.777P 2 -191.69P +1120.7
(2)
Although experimental liquids from experiments by Grove et al. [2006] at 1.2 GPa and
Till et al. [2007] at 3.2 GPa predict I atm solidus temperatures of-1200'C, we prefer to
use the 1120.7'C value from Hirschinann[2000], as it is infeasible that the anhydrous
and H20-saturated solidus cross at low pressures. For extrapolation to pressures between
6-12 GPa, a second order polynomial better describes the predicted (i.e., not tested
experimentally) solidus behavior relative to the fourth order polynomial but does not fit
the data as accurately at pressures less than 6 GPa (R2=0.9647). This quadratic is used to
determine the nominally anhydrous mineral surface in three dimensions as discussed
below.
T(0 C)
=
16.777P 2 - 149.5P + 1120.7
202
(3)
Partial melting experiments from 3-7.5 GPa by Kogiso et al. [2003] are the basis
for the garnet pyroxenite solidus used in our calculations (Figure 1). In our melting
calculations this parameterization is extrapolated to surface conditions.
T(*C)
=
-3.4524P 2 + 120.95P + 1096.8
(4)
H20-undersaturated peridotite melting is not as well understood experimentally as
the anhydrous and H20-saturated cases. To calculate H20-undersaturated solidi it is
necessary to know the H2 0 storage capacity of nominally anhydrous mantle minerals.
Here we define the H20 storage capacity as the maximum abundance of H2 0 that can be
accommodated in nominally anhydrous mantle minerals (we assume that all of the H20
resides in olivine, clinopyroxene, orthopyroxene and garnet), before the stabilization of
hydrous minerals, and use the H20-saturated peridotite solidus as a reference P-T path
for the upper bound on the storage capacity of nominally anhydrous minerals, after the
studies of Hauri et al. [2006] and Hirschmann et al. [2005]. Techniques to measure the
partition coefficients for H20 in nominally anhydrous minerals have largely developed
within the last 15 years [e.g., Bell and Rossman, 1992; Hirth and Kohlstedt, 1996; Koga
et al., 2003; Hirschmann et al., 2005]. Here we use the pressure-dependent partition
coefficients and mineral modes of upper mantle minerals near the LAB presented by
Hirschmann et al. [2009] and references within.
These partition coefficients can then be used to calculate the maximum
concentration of H20 dissolved in a peridotite melt at incipient melting using the
expression for batch melting at zero melt fraction,
cSolidus
CO
'HO
HO
0peridlliq
DH
203
(5)
where C'
0
is the bulk concentration of H20 in the peridotite, D erid Ii, is the bulk
partition coefficient for H between peridotite solid and melt and Col"us is the
concentration of H2 0 dissolved in the first infinitesimal fraction of melt. Peridotite
solidus depression was then calculated using both the cryoscopic approach benchmarked
against hydrous melting experiments of Hirschmannet al. [2009] (illustrated in their
Figure 10), and the mathematical formulation of Katz et al. [2003] based on constraints
from thermodynamic modeling and high-pressure experiments (see their equation 16).
Solidus depression follows a power-law relationship with CqO in both of these
expressions due to the greater depolymerizing effect and larger quantity of hydroxyl ions
in silicate melt compositions at low water contents (1-2 wt.% H2 0), relative to molecular
water [Stolper, 1982; 1989; Inhinger et al., 1999]. Therefore, both equations account for
the larger effect of Cliq on solidus depression at low total water contents and the
decreasing proportion and effect of C$7O on solidus temperature at higher total water
contents. The solidi for a bulk peridotite with C', 0 of 150 ppm and 450 ppm H20,
shown in Figure 2, were calculated by subtracting the solidus depression (AT) calculated
from Hirschmannet al. [2009] and Katz et al. [2003] from the anhydrous solidus (given
in (1)).
Medardand Grove [2008] conducted experiments to determine the effect of H2 0
on the liquidus temperature of olivine-saturated primitive melts. They conclude that
liquidus depression for primitive basaltic melts can be approximated as a linear function
of CH, up to 1.3 wt % H2 0 and as a third-order polynomial up to 12 wt. % H20.
Furthermore, for more silica-rich melt compositions, including low-degree mantle melts,
204
they find that the effect of H20 is smaller than for basaltic melt, and can be better
approximated as a linear function. For comparison to the previous two expressions for
solidus depression, AT for bulk peridotite melting was calculated using both the linear
and third-order expressions of Medardand Grove [2008] for liquidus depression. The
third-order polynomial produces AT within ± 9'C of the expressions for solidus
depression of Katz et al. [2003] and Hirschmann et al. [2009] at 150 ppm H2 0 and the
linear approximation yields AT within ± 3 'C at low CO (< 2 wt. %) and ±1 10 C at
higher CI 0 .
For comparison to the solidus depression at constant bulk H20 contents, we have
developed a simplified parameterization for solidus depression at a constant degree of
H20-saturation in nominally anhydrous minerals, which is only a function of pressure.
Our parameterization is based on the method of Hirth and Kohlstedt [1996] and includes
new experimental data on the H20-saturated solidus. Mantle melting in the presence of
H20 can be described using a phase diagram for a constant pressure such as that shown in
Figure 3 after Silver and Stolper [1985] and Grove et al. [2006], with bulk H20 content
increasing on the x-axis and temperature increasing on the y-axis. The green curve
illustrates the maximum storage capacity of nominally anhydrous peridotite minerals
based on data from Hauri et al. [2006]. The blue curve illustrates the simplified melting
boundary (i.e., the liquidus) for peridotite, extending from a eutectic-like H20-saturated
solidus with the maximum H20 content of the melt based on H20 solubility in silicate
melt [Grove et al., 2006] to an anhydrous liquidus [Zhang and Herzberg, 1994].
We calculate the H20-undersaturated solidi for our melting calculations by fitting
a line to the nominally anhydrous curve between the anhydrous solidus and the H 2 0-
205
saturated solidus temperatures, for a given pressure, these equations are given in Table 1.
We approximate the nominally anhydrous curve as a line based on an equivalent
formulation for the liquidus in the thermodynamic model for hydrous silicate melts
developed by Silver and Stolper [1985] for diopside-H 20 and albite-H 20, as well as the
similarity of the linear expression for liquidus depression with H20 of Medard and Grove
[2008] to the power-law expressions for solidus depression of Katz et al. [2003] and
Hirschmannet al. [2009]. The model of Silver and Stolper [1985] is derived from the
assumption of ideal mixing of hydroxyl groups, H 20 molecules and oxygen in the melt,
which is an adequate first-order approximation at low H20 contents but not likely
adequate for higher-order models (see discussion of thermodynamic models of silicate
melts in Medard and Grove [2008]). Following the model of Silver and Stolper [1985],
as well as the work of Hirth and Kohlstedt [1996] and Hauri et al. [2006], we assume an
activity of unity for H20 in the fluid. Therefore at the saturation of the nominally
anhydrous minerals, the peridotite has a water-activity (aH2o) equal to unity and melts at
the H20-saturated solidus temperature. When a peridotite contains no H2 0, it has an aH20
equal to zero and melts at the anhydrous solidus temperature. And when the peridotite
contains an amount of water less than that required for mineral saturation, the melting
temperature scales linearly with water-content between O<aH20< 1. To assume an activity
of unity for water in the fluid is a simplifying assumption, as the amount of dissolved
silicate in hydrous fluids increases with increasing pressure [e.g., Stalder et al., 2001;
Hermann et al., 2006], causing the activity of water to constantly decrease with
increasing pressure (i.e., max aH2O<1). Therefore our estimates of the temperature of the
H20-undersaturated solidus at the higher end of the range of pressures examined (-3-6
206
GPa) are maximum estimates, and in fact mantle peridotite may reach water-saturation at
even lower temperatures for a given pressure. A simple H 20-undersaturated melting
scenario for a peridotite with 150 ppm H20 (0.015 wt%) is illustrated by the red points in
Figure 3. This H20-undersaturated peridotite will melt at a temperature of 14090 C,
where the H20 content intersects the nominally anhydrous boundary. A peridotite
containing >1500 ppm H20 (0.15 wt%) at 3 GPa will melt at the temperature of the H2 0saturated solidus (825'C) and a completely anhydrous peridotite will melt at the
anhydrous solidus (1473 0 C).
Here we choose four relative values for peridotite mineral saturation (represented
as aH20
0-1, 0.3, 0.5, and 0.7) and calculate the temperature of the corresponding H2 0-
undersaturated solidus and how it varies with pressure. Each of the values of mineral
saturation corresponds to an H20 content determined from the partitioning data from
Hauri et al. [2006], which changes with pressure, as H20 solubility is pressure-dependent
as described in Table 2. For example, in the original melting scenario illustrated by the
red points in Figure 3, the peridotite contains 50 ppm H2 0 at 1 GPa, 150 ppm H2 0 at 3
GPa (the pressure illustrated in Fig. 3) and 300 ppm at 6 GPa, which corresponds to an
aH20 -1.
By choosing to parameterize our model in terms of constant H2 0 activity, we
do not mean to imply that there is a depth gradient in the H20 content of the
asthenosphere. Rather by holding H2 0 activity constant in these expressions, it facilitates
a qualitative assessment of the degree of H20 saturation required to trigger melting at a
given depth in the asthenosphere.
AH 2o =0.1:
T(0 C) = -2.848P
AH 2o =0.3:
T( 0 C)
=
2 +104.312P
+1120.7
1.659P 2 + 47.152P + 1120.7
207
(5)
(6)
AH2o =0.5: T(oC)=6.168P
AH 2O
=0.7: T(*C)
=
2
-10.01P +1120.7
10.677P 2 - 67.173P + 1120.7
(7)
(8)
The water content of melts at the H 20-saturated eutectic can be calculated using
the expression
7290P
-810T
-
24600H 20 + 1093500 =0
(9)
from Grove et al. [2006] where P is pressure in kbars, T is the temperature of the H2 0saturated solidus in 'C and H20 is the water content of the melt in wt%.
2.2 Numerical Models of Mantle Flow and Temperature
To explore the implications of the melting model, we applied it in the context of a
model of possible mantle flow beneath the northeastern U.S.. Two-dimensional
Cartesian numerical experiments on convection were conducted using the finite-element
code FEMcont [Kelemen et al., 2003; Zaranek et al., 2004]. The aim of these
experiments was to produce predictions of asthenospheric flow patterns and temperature
fields near the LAB at the edge of a lithospheric keel (Fig. 4). The model box for the
numerical experiments consists of a mesh of 201 by 101 nodes. The top of the model box
has a no-slip boundary, where v,= vz = 0, and temperature is held fixed at T=0. The
bottom of the box has a fixed temperature and velocity boundary conditions such that v,,
= V and v2 = 0. The vertical sides of the computational domain allow flow-through, such
that dvx/dx = dv2/dz = 0. The model material is an infinite Prandtl number, Boussinesq
fluid. In our Eulerian code, the model asthenosphere moves relative to the fixed
reference frame of the continental lithosphere. Thus absolute plate motion for eastern
North America is represented as a "mantle wind" in the asthenosphere and applied as a
208
constant horizontal velocity boundary condition at the base of the mantle in the model, as
illustrated in Figure 4. The models were run with mantle wind velocities from 0-3
cm/year in a west-to-east direction to represent the range of predictions for the east-west
component of North American plate motion [Conradet al., 2004].
Initial temperatures are defined linearly within the lithosphere and as a mantle
potential temperature within the asthenosphere. The experiments used an LAB profile
that approximates the results of Sp and Ps analyses [Rychert et al., 2005; Rychert et al.,
2007; A bt et al., 2010] and some surface wave tomography studies [van der Lee, 2002; Li
et al., 2003; Nettles and Dziewonski, 2008; Bedle and van der Lee, 2009] across the edge
of the craton in Pennsylvania and the continental margin in New England. The model
includes a 200 km thick continental lithosphere, a 90 km thick lithosphere beneath the
continental margin and a horizontal transition between these two lithospheric thicknesses
of 50 km. The width of this horizontal transition is difficult to precisely constrain, but a
combination of regional surface wave tomography [Li et al., 2003] and Sp and Ps
constraints [Rychert et al., 2005; Rychert et al., 2007; Abt et al., 2010] indicate that
values of 50-100 km are reasonable. Mantle flow patterns at the continental margin were
affected by changing the thickness of the continental lithosphere and the width of the
horizontal transition in lithosphere thickness (Fig. 5), but not continent length.
The model solves non-dimensional equations composed of the conservation of
mass, momentum, and energy. The thermal buoyancy of the model fluid was determined
by a dimensionless Rayleigh number,
Ra=
aATh
Y 0K
209
(10)
where p is fluid density (kg/m 3), g is gravitational acceleration (m/s 2 ), a is thermal
expansivity (K-1), AT is temperature difference across the region in question (K), h is the
height of the model box (in), puo is the reference fluid viscosity (Paes) when T=Tp and K
is thermal diffusivity (m2/s). Below a critical Rayleigh number, conduction is more
efficient than convection. Rayleigh numbers in the numerical experiments were varied
over the range of the following variables, where p=3300 kg/m 3, a=2.4-3*10-5 K,
AT=1573-1673 K, h=410-660 km, y, for the asthenosphere is 1018-21 Paes and K=10-6
m 2 /sec. The lower bound on the reference viscosity of the upper mantle is based on
deformation in response to local loads (i.e., earthquakes and lake filling) in the western
U.S. [e.g., Kaufmann andAmelung, 2000], whereas the higher estimates are from glacial
isostatic adjustment models at cratonic or global scales [Lambeck and Johnston, 1998;
Peltier,2001b; a; Dixon et al., 2004].
Asthenosphere viscosity in the model deviates from the reference viscosity
according to temperature in an Arrhenius relationship,
( Q *( T,
y
)-
*
(11)
where yo is the reference fluid viscosity (Paes), T, is the mantle potential temperature and
Q is activation energy.
Q/RT varied between 30 (a small increase in temperature
resulting in a large decrease in viscosity) and 15 in our experiments by using
experimentally-derived values of Q that vary from -200 kJ/mol for diffusion creep to
-500 kJ/mol for dislocation creep in the mantle [Karato and Wu, 1993]. To create a more
viscous lithosphere, a multiple of the asthenosphere reference viscosity was applied to the
lithosphere; lithospheric viscosity was allowed to deviate from the reference value again
using equation 11. No evidence exists for lithospheric instability beneath eastern North
210
America in the last 100 m.y., therefore a relative lithosphere-asthenosphere viscosity of
1000 was used in the numerical experiments as it produced the most stable lithosphere
geometry [Lenardic and Moresi, 1999]. Lower relative viscosities produced lithospheric
instabilities within 20 million years.
2.3 Melt Calculations:CriteriaforMelt Production
After solving for the temperature and velocity fields using FEMcont, we identify
regions of upwelling and compare the temperature and pressure of each node to the solidi
relevant to the asthenosphere described in section 2.1 to determine if melting could occur.
Particular attention is paid to vertical slices through the model box with significant
convective upwelling at the edge of the thicker continental lithosphere. The mantle is
considered to contain partial melt if it is above the solidus (i.e., temperature of first melt)
in a given time step. No compositional information is built into the FEMcont numerical
code. Composition is only expressed in the solidus chosen for the melting calculations.
We acknowledge that interaction between melting and solid flow are not considered
explicitly here, as the melting calculations are not fed back into viscosity. To incorporate
these effects in the computation requires a significant step in the complexity of the
model, which we have chosen to exclude in our effort to simply demonstrate the
plausibility of mantle melting near the LAB.
The melting calculations were run for a range of mantle potential temperatures
(Tp=1300-1550'C) based on olivine-liquid equilibria and olivine phenocryst
compositions in primary magmas from a range of tectonic settings [Kinzler and Grove,
1992a; Putirka,2005; Herzberg et al., 2007]. Anderson [1982] pioneered the idea that
211
continents can act to insulate the mantle by preventing dissipation of heat from the
Earth's interior, causing broad thermal anomalies, sometimes known as "bottom
heating." Today this idea has been substantiated in modem 2D and 3D mantle
convection models [Gurnis, 1988; Lowman andJarvis, 1996; Lowman and Gable, 1999;
Phillips andBunge, 2005; Korenaga,2007; Phillips and Coltice, 2010]. The possibility
of subcontinental insulation was considered in our models by calculating the maximum
range of additional heat retained below the North American continent [-20-40'C;
Phillipsand Bunge, 2005] and applying this additional heat to mantle that has traveled
under the continent before upwelling.
To calculate melting, we used mantle adiabats ranging from 0.29-0.39 0 /km (1013'/GPa) to reflect different estimates of the coefficient of thermal expansion and the
specific heat of the asthenosphere. Changing the adiabatic gradient had only a small
effect on the location of melting in the asthenosphere; a mantle adiabat of 13'/GPa
located the maximum depth of melting at 4-8 km deeper than an adiabat of 100/GPa for a
model with the same mantle potential temperature and initial conditions. No other effect
on the location and H20 content of melting was observed based on varying the adiabatic
gradient alone.
Once a mantle potential temperature, adiabatic gradient, composition and H2 0
content are assigned, the occurrence of asthenosphere melting in our calculations depends
only on pressure. If the asthenosphere is above the solidus at a given pressure, the
presence of edge-driven convection cells will only increase the quantity of melt produced
over time.
212
3. Results
The calculated asthenosphere flow fields produce upwelling regions viable for
adiabatic decompression melting at the margin of the lithospheric keel. The lateral extent
of upwelling is affected by both the width of the step transition from the keel to thinner
lithosphere and the thickness of the keel itself. While in eastern North America the keel
appears to be ~200 km thick, and the transition to thin lithosphere to occur over -50 km,
both thinner and thicker keels and transition widths will produce upwellings with the
potential for melting. Thicker keels and more abrupt transitions to thinner lithospheres
produce larger convective upwellings, extending almost 200 km across the thin
lithosphere (see Fig. 5). When the transition from the keel to thinner lithosphere is more
gradual (approaching a fluid dynamic shape), the convective eddy is located more under
the transition step from keel to thinner lithosphere, and extends less under the thin
lithosphere. When the mantle wind velocity is 0 cm/year, convective upwelling still
occurs via simple edge-driven convection immediately adjacent to the change in
lithospheric thickness. As the mantle wind velocity is increased, the location of
maximum upwelling migrates upwind (Fig. 7). Thus the magnitude of the mantle wind
also acts to relocate the site of the greatest chance of melting adiabatically. The size of
the melting region predicted by a given solidus also decreases with an increase in the
velocity of the mantle wind.
Our modeling and melting calculations suggest melting of anhydrous peridotite
will not occur in the asthenosphere for mantle potential temperatures less than 1425'C.
The location of melting produced by the 150 ppm H2 0 solidi calculated using the models
of Hirschmannet al. [2009] (Fig. 6e), Katz et al. [2003] (Fig. 6d) and Medard and Grove
213
[2008] (Fig. 6c), as well as the solidus for aH2O=0. I or 10% mineral saturation (Fig. 6b)
are virtually identical due to their overlap in temperature between 2 to 7 GPa (Fig. 2,
Table 3). Our melting calculations with any of these solidi suggest a mantle potential
temperature greater than approximately 1350'C is required to produce melting near the
LAB. The depth of melting in our experiments with these solidi for an asthenosphere
with a potential temperature of 1350'C is between 102-126 km, with each solidus
producing a melting region that is 21 or 24 km in height. Melting in our numerical
models occurs over a narrower range of pressures (102-121 km vs. 99-141 km) than
predicted by the original Hirth and Kohlstedt [1996] solidus at H20 concentrations of
150-200 ppm (Fig. 6a, Table 3).
The location of melting for experiments with the same mantle potential
temperature and initial conditions varies more for the calculated 450 ppm H20 solidi,
relative to those for 150 ppm H20, because of more substantial differences in solidus
temperature between the models at the relevant pressures. The 450 ppm H20 solidi of
Katz et al. [2003] (Fig. 6i) and Medard and Grove [2008] (Fig. 6h) predict melting
between -91-130 km for a mantle potential temperature of 1350'C (also see Table 3).
Our solidus for an aH2O=0. 3 or 30% mineral saturation (Fig. 6g) is closest in temperature
to the calculated 450 ppm H2 0 Hirschmann et al. [2009] solidus (Fig. 6i), and these two
models predict similar melting regions between ~91-200 km for an asthenosphere with a
potential temperature of 1350'C.
Composition (i.e., peridotite vs. garnet pyroxenite) has a small effect on the
temperature of the solidus at low pressures in this model. The anhydrous garnet
214
pyroxenite solidus [Kogiso et al., 2003] triggers a small region of melting just below the
LAB at mantle potential temperatures >1375 0 C.
4. Discussion
4.1 The LAB beneath eastern North America
As is clear from the temperature structure of the flow models presented here,
vertical gradients in temperature alone are too gradual to produce an LAB consistent with
seismic constraints. In our models, temperature gradients from the lithosphere to the
asthenosphere are distributed over more than 60 km in depth, and the maximum
temperature gradient does not exceed 14.5 0 C/km. In contrast, a temperature gradient of
more than 20'C/km over less than 11 km is required to produce the combination of Sp
and Ps phases observed at the LAB beneath stations in the northeastern U.S. and
southeastern Canada [Rychert et al., 2007]. Sp phases from the LAB observed more
widely in the eastern U.S. [Abt et al., 2010] indicate a velocity gradient that is localized
over less than 40 km in depth, and even those looser constraints are not met by an LAB
that is created purely by temperature. This conclusion is consistent with earlier
comparisons to the work of models by King and Ritsema [2000], Lenardic andMoresi
[1999], Korenaga and Jordan [2002], and Cooper et al. [2004].
Small degrees of asthenosphere melting may be a viable mechanism to explain the
observed drop in shear wave speeds across the LAB in eastern North America. The
temperature and H2 0 contents of the upwelling asthenosphere at the edge of the
continental lithosphere are the main factors that determine if melting will occur in our
numerical experiments. These asthenosphere source conditions are dependent on the
215
direction and pattern of the mantle wind (equivalent to plate motion direction) under
North America. Here we examine a range of possible source conditions because of the
uncertainty of the pattern of a North American mantle wind.
Asthenospheric mantle material saturated with H20 is most likely to melt since its
solidus is at the lowest temperature for a given pressure. However it is unlikely the
>500-3000 ppm H2 0 required to produce melt at H20-saturated conditions is present in
the asthenosphere, except perhaps above a downgoing oceanic slab at a subduction zone.
However, moderate amounts of H2 0 are likely in the asthenosphere, perhaps
heterogeneously distributed throughout [Bell and Rossman, 1992]. Estimates of H20
content for a normal MORB-source upper mantle converge on a value of 100± 50 ppm
[Dixon et al., 1988; Michael, 1988; Hirth andKohlstedt, 1996; Saal et al., 2002; Salters
andStracke, 2005; Workman and Hart,2005] and 300-900 ppm H20 for enriched
MORB- and OJB-source upper mantle [Jambon and Zimmermann, 1990; Sobolev and
Chaussidon, 1996; Dixon et al., 1997; Aubaud et al., 2005; 2006]. One plausible
scenario to produce melting in our numerical experiments is a mantle potential
temperature of ~135 0 'C (or a potential temperature of 1310-1330'C combined with 2040'C of subcontinental insulation) for an asthenosphere with -15 0 ppm H2 0. If the
mantle is significantly colder or contains significantly less H20, no melting will occur. If
there are small regions of higher H 20 content (e.g., ~450 ppm H2 0) within the convecting
asthenosphere, for example asthenosphere that was hydrated during the subduction of the
Farallon slab or another subduction event, melting is possible at mantle potential
temperatures as low as 11 90'C for our parameterization at 30% mineral saturation,
1200'C for the Hirschmann et al. [2009] 450 ppm H20 solidus, and 1215'C for the 450
216
ppm H2 0 Katz et al. [2003] solidus. Likewise, CO 2 combined with H 20 will increase the
likelihood of small degrees of partial melting in the MORB-source upper mantle
[Dasguptaet al., 2007]. Alternatively, if there is garnet pyroxenite in the convecting
asthenosphere (e.g., remnant pieces of subducted slab), small regions of melting occur at
mantle potential temperatures >1375'C. Mantle potential temperatures above 1420'C
(equivalent to what we estimate for plume source regions) are required to melt an
anhydrous asthenosphere source with no H 20 in the nominally anhydrous minerals.
How much melt would be produced in the numerical models of this study? A
simplified expression for melt production (J)per degree of temperature above the mantle
solidus is:
(T -T
df
dT
(12)
where df is the fraction of melt (wt%) produced over dT, an increment of temperature,
and T-Tsolidus is the temperature interval above the solidus. We use estimates for df/dT for
H20-undersaturated peridotite melting at 1.2 GPa by Gaetaniand Grove [1998], which
yield melt productivity rates of 0.03 wt%/1 C near the solidus to 0.173 wt%/1 C at
increased extents of melting, which are much lower than experimentally-determined melt
productivity rates from anhydrous peridotite melting of 0.26 wt%/1 C at 1.5 GPa
[Falloon andDanyushevsky, 2000] or 0.45 wt%/10 C at 3 GPa [Walter, 1998]. Using this
highly simplified expression for melt fraction, we calculate the range of in situ melt
fractions for our most reasonable case (mantle temperature=1350'C, 150 ppm H2 0) to be
<0.1 - 2.8 wt% (or 0.0 1-3.3 vol% assuming a hydrous silicate melt density of 2.7g/cc) for
an adiabatic gradient of 13'/GPa.
217
Once generated, a multitude of potential fates await these melt fractions. Any
initial melt fraction will grow at a rate proportional to its upwelling rate. The
permeability and porosity of the host mantle will depend on deformation. Lab
experiments on stress-driven melt segregation demonstrate as little as 0.5 vol% MORB
melt can segregate from an olivine host and 2-3 vol% MORB melt can overcome surface
tension at mantle grain sizes to form networks [Holtzman and Kohlstedt, 2007].
Likewise, work by Takei [2002; 2005] indicates that deformation can reduce grain
contact area, create high porosity melt sheets with a strong preferred orientation, and
dramatically reduce shear velocity. An in-depth analysis of the distribution of melt,
accounting for melt buoyancy and deformation driven melt migration, is beyond the
scope of this paper. However, in the simplest scenario, for these melt fractions to
adequately explain the sharp velocity gradient at the LAB below eastern North America,
a significant percent of the melt must remain trapped in the uppermost asthenosphere. If
the small amount of melt predicted by our model for a potential temperature of 1350'C
and 150 ppm H2 0 does indeed segregate, the work of Sparks and Parmentier[1991],
Katz et al. [2006; 2008], and Holtzman and Kohlstedt [2007] suggests this melt will form
networks in the uppermost asthenosphere adjacent to the LAB due to some combination
of dilation of the porous matrix and the high shear strains located here from corner flow
and the bordering rheological transition. Therefore, Kohlstedt and Holtzman [2009] state
that "based on the simple point that stress-driven segregation and organization will be
most effective where strain rates are highest, we hypothesize that the LAB is marked by
the location of melt-enriched shear zones that lubricate the LAB and thus the plate
boundary system." This configuration of melt would be ideal to produce the observed
218
shear wave velocity profiles across the LAB. These melt networks would quickly freeze
if the melt penetrates the cooler overlying lithosphere [Rubin, 1995]. As such we would
not expect any surface expression of these melts.
Given that the melt fractions produced in situ likely represent a lower bound on
the actual melt fraction just below the solidus-defined LAB, how do they compare to the
seismic observations? Rychert et al. [2007] constrained the seismic shear wave velocity
drops beneath two seismic stations in eastern North America to be 5-7% (HRV) and 610% (LMN), and similar velocity drops are consistent with the Sp phases observed more
widely in eastern North America by A bt et al. [2010]. According to Hammond and
Humphreys [2000] melt fractions of 1-2% are capable of producing velocity reductions
within these ranges, assuming realistic melt distribution at grain boundaries. Following
Takei [2002], Takei and Holtzman [2009ab] and Kawakatsu et al [2009], for texturally
equilibrated melt at grain boundaries, melt fractions required to produce the observed
velocity reductions would exceed 3%, but if melt occurs in horizontal melt-rich bands,
significantly smaller average mantle melt fractions would be sufficient. Therefore,
although predicted velocity drops strongly depend on deformation and the inferred
geometry of melt distribution, the in situ melt fractions at all but the lowest end of the
range extrapolated from the models in this study appear to be capable of matching the
velocity drops observed at the LAB in eastern North America. In addition, if melt
produced deeper in the super-solidus zones collects at depths just below the solidusdefined LAB, even larger seismic velocity reductions would be predicted.
4.2 The Global LAB
219
The models presented here show that small degrees of asthenospheric melting are
a possible mechanism for the sharp vertical velocity gradients observed in eastern North
America, but they also imply that asthenospheric melting may play a role in defining the
LAB in other regions where similar ranges of temperature, composition and H20 content
apply. Several recent studies have found evidence for a sharp seismic discontinuity at
depths interpretable as the LAB beneath relatively thin Phanerozoic continental
lithosphere, but not beneath thick cratonic lithosphere, on a global basis [Rychert and
Shearer, 2009], and beneath North America [A bt et al., 2010] and Australia [Ford et al.,
2010]. Our work suggests that this observation may be explained by favorable conditions
for asthenospheric upwelling and melting in regions of thin lithosphere and the
unfavorable conditions for asthenospheric upwelling and melting below thick cratonic
lithosphere. While the modeling results in this study show that melting produced by
upwelling is a viable explanation for sharp LAB velocity gradients in regions of
relatively thin continental lithosphere downwind of lithospheric keels, in some cases at
distances of up to almost 200 km from the edge of the keel, comparably sharp LAB
velocity gradients have been observed beneath continents outside this zone of influence
[e.g. Li et al., 2007; Heit et al., 2007; Chen et al., 2008; Rychert and Shearer,2009; Abt
et al., 2010; Fordet al., 2010]. In some of these regions, different forms of upwelling are
possible, but other mechanisms to produce low degrees of melt also exist. Other than
upwelling, melting can be produced by raising temperature, possibly by gradual
asthenosphereic heating from radiogenic elements while advecting beneath a thick
insulating lithosphere (i.e., bottom heating). The degree of heating from this process is
limited and would only operate under specific circumstances, when a long
220
sublithospheric traverse occurred and asthenospheric material was already near its
solidus. Melting can also be triggered by injection of incompatible elements such as
water, perhaps through dewatering of small parcels of downwelling enriched lithosphere
melting in small-scale convection at the LAB. Alternatively, the possibility that sharp
LAB velocity gradients reflect dry, depleted lithosphere over hydrated, fertile
asthenosphere cannot be ruled out.
Because the melting temperature of the asthenosphere is highly dependent on the
asthenosphere temperature and H20 content, the conditions required for melting below
the LAB may be met at some locations and times but not others. Therefore, as
temperature or H20 content fluctuate within the region of upwelling asthenosphere, the
fraction of melt produced may increase or decrease with time. Thus the magnitude of the
shear wave speed drop at the LAB produced by small degrees of partial melt in the
asthenosphere may change with time and space. For example, the amplitudes of the Sp
phases interpreted as originating from the LAB in eastern North America vary by a factor
of two, and, as a whole, LAB Sp phases observed in the tectonically active western U.S.
are significantly larger than their eastern counterparts [Abt et al., 2010]. In addition,
LAB Sp phases in Phanerozoic eastern Australia show a marked correlation between
shallower depths, larger amplitudes, and the locations of most recent, voluminous
magmatism [Fordet al., 2010].
4.3 Comparison to Other Melting Models
The location and extent of melting in our numerical experiments for an upper
mantle with 150 ppm H2 0 are nearly identical when the expressions for solidus
221
depression from Hirschmann et al. [2009], Katz et al. [2003], the expression for liquidus
depression from Medardand Grove [2008], or our own simplified parameterization for
solidus temperature at a given degree of H20 mineral saturation, are used (Table 2, Fig. 2
& 6).
The expression for solidus depression from Hirschmann et al. [2009] predicts that
a peridotite with 100 ppm H20 begins to melt 80 km along an adiabat with a potential
temperature of 1323 C, or 15 km deeper than the case of a truly dry peridotite. Despite
the differences in the formulation of our simplified parameterization, which uses a linear
approximation for the nominally anhydrous surface, this result is almost identical to our
predictions for a peridotite with 10% mineral saturation (aH2O=0.1), which melts at 73 km
along an identical adiabat, or 13 km deeper than the case of a truly dry peridotite using
the Hirschmann [2000] anhydrous solidus. In another example, using our solidus
parameterization and a potential temperature of 1350'C, a peridotite with 10% mineral
saturation melts at 84 km, or 18 km deeper than the truly anhydrous case. Likewise, when
the expression for solidus depression of Katz et al. [2003] is used with updated bulk
partition coefficients for H20 between peridotite and melt, rather than the constant bulk
partition coefficient used in their original paper, it is remarkably similar to both the
Hirschmannet al. [2009] and 10% mineral saturation expression presented here.
At higher H20 contents (e.g., 450 ppm H20) differences between the various
expressions for solidus depression with H20 increase. The Medard and Grove [2008]
expression is derived from experiments on basaltic liquidus depression and was not
originally intended to apply to solidus depression, as the expressions for Katz et al.
[2003] and Hirschmann et al. [2009] are. The predictions of our parameterization for
222
solidus temperature at 30% mineral saturation varies from the Hirschmann et al. [2009]
solidus for 450 ppm bulk H2 0 by only 2-14'C between 2-7 GPa.
4.4 Implications of the similar expressionsfor solidus and liquidus depression with H2O
The power-law expressions for solidus depression from the cryoscopic approach
of Hirschmannet al. [2009] and the mathematical model of Katz et al. [2003], although
calibrated on high-pressure experiments, do not fully account for the speciation of water
in silicate melts or the non-ideal mixing between the anhydrous and hydrous components.
Although the activity of water in silicate melt can be approximated as these power-law
expressions [Burnham, 1975; Stolper, 1982], the non-linearity of liquidus (or solidus)
depression cannot be explained by an ideal solution model for the speciation of OH and
H20 molecules in silicate liquids because it is likely related to the interaction between
H20 and other melt components [Nicholls, 1980; Medard and Grove, 2008].
Medard and Grove [2008] are able to use Margules parameters in their
thermodynamic model for liquidus depression after the method of Nicholls [1980] to
account for the non-ideality of H20 in silicate melt. Although we recognize both the
linear and cubic parameterization of liquidus depression by Medardand Grove [2008]
were never intended to apply to solidus depression, the similarity between the quantity of
liquidus depression calculated from their model and the quantity of solidus depression
calculated from models that do not account for non-ideality, is intriguing. It appears the
effect of small amounts of H20 (<200 ppm bulk H2 0) on solidus depression, as
calculated using the various expressions here, may not vary significantly from its effect
on liquidus depression. Alternatively, the amount of solidus depression calculated from a
223
yet to be determined non-ideal solution model may be dramatically different from the
expressions used here and from the amount of liquidus depression.
In the absence of sufficient high precision experimental data and a derivative
model that accounts for the non-ideality of solidus depression, there is a qualitative
agreement between the power-law expressions for solidus depression and our linear
approximation of solidus depression at low H20 contents and upper mantle pressures (< 7
GPa) when using the bulk H20-peridotite partitioning data now available.
The advantage of our parameterization is that it can be applied to any scenario
where there is a designated temperature and H20 content of mantle peridotite to test the
plausibility of mantle melting at pressures < 7 GPa, or to determine an approximate H2 0
content required for melting at a given temperature and pressure. But because this model
does not take into account phase equilibria, employ a more sophisticated activitycomposition model or account for reductions in temperature due to the heat of fusion, it is
not intended to replace a complete treatment of melting conditions or to determine the
extent of melting at a given P and T, except as a maximum estimate given by the simple
approximation in equation 12.
5. Conclusions
Melting of the asthenosphere at the edge of the continental keel may be
responsible for shear wave speed gradients observed at the LAB in eastern North
America. Numerical experiments and melting calculations predict <0.1-2.8 wt% (0.013.3 vol%) melting at a depth of 102-126 km at the edge of the continental keel for an
asthenosphere with average mantle potential temperatures and H20 contents equivalent to
224
the average MORB-source. This melting is sufficient to generate the 5-10% drop in
shear wave speed observed at the LAB below eastern North America. Such melt could in
part explain the rheological contrast observed at the LAB without producing a surface
signature.
Production of low extent melts via convective upwelling in the shallow
asthenosphere is a viable process in a variety of locations in addition to the one examined
here. Melting is likely to vary in time and space with H20 content, composition and
temperature of the asthenosphere. For example, lenses within the mantle of remnant
subducted slabs or hydrated mantle wedge will have lower melting temperatures than the
nominally anhydrous mantle. Similarly, geologic time periods with a higher mantle
potential temperature relative to today are more likely to produce melt fractions near the
LAB. Thus if small degrees of asthenospheric melting are a primary cause of the LAB,
the magnitude of shear wave gradients across this boundary are likely to be variable in
space and time.
A comparison of different expressions for peridotite solidus depression in the
presence of small amounts of H2 0 all produce very similar predictions for the location
and extent of melting, including the simplified parameterization at constant mineral
saturation presented here. Until more sophisticated models of peridotite solidus
depression that use a non-ideal solution model for H20 in silicate melts are developed,
our simple parameterization is a good starting point to test the feasibility of melting in the
mantle anywhere P, T and H 20 content are specified.
225
Acknowledgements
This research was supported by NSF EAR 0538155 and a NSF Graduate Research
Fellowship. We thank Sarah Zaranek for help with the FEMcont code, Timothy Grove
for guidance with the solidus parameterization and four anonymous reviewers for their
thoughtful comments.
Figure Captions
Figure 1. Anhydrous and H20-saturated Peridotite Solidi Used in Our Melting
Calculations. Blue bold line: fourth-order parameterization of the H20-saturated solidus
based on the illustrated experiments by Grove et al. [2006] and Till et al. [2007] (light
blue circles) and Mysen and Boettcher [1975] (dark blue circles). Closed circles
represent super-solidus experiments, and open circles, sub-solidus experiments. Black
bold line: second-order parameterization of the H20-saturated solidus for extrapolation
between 6-12 GPa. Red bold line: anhydrous peridotite solidus parameterization from
Hirschmann [2000], which is based on the compilation of anhydrous peridotite melting
experiments shown by the black circles. Also shown are anhydrous peridotite melting
experiments from Kinzler and Grove [1992b] and Grove and Juster [1989] in gray. Green
bold line and circles: garnet pyroxenite melting experiments and solidus from Kogiso et
al. [2003]. Thin gray line: mantle adiabat of 13'/GPa with a mantle potential temperature
of 1350'C.
Figure 2. H20-undersaturated Peridotite Solidi Used in Our Melting Calculations. The
H20-undersaturated peridotite solidi calculated with the expressions of Hirschmannet al.
226
[2009], Katz et al. [2003], and Medardand Grove [2008] as described in section 2.1 for
150 and 450 pm H2 0 are illustrated by the purple and blue bold lines, respectively. The
bump visible in these solidi at pressures near 2.5 GPa is related to the change in partition
cofficient for H between bulk peridotite and liquid immediately preceding and following
the onset of garnet stability. The H20-undersaturated peridotite solidi calculated at a
constant degree of saturation of the nominally anhydrous minerals from expression
presented here and by Hirth andKohlstedt [1996] are illustrated by the orange and red
dashed lines, respectively. Black bold curves denote the anhydrous and H20-saturated
peridotite solidi used in our melting calculations and the thin gray line represents a
mantle adiabat of 13 /GPa with a mantle potential temperature of 1350'C. The solidi for
10% mineral saturation (aH2O0. 1) overlap with the solidi calculated for 150 ppm H 20 for
pressures between 2 and 7 GPa and those for 30% (aH2o=0. 3 ) with the 450 ppm H2 0
solidi over the same pressure range.
Figure 3. Temperature-Water Content Phase Diagram for Peridotite Melting at Constant
Pressure after Grove et al. [2006]. Temperature and H20 contents are illustrated for 3
GPa. The green curve illustrates the maximum storage capacity of nominally anhydrous
peridotite minerals (NAM) based on partitioning data from Hauriet al. [2006] and the
blue line the simplified melting boundary for peridotite from a eutectic-like H2 0saturated solidus to an anhydrous liquidus, after Silver and Stolper [1985]. A peridotite
containing 150 ppm H2 0, illustrated by the red points, will melt at 1409'C where the
H2 0 content intersects the NAM curve. This H2 0 content corresponds to an activity of
H2 0 for this peridotite of-0. 1 (as calculated in section 2.1). At 3 GPa the H20-saturated
227
peridotite solidus is from Grove et al. [2006] and Till et al. [2007], the anhydrous solidus
Hirschmann [2000] and the liquidus from Zhang and Herzberg [1994]. Note, in this
phase diagram, the liquidus and solidus curves do not show phases that are known to be
in equilibrium except at the H20-saturated solidus.
Figure 4. FEMcont Convection Model Setup. Dimensions of the lithospheric step were
taken directly from seismic studies of eastern North America [Li et al., 2003; Rychert et
al., 2007; Abt et al., 2010]. [t. refers to the reference viscosity of the asthenosphere or
lithosphere.
Figure 5. Sensitivity of Finite Element Modeling of Asthenospheric Flow to the Depth of
the Lithospheric Keel and the Width of the Step to Thinner Marginal Lithosphere. While
in eastern North America the keel appears to be ~200 km thick, and the transition to thin
lithosphere to occur over -50 km, both thinner and thicker keels and transition widths
will produce upwellings with the potential for melting. Thicker keels and more abrupt
transitions to thinner lithospheres produce larger convective upwellings, extending almost
200 km across the thin lithosphere. When the transition from the keel to thinner
lithosphere is more gradual (approaching a fluid dynamic shape), the convective eddy is
located more under the transition step from keel to thinner lithosphere, and extends less
under the thin lithosphere.
Figure 6. Locations of Melting near the LAB Predicted by Our Numerical Experiments
and Melting Calculations. The color contours denote the temperature of the lithosphere
228
and asthenosphere in degrees C. Supersolidus regions in the asthenosphere lie within a
white contour line at which the asthenosphere temperature equals that of a given solidus.
These calculations are shown after 13 million years and use a plate velocity of 3 cm/year,
a mantle potential temperature of 1350'C, an adiabatic gradient of 13'C/GPa, a Rayleigh
number of 9.4x10 6 (pO=101 9 Paes) and a relative lithosphere-asthenosphere viscosity of
1000. a) Melting region predicted using the solidus of Hirth andKohlstedt [1996] for
nominally anhydrous mantle minerals (NAM) that are 10% saturated with H20. b)
Melting region predicted using our solidus for 10% H2 0 saturation of the NAM
(aH200. 1)- c) Melting region predicted using the solidus for a peridotite with 150 ppm
H2 0 calculated from the expression for liquidus depression of olivine-saturated basaltic
melts in the presence of H2 0 from Medardand Grove [2008] as described in section 2.1.
d) Melting region predicted using the solidus for a peridotite with 150 ppm H2 0
calculated from the expression for solidus depression of Katz et al. [2003]. e) Melting
region predicted using the solidus for a peridotite with 150 ppm H 20 calculated from the
expression for solidus depression of Hirschmann et al. [2009]. f) Calculated with the
solidus of Hirth andKohistedt [1996] for NAM that are 30% saturated with H20. g)
Calculated with our solidus for 30% H2 0 saturation of the NAM (aH2O0- 3 ). h)
Calculated for a peridotite with 450 ppm H20 using Medardand Grove [2008]. i)
Calculated for a peridotite with 450 ppm H20 using Katz et aL [2003].
a peridotite with 450 ppm H2 0 using Hirschmannet al. [2009].
229
j)
Calculated for
Figure 7. Comparison of Model Runs With and Without the Mantle Wind. A) Same
model run as illustrated in Figure 5, panel b). b) Model with identical initial conditions
except with a mantle wind velocity of 0 cm/year.
Table 1. Equation for the H20 storage capacity of the nominally anhydrous peridotite
minerals at different pressures. Y-intercept is the anhydrous solidus for a given pressure.
Temperature in 'C, pressure in GPa.
Table 2. Predicted H20 contents of mantle peridotite with pressure for different degrees
of H20 saturation of the nominally anhydrous minerals based on the partitioning data of
Hauri et al. (2006). All H20 in ppm, P in GPa.
Table 3. Location and extent of melting for the vertical slice across the melting region in
our numerical models that includes the node with the maximum temperature above a
given solidus. The location and extent of melting is also given for the node with the
minimum temperature above the solidus. All values are calculated from model runs with
a mantle potential temperature of 1350'C. Maximum extent of melting calculated using a
dfldT=O. 173 (wt%/0 C) and minimum extent of melting calculated for df/dT=0.03 as
discussed in section 4.1. Horizontal distances are measured from the right margin (see
Fig. 6).
230
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2000
1900
1800
1700
1600
1500
1400
1300
1200
1100
1000
900
800
700
0
1
2
3
4
P (GPa)
Figure 1.
240
5
6
7
H20-Undersaturated Peridotite Solidi
1900
1800
150 ppm H20 Solidi
1700
Medard & Grove (2008)
Katz et al. (2003)
1600
-iabat Hirschmann et al. (2009)
1500
u
*
450 ppm H2O Solidi
1400
1400--M Medard and Grove (2008)
1300
E
Katz et al. (2003)
Hirschmann et al. (2009)
"
1200
1100
Constant Degree of H20 saturation
in NAM
Hirth and Kohlstedt (1996)
aH20=0.1 and aH20 =0.3
1000
900
-
H 0-Saturated
800
1
2
3
4
5
This paper
aH2 o=0.1 and a12o=0.
700
0
-
6
7
P (GPa)
Figure 2.
241
8
9
10
11
12
3
100
XH20 (wt %)
Figure 3.
242
-
200
300
mantle
wind /
(0-3 cm/yr)0
200
100
Figure 4
243
300
400
150 kmn keel depth
e
E 120
-h
E
200 km
250 km
e
e
100
0
c
0
80
-
t
4/,
60
-
0
c.
A 40
-
0
~20
80
100
I11
I
I
i
30
120
140
160
Distance from step to furthest edge of upwelling [km]
Figure 5.
244
180
200
Figure 6.
1500
5o
100
1000
ISO15
200
250
300
350
400
0
Distance
(km)
0
1i00
50
100
1000
IS0
200
250
S35
300
350
40D
0
0
1500
50
100
1000
I50
200
250
S00
300
350
400
100
100
1000
150
200
250
300
350
400
0
50
100
200
250
245
Figure 7
0
1500
50
100000
9OO
150
200
-C
(D
250
(
300
350
400
300
200
100
Distance (km)
0
1o
50
100
3
1502
200
250
Soo
300
350
40400
300
200
100
246
0
0
Table 1. Equation for the H20 storage capacity of the
nominally anhydrous peridotite minerals at different
pressures. Y-intercept is the anhydrous solidus for a
given pressure. Temperature in *C, pressure in GPa.
Pressure
1
T = -6080 P + 1248.5
2
3
4
5
6
T=
T=
T=
T=
T=
-490 P+1366
-4322.7 P + 1473.4
-3808 P + 1570.6
-3390 P + 1657.6
-3081.3 P + 1734.4
247
Table 2. Predicted H2 0 contents of mantle peridotite with pressure for different degrees
of H2 0 saturation of nominally anhydrous minerals based on the parititioning data of
Hauri et al. (2006). All H20 in ppm, P in Gpa.
Pressure
aH2O=1.0*
aH2O=0. 7
aH2O=0.5
aH2O=0. 3
aH20o0-1
1
500
350
250
150
50
2
1000
700
500
300
100
3
1500
1050
750
450
150
4
2000
1400
1000
600
200
5
2500
1750
1250
750
250
6
3000
2100
1500
900
300
* equivalent to XH20 at NAM saturation
248
Table 3. Location and extent of melting for the vertical slice across the melting region in our numerical models that includes
the node with maximum temperature above a given solidus. The location and extent of melting is also given for the node
with the minimum temperature above the solidus. All values are calculated from models with a mantle potential temperature
of 1350*C. Maximum extent of melting calculated using a df/dT=0.173 (wt%/*C) and minimum extent of melting calculated
for df/dT=0.03 as discussed in section 4.1. Horizontal distances are measured from the right margin (see Figure 5).
150 ppm H20
Katz et al. 2003
Depth (kin)
Horizontal Location (km) T above solidus (*C) f (wt%)
102
184
1.8
114
184
14.51 (max)
2.51
126
184
4.3
122
216
0.12 (min)
0.0036
Hirschmann et a!., 2009
102
184
3.1
114
184
15.96 (max)
2.75
126
184
5.9
102
139
.08 (min)
0.0024
Medard and Grove, 2008
105
184
7.2
114
184
12.71 (max)
2.2
126
184
2.6
121
143
0.05 (min)
0.0015
105
114
126
184
184
184
0.74
8.38 (max)
0.79
1.4
0.001
Constant ActivityaH20=0.1
This Paper
121
200
0.03 (min)
99
117
141
98
188
188
188
76
8.75
40.25 (max)
7.59
0.14 (min)
Depth (km)
91
114
181
118
Horizontal Location (km)
184
184
184
253
T above solidus (*C) f (wt%)
74.23
131.43 (max)
22.73
5.13
0.19 (min)
0.0057
91
114
197
197
184
184
184
139
92.7
151.18 (max)
0.13
0.04 (min)
91
114
161
130
184
184
184
143
47.24
100.52 (max)
6.98
0.09 (min)
91
122
205
188
188
188
265
77.97
163.44 (max)
2.47
0.16 (min)
184
184
184
76
81.85
164.88 (max)
5.4
0.04 (min)
Hirth and Kohlstedt, 1996
450 ppm H20
Katz et at. 2003
6.96
0.0042
Hirschmann et at., 2009
26.16
0.0012
Medard and Grove, 2008
Constant ActivityaH20=0.
This Paper
3
17.39
0.0028
28.2
0.0048
Hirth and Kohistedt, 1996
91
118
205
185
249
28.52
0.0012